 

 

 

 

mm;

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 
 

 

The Cenozoic Evolution of the
San Joaquin Valley, California

By J. ALAN BARTOW

 

U.S. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1501

 

 

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON : 1991

DEPARTMENT OF THE INTERIOR

MANUEL LUJAN, 1m, Secretary

U.S. GEOLOGICAL SURVEY

Dallas L. Peck, Director

Any use of trade, product, or ﬁrm names in this
publication is for descriptive purposes only and
does not imply endorsement by the U.S. Government

 

Library of Congress Cataloging-in-Publication Data

Bartow, J. Alan.
The Cenozoic evolution of the San Joaquin Valley, California / by J. Alan Bartow.
p. cm.—(U.S. Geological Survey professional paper ; 1501)
Includes bibliographical references.
Supt. of Docs. no.: I 19.161501
1. Geology, Stratigraphic—Cenozoic. 2. Geology—California—San Joaquin River Valley. 3. Sedimentation and depo—
sition—California-San Joaquin River Valley. I. Title. II. Series.
QE690.B34 1991 551.7’8’097948—dc20 89—600313
CIP

 

For sale by the Books and Open-File Reports Section, U.S. Geological Survey,
Federal Center, Box 25425, Denver, CO 80225

CONTENTS

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

Page Page
Abstract 1 The sedimentary record—Continued
Introduction 2 Paleogene—Continued
Setting 2 Upper Paleocene and lower Eocene sequence ———————— 17
Previous work 2 Eocene sequence 18
Purpose and approach 3 Lower Oligocene sequence ---------------------- 19
Geology 3 Upper Oligocene sequence —————————————————————— 19
Northern Sierran block 3 Neogene and Quaternary 20
Southern Sierran block 6 Lower and middle Miocene sequence ——————————————— 20
Northern Diablo homocline 8 Middle and upper Miocene sequence ——————————————— 21
West-side fold belt 9 Upper Miocene, Pliocene, and Pleistocene sequence———- 23
Maricopa-Tejon subbasin and south-margin deformed Upper Pleistocene and Holocene deposits ——————————— 23
belt 10 Basin evolution 24
Major controls on sedimentation ———————————————————————— 12 Paleocene 25
Tectonics 12 Eocene 25
Plate tectonics 12 Oligocene 28
Regional tectonics 13 Miocene 29
Sea-level change 15 Pliocene 31
Climate 16 Pleistocene 31
The sedimentary record 16 Holocene 33
Paleogene 16 Conclusions 33
Great Valley sequence 16 References cited 34
ILLUSTRATIONS
[Plates are in pocket]
PLATE 1. Generalized geologic map and cross sections of the San Joaquin Valley area, California.
2. Correlation of the Cenozoic stratigraphic units of the San Joaquin basin with tectonic, volcanic, and sea-level events.
Page
FIGURE 1. Index map of central California 4
2. Diagram showing interrelation of factors affecting eustasy 16
3. Diagrammatic cross section of the northern San Joaquin Valley showing stratigraphic relations of Quaternary
alluvial deposits 24
4. Maps showing evolution of regional stress patterns in the San Joaquin Valley area 26
5—13. Maps of the San Joaquin basin showing:
5. Late Paleocene paleogeography 27
6. Early to middle Eocene paleogeography 28
7. Middle Eocene paleogeography 28
8. Oligocene paleogeography 29
9. Early Miocene paleogeography 30
10. Middle Miocene paleogeography 30
11. Late Miocene paleogeography 31
12. Pliocene paleogeography 32
13. Pleistocene paleogeography (San Joaquin Valley) 32

 

THE CENOZOIC EVOLUTION OF THE
SAN JOAQUIN VALLEY, CALIFORNIA

By j. ALAN BARTow

ABSTRACT

The San Joaquin Valley, which is the southern part of the
700-km-long Great Valley of California, is an asymmetric structural
trough that is ﬁlled with a prism of upper Mesozoic and Cenozoic
sediments up to 9 km thick; these sediments rest on crystalline
basement rocks of the southwestward—tilted Sierran block. The San
Joaquin sedimentary basin is separated from the Sacramento basin to
the north by the buried Stockton arch and associated Stockton fault.
The buried Bakersﬁeld arch near the south end of the valley separates
the small Maricopa-Tejon subbasin at the south end of the San Joaquin
basin from the remainder of the basin. Cenozoic strata in the San
Joaquin basin thicken southeastward from about 800 m in the north to
over 9,000 m in the south.

The San Joaquin Valley can be subdivided into ﬁve regions on the
basis of differing structural style. They are the northern Sierran block,
the southern Sierran block, the northern Diablo homocline, the west-
side fold belt, and the combined Maricopa-Tejon subbasin and south-
margin deformed belt. Considerable facies variation existed within the
sedimentary basin, particularly in the Neogene when a thick section of
marine sediment accumulated in the southern part of the basin, while a
relatively thin and entirely nonmarine section was deposited in the
northern part. The northern Sierran block, the stable east limb of the
valley syncline between the Stockton fault and the San Joaquin River,
is the least deformed region of the valley. Deformation consists mostly
of a southwest tilt and only minor late Cenozoic normal faulting. The
southern Sierran block, the stable east limb of the valley syncline
between the San Joaquin River and the Bakersﬁeld arch, is similar in
style to the northern part of the block, but it has a higher degree of
deformation. Miocene or older normal faults trend mostly north to
northwest and have a net down-to-the-west displacement with individ-
ual offsets of as much as 600 m. The northern Diablo homocline, the
western limb of the valley syncline between the Stockton arch and
Panache Creek, consists of a locally faulted homocline with northeast
dips. Deformation is mostly late Cenozoic, is complex in its history, and
has included up-to-the—southwest reverse faulting. The west-side fold
belt, the southwestern part of the valley syncline between Panoche
Creek and Elk Hills and including the southern Diablo and Temblor
Ranges, is characterized by a series of folds and faults trending slightly
oblique to the San Andreas fault. Paleogene folding took place in the
northern part of the belt; however, most folding took place in Neogene
time, during which the intensity of deformation increased southeast-
ward along the belt and southwestward toward the San Andreas fault.
The Maricopa-Tejon subbasin and the south-margin deformed belt are
structurally distinct, but genetically related, regions bounded by the
Bakersﬁeld arch on the north, the San Emigdio Mountains on the south,
the Tehachapi Mountains on the east, and the southeast end of the fold
belt on the west. This combined region, which is the most deformed part

Manuscript approved for publication, November 16, 1988.

 

of the basin, has undergone signiﬁcant late Cenozoic shortening
through north-directed thrust faulting at the south margin, as well as
extreme Neogene basin subsidence north of the thrust belt.

The sedimentary history of the San Joaquin basin, recorded in
terms of unconformity-bounded depositional sequences, has been con-
trolled principally by tectonism, but it has also been controlled by
eustatic sea-level changes and, to a lesser degree, by climate. Plate
tectonic events that had an inﬂuence on the basin include (1) subduction
during the early Tertiary that changed from oblique to normal
convergence in the later part of the Eocene, (2) the mid-Oligocene
encounter of the Pacific-Farallon spreading ridge with the trench, and
the consequent establishment of the San Andreas transform, (3) the
northwestward migration of the Mendocino triple junction that induced
extensional tectonism and volcanism in adjacent areas, and (4) the
change in plate motions at 5 Ma that resulted in an increased component
of compression normal to the San Andreas transform. Other tectonic
events of a more regional scale that affected the San Joaquin basin
include (1) clockwise rotation of the southernmost Sierra Nevada and
tectonism that produced large en echelon folds in the southern Diablo
Range, both perhaps related to Late Cretaceous and early Tertiary
right slip on the proto-San Andreas fault, (2) uplift of the Stockton arch
in the early Tertiary, (3) regional uplift of southern California in the
Oligocene that was a precursor to the ridge-trench encounter, (4)
extensional tectonism in the Basin and Range province, particularly in
the Miocene, (5) wrench tectonism adjacent to the San Andreas fault in
the Neogene, (6) northeastward emplacement of a wedge of the
Franciscan Complex at the west side of the Sierran block and the
associated deep-seated thrusting in the Cenozoic, and (7) the acceler-
ated uplift of the Sierra Nevada beginning in the late Miocene.

The early Cenozoic sedimentary history of the San Joaquin basin
differs from that of the later Cenozoic: the former is characterized by a
few long-lasting basinwide depositional sequences, whereas the latter is
characterized by shorter sequences of more local extent. This change in
style of sedimentation took place during the Oligocene at about the
beginning of the transition from a convergent continental margin to a
transform margin. Paleogene basin history was controlled principally
by subduction-related and proto—San Andreas fault-related tectonics
and, to a lesser extent, the effects of changing eustatic sea level. A
eustatic fall in sea-level was probably the principal cause for the
regression recorded at the end of the upper Paleocene and lower Eocene
depositional sequence, and it contributed to most other Paleogene
regressions. Tectonic events related to the approach of the Paciﬁc-
Farallon spreading ridge became important in the Oligocene. Neogene
basin history was controlled principally by the tectonic effects of the
northwestward migration of the Mendocino triple junction along the
California continental margin and by the subsequent wrench tectonism
associated with the San Andreas fault system. Compression normal to
the San Andreas in the latest Cenozoic, resulting from changes in
relative plate motion, contributed to compressional deformation at the
west side of the valley. Eustatic sea—level effects are less discernible in

1

2 THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

the Neogene, but a middle Miocene highstand probably contributed to
the Widespread transgression at that time. Climate was an important
factor in basin history only in the latest Cenozoic (late Pliocene and
Pleistocene), when alpine glaciers in the Sierra Nevada and a pluvial
climate inﬂuenced sedimentation.

The San Joaquin basin, which at the end of the Mesozoic formed
the southern part of an extensive forearc basin, evolved during the
Cenozoic into today’s hybrid intermontane basin. Its evolution com—
prises the gradual restriction of the marine basin through uplift and
emergence of the northern part in the late Paleogene, closing off of the
western outlets in the Neogene, and ﬁnally the sedimentary inﬁlling in
the latest Neogene and Quaternary.

INTRODUCTION
SETTING

The Great Valley of California, a 700-km-long by up to
100—km-wide alluvial plain between the Sierra Nevada on
the east and the Coast Ranges on the west, is divided into
the Sacramento Valley in the north and the larger San
Joaquin Valley in the south. The south-ﬂowing Sacra-
mento River, draining the Sacramento Valley, and the
north-ﬂowing San Joaquin River, draining the northern
part of the San Joaquin Valley, join in the Sacramento-
San Joaquin Delta near Stockton (pl. 1, ﬁg. 1); the delta
in turn drains westward into San Francisco Bay. The
Tulare Lake basin in the south-central San Joaquin
Valley and the Buena Vista Lake basin at the extreme
south end of the valley contain closed depressions, and
both receive part of the drainage, at times, from the
Kings and Kern Rivers.

Geologically, the San Joaquin Valley is an asymmetric
structural trough with a broad, gently inclined, and
little-deformed east ﬂank and a relatively narrow west
ﬂank; the west ﬂank is a steep homocline in the northern
part of the valley, but becomes a belt of folds and faults
in the southern part of the valley (pl. 1). The trough is
ﬁlled with a prism of upper Mesozoic and Cenozoic
sediments that reaches a thickness of over 9 km in the
west-central part of the valley and at the south end. The
sedimentary prism represents, in a broad sense, the ﬁll of
the San Joaquin sedimentary basin; however, this basin
is, in a strict sense, a composite of a late Mesozoic and
early Cenozoic forearc basin that was largely open to the
Paciﬁc Ocean on the west, and a later Cenozoic trans-
form-margin basin. The basin-ﬁlling sediments rest on a
westward-tilted block of crystalline basement composed
of Sierra Nevada plutonic and metamorphic rocks under
the eastern part of the valley and maﬁc and ultramaﬁc
rocks of a presumed ophiolite of Jurassic age under the
central and western parts of the valley (Cady, 1975;
Page, 1981). On the west side of the valley, the Mesozoic
and early Tertiary Great Valley sequence, together with
the conformably underlying ophiolite, is juxtaposed with

 

the Franciscan Complex along a boundary fault termed
the Coast Range thrust as ﬁrst proposed by Bailey and
others (1964). For many years, the Coast Range thrust
was interpreted as a fossil subduction zone, but recent
work based on seismic reﬂection and refraction suggests
that the Franciscan has been thrust eastward as a wedge
between crystalline basement below and the Great Val-
ley sequence above (Wentworth and others, 1983; Went-
worth and others, 1984) (pl. 1, sections A, B). The
boundary between the Franciscan and the Great Valley
sequence, then, becomes the roof thrust of the wedge
and, unlike the hypothesized Coast Range thrust, ex-
tends no farther east than the tip of the wedge.

The Great Valley sedimentary basin is divided by the
buried, transverse Stockton arch and Bakersﬁeld arch.
The Stockton arch, which is a broad structure that is
bounded on the north by the Stockton fault (pl. 1) but has
a poorly deﬁned southern limit, separates the San
Joaquin and Sacramento sedimentary basins. The
Bakersﬁeld arch separates the Maricopa—Tejon subbasin
at the south end of the San Joaquin Valley from the
remainder of the San Joaquin sedimentary basin. Neither
arch has appreciable structural relief, but they did have
an inﬂuence on sedimentation, as will be shown. The
Tertiary depocenters of these basins are approximately
coincident with the Pleistocene and Holocene Buena
Vista and Kern Lakes basins in the south and the Tulare
Lake basin in the central part of the valley. The Teha-
chapi-San Emigdio Mountains uplift that bounds the
valley on the south might be considered a third trans-
verse structure. Cenozoic strata in the San Joaquin
Valley thicken southeastward from about 800 m over the
western part of the Stockton arch to over 9,000 m in the
Maricopa—Tejon subbasin in the south (pl. 1, section D).
The Mesozoic and early Tertiary Great Valley sequence,
on the other hand, thins southeastward and is apparently
absent south of the Bakersﬁeld arch.

PREVIOUS WORK

Although geological observations of rocks bordering
the San Joaquin Valley date back to the late 1800’s
(Shedd, 1932), the valley itself received scant attention at
that time. Geological study of the sedimentary deposits in
the valley was spurred on in the early 1900’s following the
discovery of oil at the McKittrick (1887), Coalinga (1887),
and Kern River (1901) oil ﬁelds. Knowledge of the
geology of the valley accumulated as oil exploration
progressed in the years preceding, and especially those
following, World War II.

The ﬁrst general review of the Cenozoic history of the
valley was that of Hoots and others (1954), although Reed
(1933) and Reed and Hollister (1936) touched on the San
Joaquin Valley in their broader summaries of California

GEOLOGY 3

and southern California geologic history, respectively.
The Cenozoic history of the valley was updated by
Repenning (1960) in a succinct, but stratigraphically
comprehensive, summary and was further revised by
Hackel (1966).

Bandy and Arnal (1969) applied a new technique—
based on the analysis of benthic foraminiferal faunas from
marine Tertiary rocks in the southern part of the
valley—in an attempt to quantify basin subsidence and
uplift. Foss’ (1972) interpretation of the Tertiary marine
stratigraphy emphasized the apparent synchroneity of
depositional cycles on the east and the west sides of the
valley, even though a different type of depositional
sequence characterizes each area. ‘

More recently, geologists have gone beyond describing
the basin stratigraphy and history and have sought
tectonic mechanisms that have controlled different as-
pects of basin evolution. Nilsen and Clarke (1975) dis-
cussed the tectonic setting for Paleogene sedimentation
in the Great Valley together with other regions of
California. Harding (1976) tied the structural evolution of
the west-side fold belt to the history of movement on the
San Andreas fault. Subsequently, others (notably Blake
and others, 1978; Dickinson and Snyder, 1979; Howell
and others, 1980) have related the origin of the San
Joaquin basin, along with other California Neogene
basins, to plate tectonic processes.

PURPOSE AND APPROACH

The theory of plate tectonics, since its emergence in
the 1960’s, has given a new perspective to the interpre-
tation of regional tectonics (Blake and others, 1978;
Howell and others, 1980; Page and Engebretsen, 1984).
Concurrently, revisions in regional stratigraphy, result-
ing mostly from the application of recent improvements
, in global correlations and the work toward a standard

global chronostratigraphic scale, allow more precise
comparison of the timing of depositional events from
widely separated parts of the San Joaquin basin. The goal
of this report is to interpret the Cenozoic sedimentary
record of the San Joaquin Valley in terms of external
controls on sedimentation and to speculate, where possi-
ble, on the nature of the tectonic events responsible.

This report is chieﬂy a review of the Cenozoic geologic
history of the San Joaquin basin in the light of current
ideas on plate tectonics, regional tectonics, and eustatic
sea-level change, although no exhaustive effort has been
made to summarize all the extensive geological literature
on the basin. It incorporates the results of several years
of U.S. Geological Survey research by myself and others
(notably D.E. Marchand, B.F. Atwater, J .W. Harden,
and W.R. Lettis) on the San Joaquin Valley, the purpose
of which was to elucidate the regional tectonic setting of

 

the valley as background for more speciﬁc geologic
hazards studies required in the siting, design, and
construction of nuclear powerplants. A basin-study ap-
proach to understanding regional tectonic history was
adopted. This assumes a basic cause and effect relation
between tectonics and sedimentation, and also that the
sedimentary ﬁll in a basin is a record of tectonic activity.

GEOLOGY

Although the San Joaquin Valley, as the southern part
of the Great Valley, constitutes part of a discrete
geomorphic and structural province within the western
Cordillera of North America, the geology is internally
variable in both stratigraphy and style of deformation.
Stratigraphically, the greatest variation is in the N eo-
gene deposits, which comprise a thick section of marine
sediments in the southern part of the basin but a
relatively thin and entirely nonmarine section in the
northern part. Structurally, the greatest differences are
between the west-side fold belt and the little-deformed
sedimentary cover of the Sierran block on the east side of
the valley.

To facilitate description of the geology, the valley is
subdivided into ﬁve regions on the basis of structural
style (ﬁg. 1, pl. 1). Although each region is structurally
distinct in style of deformation and tectonic history, the
boundaries between areas are necessarily arbitrary.

NORTHERN SIERRAN BLOCK

The northern Sierran block region of the San Joaquin
Valley consists of the stable east limb of the valley
syncline from the Stockton fault on the north to about the
San Joaquin River on the south. The region, which is the
least deformed part of the San Joaquin basin, includes as
its dominant element the broad and poorly deﬁned
Stockton arch.

The Stockton arch is evident principally as an area
where Paleogene and uppermost Cretaceous strata have
been erosionally truncated (Hoots and others, 1954).
There is little evidence of arching in overlying Tertiary
units (Bartow, 1985) and no evidence of basement arching
(Bartow, 1983) (pl. 1, section D). This structure probably
formed initially in the latest Cretaceous or Paleocene,
perhaps by local thickening of the Cretaceous section,
and had a major period of uplift in the Oligocene. Its
origin will be discussed further in the section “Regional
Tectonics.” The structure was a low-relief positive fea-
ture through most of the Paleogene.

The stratigraphy of the Modesto-Merced area is typical
of that on the northeast side of the valley (pl. 2, col. 3);
farther west, the stratigraphy is similar to that of the

THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

 

 

  
  

      

 

   
   
 

 

 

 

 

 

8 8
30° {‘1 5’90 {‘1
/\ /
(D
I I -'
‘ l I rn
San FranCIsco I 33
' ' E
:0 \
3 39°
F ’ o ]>
—I
D ‘n
E/ E
a» 0 CI B‘PLANATION <3
A; \ n ------ Boundary of structural region i‘
N O E —Dashed where approximate
—-' Fault —Dashed where a r ximatel
O < DP 0 )1
IV 5‘: E) E located; dotted where concealed
>
(J U)
I > 5: A 1
O ‘7 m >
N 5 7 E
>11 "’ 2 SACRAMENTO
VALLEY
N
\
O 390
SAN
JOAQUIN
\‘E Z VALLEV
5s ,.
36° \ -; 017,735."
c lye .....
I (A 7’ .
| m ...... INDEX MAP OF CALIFORNIA
I 33 y
I V m
t ‘ Z Z
.1" l ‘ O m :8
I. m < a
I m '“ )>
O I \ 1; O \ 9).,
('3 r >
In I; a
D
050 > If
\ 2 E
0
U)
0
I
l
l
\
96°
(5 4/
x I' . ountalns
J40 \ goo/<3 Tenachap‘ GARLOCK FAULL , a ’ '
K
% ‘K
B x o 10 20 30 40 so KILOMETERS
I_I_I__I___._I._I
\
.3
/ / / 5°
& R
o 5% s

FIGURE 1. —Index map of central California showing the ﬁve structural regions of the San Joaquin Valley in relation to
principal geographic and structural features. NSB, northern Sierran block; SSB, southern Sierran block; NDH,
northern Diablo homocline; WFB, west-side fold belt; M-TS, Maricopa-Tejon subbasin and south-margin deformed

belt.

GEOLOGY 5

Orestimba Creek area on the west side of the valley (pl.
2, col. 1). Over the Stockton arch (pl. 1), Paleogene strata
are absent and nonmarine Neogene strata rest directly
on the Great Valley sequence (Church and Krammes,
1958; Bartow, 1985). The Cenozoic deposits in this part of
the valley are relatively thin (about 1,100 m) compared to
the southern part of the valley, whereas the underlying
Great Valley sequence is much thicker (over 3,000 m)
(Hoffman, 1964) (pl. 1, section A).

Cenozoic deformation has consisted mostly of west or
southwest tilting of the rigid Sierran block as evidenced
by the subtle angular unconformities with discordances of
generally less than 1° that separate the Cenozoic units
along the northeast side of the valley (Grant and others,
1977; Marchand and Allwardt, 1981). Evidence of the
earliest Cenozoic tilting, however, is provided by the
truncation of Upper Cretaceous and Paleocene units by
Eocene strata in the subsurface (Bartow, 1985). Discord-
ance between Eocene and younger units is much less
apparent than that between Eocene and pre-Eocene
units, but there is some suggestion of tilting in the
Oligocene, based on the differences in gradient of depo-
sitional surfaces in the Ione and Valley Springs Forma—
tions at the eastern edge of the valley (Marchand, 1977).
These data are questionable, however, because a lack of
traceable markers and erosional relief at the contacts
makes it difﬁcult to reliably determine gradients or dips
in these units. Although there is little reliable evidence of
Oligocene westward tilting, truncation of older units over
the Stockton arch indicates uplift of that area, which, as
the southern part of the basin continued to subside during
the Oligocene, produced a southward tilt rather than a
westward tilt. A hiatus representing most of the Oligo-
cene is evidence that there was negligible subsidence in
the western part of the block during that interval. In the
Miocene section, there is little evidence of discordance
between the Valley Springs and Mehrten Formations,
although this evidence is again difﬁcult to assess accu-
rately. Later Neogene and Quaternary units, however,
do show appreciable differences in gradient (Marchand,
1977; Marchand and Allwardt, 1981). It seems most
probable that, while there may have been regional uplift
or southward tilting of the northern Sierran block in the
middle Tertiary, there was little southwest tilting until
the late Miocene or Pliocene. Tilting continues to the
present, probably at an accelerating rate.

Most of the Cenozoic faulting is localized along the
Foothills fault system, consisting of the Bear Mountains
and the Melones fault zones (pl. 1, ﬁg. 1). This fault
system originated as a major Mesozoic (pre-batholith)
shear zone or suture and has been locally reactivated in
the Cenozoic. Studies of the fault system by Woodward—
Clyde Consultants for Paciﬁc Gas and Electric Co}
revealed several locations where Cenozoic deposits are
offset across the fault zones. Displacement at one locality

 

on the Melones fault zone has been demonstrated to be
younger than about 4 Ma (Bartow, 1980), and Quaternary
movement is indicated elsewhere by shearing or offset of
soils or Pleistocene colluvial deposits (Marchand, 1977;
Schwartz and others, 1977). There is, however, no
indication of present-day seismicity along these zones
(Wong and Savage, 1983). Cenozoic normal faulting in the
foothills belt, and elsewhere near the tectonic hinge line,
was probably approximately coincident with the tilting;
the faulting suggests that the west or valley side of the
Sierran block was subsiding faster than the Sierra
Nevada was rising, resulting in tensional faulting near
the hinge.

Features possibly related to the Cenozoic normal
faulting are the numerous northwest-trending linea-
ments in the eastern part of the valley (Marchand and
Allwardt, 1978; Hodges, 1979). Most of these are proba-
bly not faults, but a few have 1 to 2 m of normal
displacement of Pleistocene units (Marchand, 1977).
Small normal faults with offsets of a few meters are
present at a few localities in exposed Tertiary strata as
well. There also appears to have been substantial dis-
placements of the Eocene Ione Formation along a system
of lineaments east of Merced (Marchand, 1977; Marchand
and Allwardt, 1978), but detailed gravity proﬁles across
one of the more prominent lineaments do not support
signiﬁcant offset of the basement surface. Data from one
proﬁle permit 2 m of fault offset of the basement surface
Within a few meters of the lineament; however, the
interpretation of these gravity data is ambiguous because
of uncertainties about erosional relief on the basement
surface, lateral density variations in the basement, and
density variations in overlying sediment (A. Griscom,
written commun., 1978).

Few subsurface faults have been recognized in the
northern part of the San Joaquin Valley; the largest of
the known subsurface faults is the Stockton fault, which
bounds the Stockton arch on the north (pl. 1). The
Stockton fault is a south-dipping reverse fault that trends
transversely to the regional structure. The fault, which
appears to have a complex history, has a total down-
to-the-north dip slip of as much as 1,100 m (Hoffman,
1964). It may have originated in the Late Cretaceous as
a normal fault, or possibly a left-lateral strike-slip fault
with a south-facing scarp (Teitsworth, 1964). It was
reactivated as a reverse fault in the latest Cretaceous or
Paleogene and was probably active through the early
Miocene, although most of the down-to-the-north offset
occurred during the Oligocene (Hoffman, 1964;
Teitsworth, 1964; Bartow, 1985). In the Merced-
Chowchilla area, another west-northwest—trending fault

1Unpublished report on geology of proposed Stanislaus nulear project by Woodward-Clyde
Consultants for Paciﬁc Gas and Electric Co., 1977.

6 THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

is recognized mostly on the basis of apparent offset of the
post-Eocene unconformity (Bartow, 1985). Its inferred
trace appears to coincide with a diffuse surface lineament
visible on satellite images (Antonnen and others, 1974;
Hodges, 1979). The lineament, termed the Kings Canyon
lineament, crosses the valley north of Chowchilla, paral-
lels the south fork of the Kings River in the Sierra
Nevada, and continues southeastward nearly to Death
Valley. The ophiolite remnant at Del Puerto Creek and a
major bend in the Ortigalita fault lie on the northwest
projection of the lineament in the Diablo Range.

The reverse faulting on the Stockton fault, and the
consequent elevation of Cretaceous rocks south of the
fault to form the Stockton arch, indicates a general
north-south compression from about the Paleocene
through the early Miocene. Late Cenozoic southwest
tilting of the Sierran block and subsidence of the west
side, with concurrent normal faulting near the hinge,
suggest east-west to northeast-southwest extension. The
normal faulting, however, insofar as it represents ﬂex-
ural stress as the west side of the block subsided faster
than the east side rose, is not necessarily evidence of
regional extension. Present-day seismicity in the south—
ern part of the northern Sierran block indicates north-
south to northeast—southwest compression producing
predominantly strike—slip and reverse faulting at depths
greater than 12 km (Wong and Savage, 1983).

SOUTHERN SIERRAN BLOCK

The southern Sierran block comprises the southern
part of the stable and little—deformed east limb of the
valley syncline. Its south boundary is the crest of the
Bakersﬁeld arch, a broad southwest-plunging ridge of
basement rock, and the north boundary is arbitrarily
placed at the San Joaquin River (ﬁg. 1, pl. 1). Both
Cenozoic and Mesozoic sedimentary deposits thicken
gradually southward and together total more than 5,000
m in the area south of Tulare Lake. Cenozoic strata alone
reach a thickness of more than 4,500 m, whereas the
Mesozoic rocks thin southeastward and pinch out or are
truncated against the north ﬂank of the arch (pl. 1,
section D).

The stratigraphy of the Bakersﬁeld arch area (pl. 2,
col. 9) is typical of the southern part of this area. Outside
the arch area, Tertiary rocks, particularly the older
Tertiary, are not well known because of the absence of
Tertiary outcrops between the San Joaquin River and the
Tule River and because of the sparsity of deep wells. For
much of the Neogene, the central part of the Southern
Sierran block was a broadly ﬂuctuating zone of transition
between nonmarine deposition on the north and east and
marine deposition on the south and west. The Cenozoic

 

stratigraphy in the subsurface of the Hanford-Tulare
area (pl. 2, col. 6) probably has some similarity to both the
Bakersﬁeld arch area to the south and the Kettleman
Hills area to the west.

The San Joaquin Valley part of the southern Sierran
block is structurally similar to the valley part of the
northern Sierran block; differences are principally in
degree of deformation rather than in style of deforma-
tion. The southwest-to—west tilt of the entire Sierran
block increases southward so dips of exposed Tertiary
units in the Bakersﬁeld area average 4°-6°, in contrast to
the 1°—2° dips in the north (pl. 1, sections A, B, C).
Further indication of the increased tilt is provided by the
greater height of the southern Sierra Nevada and the
greater depth to basement in the southwestern part of
the valley, although part of this difference in elevation is
a result of normal faulting along the west edge of the
southern Sierra Nevada.

Truncation of Cretaceous and Paleocene(?) strata indi-
cates tilting prior to the middle Eocene of the southern as
well as the northern Sierran block. Minor angular dis-
cordance between the Walker Formation or Vedder Sand
and overlying J ewett Sand is evidence of tilting near the
end of the Oligocene. An unconformity at the base of the
“Santa Margarita” Formation that truncates older units,
and a more extensive unconformity and truncation at the
base of the Kern River Formation are evidence of the
accelerating uplift and westward tilt of the southern
Sierran block beginning in the late Miocene.

Normal faults along the east side of the valley are
concentrated in the area of the Bakersﬁeld arch. These
faults generally trend northwest to north, although a
secondary west to west-northwest trend is apparent (pl.
1). The net displacement is down to the southwest,
although down-to-the-northeast faults are present (Bar-
tow, 1984). One of the principal faults of this group is the
Kern Gorge fault, along which basement rocks to the
southwest have been downdropped more than 600 m. An
important exception to the northwesterly fault trend is
the Poso Creek fault that trends in a westerly direction
through the Tertiary outcrop belt and then curves to the
northwest to merge with the subsurface Pond fault.

Faulting appears to die out northwestward along the
east edge of the valley, partly because Quaternary
deposits overlap the Tertiary strata onto the basement
rocks north of Deer Creek. Buried (or partially buried)
faults that offset the basement rock surface have been
inferred in the area between Porterville and Dinuba
(Croft and Gordon, 1968) and near Clovis northeast of
Fresno (Page and LeBlanc, 1969). The presence of these
inferred faults is based principally on surface lineaments
and a steeply west-sloping basement surface, but no fault
offsets have been convincingly demonstrated. These
faults are not included on plate 1.

GEOLOGY 7

Many subsurface faults have been inferred west of the
Tertiary outcrop belt by others (Los Angeles Depart-
ment of Water and Power, 1975, ﬁg. 2.5.1—7A). Most of
these faults seem to be small and have a predominant
northwest trend; they have been recognized only where
well density is sufﬁcient for delineation of faults, that is,
mainly in oil ﬁelds. No attempt has been made to
generalize these on plate 1. The Pond fault and Greeley
fault system are, however, major structures. The Pond
fault, actually a zone of subparallel southwest-dipping
normal faults up to 2 km Wide, apparently joins the Poso
Creek fault to the southeast (Los Angeles Department of
Water and Power, 1975). Down-to-the-southwest offsets
decrease upward from a maximum of over 500 m on the
basement surface. Near the town of Pond, a zone of
cracks extends to the ground surface, which has been
downdropped as much as 23 cm across the fault (Holzer,
1980).

The buried Greeley fault system consists of an en
echelon set of northwest-trending normal faults (pl. 1).
The basement surface is downdropped on the northeast
as much as 615 m, but offsets decrease upward so there
is no apparent offset of strata younger than late Miocene
(Los Angeles Department of Water and Power, 1975).
The Greeley fault has been reported to have a large
component of lateral displacement (Sullivan and Weddle,
1960). Webb (1977) inferred 670 m of right slip on the
basis of apparent offset of Miocene channel sands, but he
later reinterpreted (Webb, 1981) the apparent offsets as
meanders in the channels. The Greeley fault system is
paralleled on the southwest by a series of short low-
amplitude folds that have their strongest expression in
early Miocene and older strata. This folding could be (1)
genetically related to strike-slip faulting, (2) due to
compression that also would have produced reverse
faulting, as is found on faults with similar trend farther
west in the basin, or (3) a result of drape over a buried
fault scarp. An intensive study of the Greeley fault
system for a proposed nuclear powerplant site astride the
northern segment of the fault, based on seismic-reﬂection
proﬁles and oil well data (Los Angeles Department of
Water and Power, 1975), concluded that the movement
has been normal (down-to-the—east) and that there is no
evidence of lateral displacement. This study apparently
did not consider the possibility of reverse movement;
however, the geometry of the Greeley structure, as seen
on seismic-reﬂection sections, is sufﬁciently different
from reverse fault structures like Semitropic anticline to
seriously weaken the reverse fault hypothesis for the
origin of the Greeley structure.

A large number of northwest-trending surface linea—
ments in the Kern River area were described by Warne
(1955), who implied a relation to deep lateral faulting.
More recent trenching of selected lineaments, however,

 

shows no evidence of near-surface faulting (Los Angeles
Department of Water and Power, 1975). The surface
lineaments in the southeastern part of the valley are
similar to those (described above) in the northeastern
part of the valley. Their origin is unclear, but in neither
case do they seem to be related to bedrock faults.

In contrast with the northern Sierran block Where the
normal faulting is mostly late Cenozoic in age, the
faulting appears to be mostly Miocene or older in the
southern part of the block. Although it is difﬁcult to
determine the time of inception of the normal faulting,
subsurface evidence from both well sections (Bartow,
1984) and seismic sections (Los Angeles Department of
Water and Power, 1975) indicates greater offset of the
basement surface than of late Miocene or Pliocene
horizons. Faults with a general north trend (northwest to
north-northeast) and those, such as the Poso Creek fault,
with a general west trend (west to west-northwest) seem
to be similar in that offset decreases upward. In general,
north-trending faults were active in the early Tertiary
and again beginning in the late Miocene. West-trending
faults may have had their origin in the latest Oligocene
and early Miocene like those in the Maricopa-Tejon
subbasin at the south end of the San Joaquin Valley, as
Will be shown in the section “Maricopa-Tejon Subbasin
and South-Margin Deformed Belt,” and were probably
active until about the late Miocene. The Greeley fault,
farther west and near the center of the valley, may have
had its origin in the Cretaceous or earliest Tertiary and
shows no offset of horizons younger than early Miocene
(Los Angeles Department of Water and Power, 1975).
Part of the movement of the normal faults in the southern
Sierran block was, then, concurrent with the late Mio-
cene to recent uplift of the Sierran block (pl. 2). However,
a signiﬁcant part of the offset was pre-uplift, and faulting
seems to have ended in the Pliocene while uplift presum—
ably continued. There are few faults on the north side of
the Bakersﬁeld arch that offset Quaternary deposits. The
exceptions seem to be largely due to subsurface compac-
tion as a result of ﬂuid withdrawal—oil in the case of the
Kern Front and Premier faults (Castle and others, 1983;
Bartow, 1984) and ground water in the case of the Pond
fault (Holzer, 1980). In addition to the normal faults
involving basement rocks, a number of syndepositional
growth faults formed during late Miocene sedimentation
in the area west of Bakersﬁeld (MacPherson, 1978).

North-trending normal faults of pre-Miocene age (in-
cluding northwest-trending faults like the Greeley and
Pond) indicate approximate east-west to northeast-
southwest extension in the early Tertiary; these faults
may be the only manifestation of a north-south regional
compressive stress at that time, although there may be
some possibility of minor right-lateral strike slip on
northwest—trending faults like the Greeley fault. East-

8 THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

trending normal faults indicate north-south extension for
the period during which they were active, probably about
late Oligocene to late Miocene.

NORTHERN DIABLO HOMOCLINE

The northern Diablo homocline consists of the west
limb of the valley syncline from the Stockton arch in the
north to Panoche Creek in the south. It includes the
northeast ﬂank of the northern Diablo Range (ﬁg. 1, pl. 1).

The stratigraphy of the Orestimba Creek area and the
Los Banos-Oro Loma area (pl. 2, cols. 1 and 2) are
representative of the Diablo homocline. Approximately
1,400 m of Cenozoic deposits in that area thin northwest-
ward toward the Stockton arch, mostly through trunca-
tion of the marine older Tertiary units (Hoffman, 1964;
Hackel, 1966; Bartow and others, 1985). Paleogene
marine rocks, however, reappear in the Corral Hollow-
Lone Tree Creek area southwest of Tracy. The present
Diablo Range has resulted principally from Neogene
tectonism, although there is some evidence that the
northern Diablo Range existed as a positive area as far
back as the Paleogene (Clarke and others, 1975; Nilsen
and Clarke, 1975; Bartow and others, 1985). The relation
between the Paleogene Diablo uplift and Stockton arch is
unclear, but Neogene structures in the northern Diablo
Range appear to be superimposed on the older positive
areas.

The Cenozoic rocks of the northern Diablo Range form
a northeast-dipping homocline in which the dips of the
Tertiary strata generally range from 30° to 50° (pl. 1).
Subordinate structures are principally faults, but folds
are associated with the Vernalis and Black Butte faults
near Tracy at the west end of the Stockton arch and a
small anticline near Patterson produces a local reversal of
dip in the homocline. Near Gustine, the dips of Tertiary
strata ﬂatten abruptly to 10° or less northeast of a
northwest-trending fault.

Folding or tilting is mostly of Neogene age, although
some deformation did take place in the Paleogene. The
slight unconformity between the Great Valley sequence
and overlying Paleogene units is evidence of mild defor-
mation in earliest Tertiary time. Truncation of Paleogene
units at the base of the Valley Springs Formation is
apparent along the Diablo homocline, as it is elsewhere
along the south side of the Stockton arch; this relation
indicates post-Eocene uplift or tilting of the arch. Al—
though the beginning of Neogene uplift of the Diablo
Range is evidenced by the formation of coarse alluvial fan
deposits derived from the range in the late middle to late
Miocene (pl. 2), the angular unconformity at the base of
the latest Pliocene and Pleistocene Tulare Formation
marks the principal uplift of the range (Bartow, 1985).

 

The principal Cenozoic faults or fault zones of the
northern Diablo homocline are (1) the Black Butte fault,
a northwest-trending fault west of Tracy, (2) the Vernalis
fault, a subsurface fault that parallels the Black Butte
fault and trends at a right angle to the Stockton fault near
its west end, (3) the Tesla-Ortigalita fault zone, the west
boundary of the Diablo homocline and the present bound-
ary between the Franciscan Complex and the Great
Valley sequence, and (4) the San Joaquin fault zone,
which lies along the west edge of the valley. The history
of faulting is complex and, for some faults, not well
understood. As with the folding, most Cenozoic faulting
seems to be Neogene in age and it is difﬁcult to identify
speciﬁc structures as Paleogene in age because of the
strong overprint of Neogene tectonism.

The Black Butte and Vernalis faults are subparallel,
southwest-dipping reverse faults, eachwith an associ-
ated anticline on the upthrown side. The Black Butte
fault involves units as young as the Tulare Formation
and, therefore, must have been active as recently as the
Pleistocene (Raymond, 1969). The age of the Vernalis
fault is less well known because it is an entirely subsur-
face structure. Large offset at the base of the Valley
Springs Formation (Bartow, 1985) suggests that most of
the movement took place in the Miocene or later.
Although the upper limit of faulting is not known and
there is no evidence of Quaternary movement, the
fault-plane solution for a 1977 magnitude 3.5 earthquake
near Patterson approximately on-trend to the southeast
indicates the same style of faulting as for the Black Butte
fault (Wong and Ely, 1983).

The Tesla-Ortigalita fault is a zone of high-angle faults
with a net down-to-the-east displacement that may total
thousands of meters. The dip of the fault plane is not
everywhere known, but it is locally a southwest-dipping
reverse or thrust fault (Briggs, 1953). In the big bend
segment of the Tesla-Ortigalita fault between Hospital
and Del Puerto Creeks, the fault plane generally dips
steeply in the direction of the downthrown block, which
suggests that it is locally a normal fault (Maddock, 1964;
Raymond, 1969). As Maddock pointed out, however, the
fault may have been folded subsequent to its formation.
Although movement on the fault zone has been predom-
inantly dip slip, it is not known how much was in a
reverse sense and how much might have been in a normal
sense. The southern part of the zone from Quinto Creek
to Little Panoche Valley shows evidence of right-lateral
strike-slip displacement (Lettis, 1982; Anderson and
others, 1982), and numerous fault-plane solutions for this
part of the zone show chieﬂy right-lateral displacement
(LaForge and Lee, 1982). The Holocene strike slip,
however, may have been only recently superimposed on
the predominant dip slip.

GEOLOGY 9

The history of the Tesla-Ortigalita fault zone is very
poorly understood. What seems to be a nearly continuous
fault zone may actually be an aggregate of fault segments
having different origins and different histories. Some
segments may have originated during Paleogene uplift,
but most of the dip slip occurred in the Neogene. Similar
elevations for an isolated 9-Ma basalt ﬂow east of the
fault zone near San Luis Reservoir and the base of the
upper Miocene Quien Sabe volcanic field west of the fault
zone led Lettis (1982, 1985) to conclude that there had
been no appreciable differential vertical movement since
the late Miocene. More recent work, however, suggests
that the isolated ﬂow was derived from a local vent in or
near the fault zone itself and had no connection with the
Quien Sabe Volcanics (D.H. Sorg, oral commun., 1986).
The absence of vertical offset of Quaternary units across
the southern segment of the fault zone indicates that
there has been no appreciable Quaternary dip slip, but
this segment does show evidence of late Cenozoic strike
slip (Lettis, 1982; Anderson and others, 1982). The
northern segment of the fault zone, the Tesla fault
proper, juxtaposes upper Miocene deposits with the
Franciscan and, therefore, must have had considerable
post-late Miocene dip slip. The history of the big bend
segment of the fault between Hospital and Del Puerto
Creeks is much more difﬁcult to assess. The present bend
in the fault trace might be due to folding since the time of
fault formation. If so, the timing of the folding is not
known.

The San Joaquin fault zone is marked by a series of
east-facing scarps and offset Quaternary depositional
surfaces that were interpreted by Herd (1979b) as
evidence of down-to-the-east normal faulting. Along
much of the zone’s length, however, the inferred faults
are covered by upper Pleistocene and Holocene alluvium,
so that there is some question about both the continuity
and the dip of the faults. Bartow (1985) reinterpreted the
zone as a series of reverse faults, a conclusion which
seems more compatible with the regional framework.
Available evidence, which shows that units as young as
Pleistocene are offset, suggests that the San Joaquin
fault zone may have been active at the same time as the
Black Butte and Vernalis faults.

A set of subparallel faults between the Tesla-Ortigalita
and the San Joaquin fault zones south of San Luis
Reservoir (greatly generalized on pl. 1) was termed the
O’Neill fault system by Lettis (1982, 1985). This fault
system consists of numerous northeast-dipping faults
that offset Quaternary pediment surfaces by as much as
100 m (Lettis, 1982, 1985). The faults are apparently
bedding-plane slips in the underlying Great Valley se-
quence that formed in response to the strong bending of
the upturned strata. These faults caused offsets of

 

Quaternary erosion surfaces and their associated depos-
its that lie across the beveled edges of the Great Valley
sequence.

One of the most fundamental structural features of the
Diablo Range, as well as elsewhere in the Coast Ranges,
is the contact between the Franciscan Complex and the
Great Valley sequence (pl. 1, sections A, B). The original.
contact, although it has been greatly modiﬁed by younger
faults like the Tesla-Ortigalita, is apparently tectonic
(Page, 1981). This fault contact, commonly termed the
Coast Range thrust, is interpreted as the roof thrust of a
Franciscan wedge (Wentworth and others, 1983; Went-
worth and others, 1984) that has had an inﬂuence on
Cenozoic regional tectonics since, perhaps, as early as the
Paleogene.

Paleogene deformation of the northern Diablo Range,
which seems to have consisted mostly of broad regional
uplift, implies northeast-southwest compression; how-
ever, the orientation of the stress cannot be determined
with any certainty. Neogene structures reﬂect a general
northeast-southwest compression, but latest Cenozoic
right—lateral strike slip and seismicity on the southern
segment of the Tesla-Ortigalita fault indicate a north-
south or north-northeast—south-southwest compression
producing a northwest-southeast shear, at least for the
area between San Luis Reservoir and Panoche Valley.

WEST-SIDE FOLD BELT

The west-side fold belt extends along the southwest
side of the valley syncline from about Panoche Creek on
the north to the Elk Hills in the southwesternmost San
Joaquin Valley. The belt includes the southern Diablo
and Temblor Ranges and is characterized by Cenozoic
folds and faults that trend, for the most part, slightly
oblique to the San Andreas fault on the southwest (pl. 1).

The stratigraphy of the west-side fold belt is variable,
as might be expected in a tectonically active area.
Stratigraphic columns for four separate areas—the Val-
lecitos syncline, Kettleman Hills north dome, Lost Hills-
Devils Den area, and Elk Hills area (pl. 2, cols.
4,5,7,8)—provide some indication of the variation. Total
thickness for the combined Mesozoic and Cenozoic section
may be over 9,500 m near the San Joaquin Valley syncline
axis. As with the southern Sierran block part of the
valley, there is a northward-thinning trend for the
Cenozoic (pl. 1, sections B, C) and, particularly for the
Neogene, a northward trend toward shallower marine
and nonmarine facies. Middle Tertiary deposits repre-
senting some of the deepest water in the San Joaquin
basin are found in the southern Temblor Range. Older
rocks are not as well known in the southern part of the

10 THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

fold belt because of the absence of outcrops and the
sparsity of wells that reached Paleogene strata.

The northernmost fold in the west-side fold belt is the
Vallecitos syncline, located just south of Panoche Creek.
The southeast boundary of the fold belt is arbitrarily
placed east and south of Elk Hills Where the fold trends
change from northwest to west. The east boundary
deviates from the valley syncline axis near Cantua Creek
and south of Kettleman Hills to include the subdued
Turk, Buttonwillow, Bowerbank, and Semitropic anti—
clines that, although lying east of the valley axis, are
structurally more akin to the west-side fold belt than to
the less-deformed southern Sierran block.

The intensity of deformation increases southeastward
along the fold belt as well as southwestward across the
belt toward the San Andreas fault (pl. 1). The increased
intensity is evidenced by tighter folds and an increased
number of reverse and thrust faults (Vedder, 1970;
Dibblee, 1973a). Thrust faults seem to be predominantly
west dipping, although the faulting in the interior of the
Temblor Range is complex. Recent thrust-fault-gener-
ated earthquakes at the Coalinga anticline (May 1983)
(Eaton, 1985b) and Kettleman Hills (August 1985) (Went-
worth, 1985) are evidence of thrusting beneath major
west-side folds and indicate the style of Holocene defor-
mation along the east side of the fold belt (Wentworth
and others, 1983; Wentworth and others, 1984; Namson
and Davis, 1984; Medwedeff and Suppe, 1986). The
west-side folds are, then, partly a reﬂection of deep-
seated thrust deformation that is related to emplacement
of a wedge of the Franciscan Complex.

Deﬂection of the shaleout line of the subsurface, lower
Eocene Gatchell sand (of local usage) around the down-
plunge end of the Coalinga anticline in the northern part
of the fold belt provides evidence that the anticline
probably formed in the Paleocene or early Eocene
(Harding, 1976). Paleogene deformation is difﬁcult to
identify in the southern part, however, because of the
deep burial of Paleogene rocks and the strong overprint
of Neogene deformation. Harding (1976) outlined the
Neogene development of the fold belt in relation to the
history of strike slip on the San Andreas fault. The ﬁrst
en echelon folds in the Temblor Range or southern part
of the fold belt date from the late early Miocene (near the
Saucesian-Relizian boundary), whereas the easternmost
anticlines in the fold belt (Buttonwillow, Bowerbank, and
Semitropic) are entirely Pleistocene in age. The age of
faulting in the fold belt is not well constrained, but
eastward-verging thrust faults seem to have formed
fairly late in the deformation history in the more tightly
folded area near the San Andreas fault (Harding, 1976)
and are still active at the west margin of the valley.

A fault along the southwest side of the Semitropic
anticline has been interpreted as a normal fault (Los

 

Angeles Department of Water and Power, 1975), but the
asymmetry of the fold and its abrupt southeast boundary,
as seen on proprietary seismic-reﬂection sections, sug-
gest that it may be a northeast-dipping reverse fault. A
structure interpreted as a reverse fault by Wentworth
and others (1983) appears on a seismic-reﬂection section
in a position on-trend with the Semitropic anticline fault
to the northwest. The association of the fault at Semi-
tropic With a fold is, in itself, suggestive of compressive
deformation. The age of the fault is difﬁcult to assess, but
if it is genetically related to the fold, it would be Pliocene
or Pleistocene in age. It may, however, be an older
structure that has merely served to control the location of
the younger fold.

The structures of the west-side fold belt cumulatively
indicate north-south to northeast-southwest compression
through the Cenozoic. During the early Paleogene and
most of the Neogene, this compression was apparently
manifested as a northwest-southeast shear couple. A
tendency for Pliocene and Pleistocene structures to be
oriented more parallel to the San Andreas fault indicates
an increasing component of compression normal to the
fault in the latest Cenozoic. Present-day seismicity at the
San Joaquin Valley-Coast Ranges boundary in the north-
ern part of the fold belt indicates continuing northeast-
southwest compression (Eaton, 1985a).

MARICOPA-TEJON SUBBASIN AND SOUTH-MARGIN
DEFORMED BELT

The Maricopa-Tejon subbasin and the south-margin
deformed belt are structurally distinct areas, but they
are probably genetically related. The Maricopa-Tejon
subbasin is located at the extreme south end of the San
Joaquin basin between the Bakersﬁeld arch on the north
and the deformed belt of the north ﬂank of the San
Emigdo Mountains on the south. These areas are
bounded on the east by the Tehachapi Mountains of the
southernmost Sierra Nevada and merge westward with
the southeast end of the west-side fold belt (ﬁg. 1, pl. 1).

The western part of the Maricopa-Tejon subbasin, the
Maricopa subbasin proper, is characterized by its great
depth—probably more than 9 km to basement in the
central part. The south-margin deformed belt is charac-
terized by the northward—directed thrust faulting at the
south edge of the basin. Extreme Neogene subsidence,
together with thrust faulting that resulted in several
kilometers of crustal shortening in the late Cenozoic
(Davis, 1983), is evidence that the south end of the San
Joaquin basin is the most highly deformed part (pl. 1,
section D).

The Maricopa subbasin contains the thickest Cenozoic
deposits in the San Joaquin basin. Neogene and Quater-
nary strata are more than 6,100 m thick at the Paloma oil

GEOLOGY 11

ﬁeld, a few kilometers east of Buena Vista Lake Bed (pl.
1). The thickness of Paleogene strata in the central part
of the basin is not known because few wells have reached
the Paleogene and none have reached basement; how-
ever, more than 1,750 m of Paleogene strata crop out
near San Emigdio Creek on the south side of the basin (pl.
2, col. 10) and a greater thickness might be present
downdip to the north.

There are no known Cretaceous or Paleocene deposits
south of the Bakersﬁeld arch. Eocene strata rest on
basement rocks in the San Emigdio Mountains and at
South Coles Levee oil ﬁeld at the west end of the
Bakersﬁeld arch (Church and Krammes, 1957), but no
wells have reached the basement in the intervening area.

Paleobathymetries recorded in the middle Tertiary
deposits of the Maricopa-Tejon subbasin are the deepest
found in the San Joaquin basin. Abyssal depths (about
1,800 m) were reached in the Zemorrian, Saucesian and
Luisian Stages (Bandy and Arnal, 1969). Paleogene
nonmarine strata were deposited on the east and south-
east, and the basin gradually shallowed through the late
Neogene and became entirely nonmarine in latest Plio-
cene time.

Structural trends are variable in the Maricopa-Tejon
subbasin and south-margin deformed belt, but there is a
general west trend along the south margin of the basin.
The northwest fold trends of the west-side fold belt
change to west-northwest where that region merges with
the deformed belt at the south end of the valley. The folds
and faults of the San Emigdio Mountains, dominated by
the Pleito thrust fault system, form a northward-directed
salient with an average west fold trend. To the north and
northeast, the northeast-trending White Wolf fault is the
dominant structure. The White Wolf fault and the
smaller Springs fault to the southeast both trend approx-
imately parallel to the Garlock fault, which lies along the
southeast side of the Tehachapi Mountains. Both faults,
like the Garlock, show some geologic evidence of left-
lateral movement. Farther northeast, the northwest- to
west-trending Edison fault is an older Tertiary normal
fault with down-to—the-north offset of over 1,500 m
(Dibblee and Chesterman, 1953; Bartow, 1984).

The south margin of the San Joaquin basin, in addition
to being the most highly deformed part of the basin,
probably has the most complex tectonic history. Evi-
dence of possibly the earliest deformation is provided by
paleomagnetic data that indicate a clockwise rotation of
the Tehachapi Mountains of 45° to 60° that took place
between 80—100 Ma and 16 Ma (Kanter and McWilliams,
1982; McWilliams and Li, 1985). Some of this rotation
probably occurred in the Late Cretaceous or early
Tertiary and the remainder took place after eruption of
volcanic rocks in the earliest Miocene (Plescia and Cal-
derone, 1986). Geologic evidence of the earliest Cenozoic

 

deformation is a major angular unconformity in the
western San Emigdio Mountains where upper Oligocene
and lower Miocene sediments of the Temblor Formation
overlap truncated older Tertiary units and rest on
basement rocks (Nilsen and others, 1973; Davis, 1986).
Although no speciﬁc faults can be positively identiﬁed as
having been active during the period of tectonism, which
may have begun in the late Eocene and extended into the
Oligocene, Davis (1986) suggested that Oligocene uplift of
the San Emigdio Range was produced largely by a major
south-verging thrust fault. The Caballo Canyon fault,
identiﬁed by Davis (1986) as the Oligocene thrust, is an
obscure fault (not shown on pl. 1) that has been subject to
other interpretations (Davis, 1983), so the hypothesized
Oligocene thrusting remains somewhat questionable.

Normal faults at the south margin of the basin have
generally west trends (from northwest to northeast) and
occur mainly in the subsurface (Hirst, 1986; Davis, 1986).
These faults were active during the latest Oligocene and
early Miocene, concurrent with volcanism dated at 22.1
to 22.9 Ma2 (Turner, 1970) and basin subsidence (Hirst,
1986; Davis, 1986). The mostly west-trending Edison
normal fault, as well as other normal faults of general
west trend in this region of the basin, was also active at
that time.

Most of the deformation of the San Emigdio Moun-
tains, including uplift and folding, is late Cenozoic and is
directly related to thrust faults of the Pleito fault system
(Davis, 1986). These thrusts date only from the Pliocene
and, on the basis of the earliest appearance of coarse
detritus in the basin to the north, most of the uplift was
in the late Pliocene and Pleistocene (Davis, 1986). Al-
though the basin continued to subside through the
Miocene, subsidence accelerated during the Pliocene
(Davis, 1986; Hirst, 1986).

The White Wolf fault, which was the locus of the
magnitude 7.2 Arvin-Tehachapi earthquake of July 1952,
is a southeast-dipping reverse fault (Oakeshott, 1955;
Stein and Thatcher, 1981). Total vertical separation on
the basement surface has been at least 3,600 m (Stein and
Thatcher, 1981) or possibly more than 4,600 m (Davis,
1983). Although seismologic data from the 1952 earth-
quake indicate a component of left-lateral slip (Oake-
shott, 1955; Stein and Thatcher, 1981), evidence for large
cumulative left-lateral displacement is ambiguous and the
total lateral offset may be small relative to the large
vertical offset (Davis, 1986).

The early history of the White Wolf fault is uncertain,
but it may have originated as a down-to-the-northwest
normal fault during the late Oligocene and early Miocene

2Dates have been converted from old to new (1977) constants according to Dalrymple (1979).

12 THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

period of normal faulting (Davis, 1986). There is strati-
graphic evidence of continued down-to—the-northwest
movement accompanying basin subsidence through the
Miocene, probably as a normal fault (Davis, 1986). More
recently, probably during the Pliocene or Pleistocene,
the conﬁguration of the fault was apparently changed to
the southeast-dipping reverse fault that it is today.

The stress regime in which the early Oligocene defor-
mation took place is obscure, but it presumably involved
an approximate north-south compression. The latest
Oligocene and Miocene normal faulting and volcanism, in
contrast, clearly indicates extension, probably with a
general north-south orientation. The Pliocene to H010-
cene thrust faulting is also a clear indication of strong
compressive tectonism, again with a north-south orien—
tation. The tectonic history of the south end of the San
Joaquin basin seems, therefore, to be one of alternating
periods of compression and extension, all in a north-south
direction.

MAJOR CONTROLS ON SEDIMENTATION

The Cenozoic stratigraphic record in the San Joaquin
basin seems, at ﬁrst glance, to be a far-from-ideal record
of basin history. It is incomplete, particularly in the
northern part of the basin; it is poorly dated in many
places; and, it exhibits wide variations in facies from area
to area. The gaps and complexity, however, are impor-
tant parts of the record and they can provide important
clues to basin history.

Stratigraphic sections from different parts of the basin
(pl. 2) can be informally divided, where they consist
mostly of marine deposits, into depositional sequences
that are composed of transgressive-regressive couplets
that formed principally in response to relative changes in
sea level in combination with sedimentation. A relative
rise in sea level, that is, the eustatic rise plus the effect
of subsidence, does not necessarily equate with trans—
gression, nor does a relative fall in sea level necessarily
equate With regression. If, for example, sedimentation
exceeds the relative rise of sea level, then a regression
will result.

The marine sedimentary sequences of this report,
which are commonly but not necessarily unconformity
bounded, facilitate correlation from area to area within
the basin. In a few cases an unconformity-bounded
marine sequence may be approximately correlated with a
nonmarine sequence on the basis of the bounding uncon-
formities. The fact that correlatable sequences are
present throughout the basin provides evidence for
external control of the sedimentary record, and although
the sequence is the primary record of basin history, the
unconformities bounding the sequence are equally impor-

 

tant. The goal in deciphering basin history is to identify
the cause or causes of each event.

The major external controls on sedimentation are
tectonism, eustatic sea-level change, and climate. The
sedimentary record represents the complex interplay of
all of these factors, although tectonism is clearly domi-
nant. Any thick accumulation of sediments, such as that
found in the southern San Joaquin basin, clearly implies
tectonic subsidence. Furthermore, the location of this
basin along an active continental margin virtually assures
tectonic activity in some form, throughout the Cenozoic.
The other factors, sea-level change and climate, play
important roles as well, perhaps more so than previously
recognized.

TECTONICS

Tectonics, as it applies to the Cenozoic San Joaquin
basin, includes basin subsidence, uplift of the adjacent
Sierra Nevada and Coast Ranges, and contemporaneous
deformation of the basin itself such as faulting and
folding. Closely associated with tectonism is volcanism at
the margins of the basin and in adjacent regions.

The various aspects of Paciﬁc Coast Cenozoic tectonics
should be explicable in terms of the interactions of the
crustal plates at the western edge of North America.
Since the ascendency of the plate tectonics paradigm,
knowledge of how plate interactions have inﬂuenced
regional tectonics on the California margin has grown
steadily. Before discussing the regional tectonic events
that played a part in the evolution of the San Joaquin
basin, it is appropriate to brieﬂy review the plate tectonic
events that are most relevant to the central California
part of the continental margin.

PLATE TECTONICS

A subduction zone has existed at the western margin of
North America throughout the Cenozoic (pl. 2). From the
Late Cretaceous until about the middle or late Eocene (75
to 40 Ma), the relative plate motions resulted in oblique
subduction of an oceanic plate, probably the Kula plate
(Page and Engebretson, 1984). Rapid convergence rates
during this period produced a low-angle subduction zone
and consequent displacement of arc magmatism eastward
from the Sierra Nevada into Colorado (Lipman and
others, 1972; Snyder and others, 1976; Coney and Rey-
nolds, 1977; Cross and Pilger, 1978). Oblique subduction
at the central California margin continued until nearly
the end of the Eocene when the Farallon plate, which had
a more normal component of motion relative to North
America, supplanted the Kula plate at central California
latitudes (Page and Engebretson, 1984). Slowdown in the

MAJOR CONTROLS ON SEDIMENTATION 13

convergence rates from the late Eocene into the Oligo-
cene resulted in steepening of the subduction zone and
the consequent migration of volcanism southwestward
from Idaho and Montana into Nevada (Lipman and
others, 1972; Snyder and others, 1976; Cross and Pilger,
1978).

The San Andreas transform originated in mid-Oligo-
cene time (28-30 Ma) (Atwater and Molnar, 1973; Enge-
bretson and others, 1985) when the ancestral East Paciﬁc
rise ﬁrst encountered the subduction zone. The term
“San Andreas transform” is used here, as it was by
Dickinson and Snyder (1979), for the whole system of
subparallel faults that constitute the plate boundary. The
initial slip was probably offshore on faults at or near the
continental margin (Garfunkel, 1973; Dickinson and Sny-
der, 1979). Slip on the San Gregorio-Hosgri fault zone and
San Andreas fault proper probably did not begin until
nearly middle Miocene time (about 16—17 Ma); most of the
initial slip was on the San Gregorio-Hosgri fault (Gra-
ham, 1978).

The transform lengthened as paired triple junctions
migrated northwestward and southeastward along the
continental margin. Relative positions of the North
American and Paciﬁc plates in the Neogene, recon-
structed according to the global-circuit method (Atwater
and Molnar, 1973), differ somewhat from the reconstruc-
tion by the hot-spot method (Engebretson and others,
1985). Although the difference between the two recon-
structions is within the limits of probable error (Enge-
bretson and others, 1985), the northward migration
history of the Mendocino triple junction from the global-
circuit method seems to provide the best ﬁt to geologic
history. The unstable conﬁguration of the migrating
triple junction (trench and transform not colinear) in-
duced a wave of extensional tectonism in nearby regions
(Dickinson and Snyder, 1979; Ingersoll, 1982). Local
volcanism in west-central California was approximately
coincident with the passage of the triple junction and is a
further manifestation of the extensional regime (Dickin-
son and Snyder, 1979; Johnson and O’Neil, 1984; Fox and
others, 1985).

The plate reconstructions of Atwater and Molnar
(1973) suggest an increase in relative motion between the
North American and Paciﬁc plates at about 10 Ma,
although this increase is expressed as a change in average
rates for the period of 21 to 10 Ma versus 10 to 4.5 Ma and
the change may have been a gradual one over several
million years. Page and Engebretson (1984) showed an
increase in slip rate at about 15 Ma, and Cox and
Engebretson (1985) inferred a small change in motion at
8.5 Ma, but nothing at 10 Ma. The differences may be
more apparent than real and may be simply a result of the
different methods used in the reconstructions. In any
case, there was an acceleration of the slip rate on the San

 

Andreas fault at 10 to 12 Ma (Huffman, 1972; Graham,
1978).

At about 5 Ma the motion of the Paciﬁc plate changed
to a more northerly direction, resulting in a component of
compression normal to the San Andreas transform (Min-
ster and Jordan, 1984; Page and Engebretson, 1984; Cox
and Engebretson, 1985). Opening of the Gulf of California
at about the same time, 5.5 Ma according to Moore and
Curray (1982), indicates increased coupling between the
sliver of former continental terrane west of the transform
and the Paciﬁc plate. This increased coupling resulted in
an acceleration of the slip rate on the San Andreas fault.

REGIONAL TECTONICS

The earliest Cenozoic tectonic events that affected the
region of the San Joaquin basin were probably related to
movements of the proto-San Andreas fault during the
Paleocene and possibly early Eocene. Right-lateral
strike-slip movement on the proto-San Andreas was
concurrent with oblique subduction and is generally
believed to have ended by about the end of the Paleocene
(Nilsen and Clarke, 1975; Dickinson and others, 1979).
Inasmuch as oblique subduction at the central California
margin continued until nearly the end of the Eocene,
strike-slip movement could conceivably have continued
well into the Eocene, although the evidence is equivocal.
A clockwise rotation of the southernmost Sierra Nevada,
demonstrated by paleomagnetic data, has been inferred
to reﬂect oroclinal bending due to right-lateral shear
along the proto-San Andreas fault (Kanter and
McWilliams, 1982; McWilliams and Li, 1985). This oro—
cline might also be considered a tectonic effect of the
accretion of the Tujunga terrane (part of the Salinia
composite terrane) to the North American craton in the
Mojave region near the end of the Paleocene (Howell and
others, 1987; Nilsen, 1987). The large folds at the
northwest end of the fold belt—that is, the Vallecitos
syncline, Joaquin Ridge anticline, and White Creek
syncline—are apparently of the right age and orientation
to have originated as an en echelon fold set associated
with right slip on the proto-San Andreas fault (Harding,
1976). Early Paleogene fold growth might also be con-
sidered an indication of thrusting associated with early
eastward movement of a Franciscan wedge at depth. The
evidence is insufﬁcient to make a deﬁnitive statement,
but both processes may have been operative.

The Stockton arch at the north end of the San Joaquin
basin has a more enigmatic origin. It has been suggested
that it formed by crustal buckling at the tectonic transi-
tion between a region of oblique subduction with proto-
San Andreas strike slip to the south, and a region of
oblique subduction without strike slip to the north
(Nilsen and Clarke, 1975; Dickinson and others, 1979).

14 THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

This origin is unlikely because of the absence of basement
arching (pl. 1, section D). The broad high of Cretaceous
rocks made apparent by later truncations is probably a
result of localized structural or sedimentary thickening of
the Cretaceous section. Structural thickening, although
possible, seems less likely because there is no evidence of
stratigraphic repetition by thrust faults or any other
deformation within the Cretaceous section. Sedimentary
thickening, on the other hand, seems more likely because
there is some evidence of thickness changes associated
with the arch. An isopach map of the Lathrop sand (of
local usage), for example, shows an abrupt northward
thinning across the trend of the Stockton fault (Hoffman,
1964). The total Cretaceous section is also thicker south
of the fault, even though some of the section has been
removed by post-Eocene erosion (pl. 1, section D). Well
data indicate that the basement is higher north of the
Stockton fault than it is to the south, under the arch
(Hoffman, 1964; Teitsworth, 1964; Bartow, 1983), and
thickness changes Within the Cretaceous section imply
down-to—the-south faulting concurrent with sedimenta-
tion. Inasmuch as Cenozoic movement on the fault has
been down to the north, there must have been a major
south-side-down offset of the basement surface during
the Cretaceous. The high basement on the north and the
increased Cretaceous thickness on the south can be
alternatively attributed to post-Cretaceous left-lateral
slip on the Stockton fault, but the amount of offset
required (several kilometers) makes this an unlikely
interpretation. Latest Cretaceous or Paleocene south-
side—up movement on the Stockton fault would reduce the
Cretaceous throw on the basement surface and raise the
Cretaceous rocks in the area to the south. The feature
that has come to be known as the Stockton arch might be,
then, simply an up—tilted fault block. Whatever its origin,
the Stockton arch area persisted as a positive element
throughout the remainder of the Cenozoic.

The Franciscan-Great Valley sequence fault boundary
(the Coast Range thrust) may have been reactivated in
the Tertiary, initially during the period from about 60 to
50 Ma when plate convergence was at a maximum (Page
and Engebretson, 1984). Cenozoic activity at this bound-
ary probably resulted from the emplacement of a wedge
of Franciscan between the Great Valley sequence above
and crystalline basement below, as proposed by Went—
worth and others (1984) (pl. 1, sections A, B).

Oligocene events in southern California had an effect
on the southernmost San Joaquin basin. A regional uplift
in the south, together with the formation of fault-
bounded alluvial basins, has been ascribed to the ap-
proach of the ancestral East Paciﬁc rise to the North
American plate and subduction of young, buoyant litho-
sphere somewhat in advance of the actual arrival of the
spreading ridge at the trench in mid-Oligocene time

 

(Nilsen, 1984; Crowell, 1987). The principal effect for the
San Joaquin basin seems to have been uplift and faulting
of the south end of the basin, particularly the San
Emigdio Mountains and southernmost Sierra Nevada
(Davis, 1983, 1986).

The evolution of extensional deformation in the Basin
and Range province probably also had an indirect effect
on the San Joaquin basin. The beginning of the exten-
sional stress regime in the Basin and Range during the
Oligocene is related to the evolution of the arc-trench
system. The transition from compression to intra-arc
extension and then to back-arc extension took place as
the subduction angle steepened and the eastern limit of
magmatism consequently migrated southwestward
(Eaton, 1979). The intra-arc and back-arc extension was
oriented at right angles to the trend of the trench (Zoback
and others, 1981) and may have produced compression in
the region (including the San J oquin basin) between the
Basin and Range province and the trench. '

Wrench tectonics in conjunction with deep-seated
thrusting along the southwest side of the basin adjacent
to the San Andreas fault dominated the Neogene. The
earliest conclusive evidence of en echelon folding appears
in the stratigraphic record near the end of Saucesian time
(about 16-17 Ma) (Harding, 1976). Additionally, the
provenance and distribution of Miocene sandstone in the
Temblor Range suggest that strike slip may have begun
by late early Miocene time on the central California
portion of the San Andreas fault (Graham and others,
1986). This timing is consistent with the fault offset
history of Huffman (1972, ﬁg. 13). Progressive basinward
expansion of the fold belt, together with cessation of
folding near the San Andreas while it continued farther
east, suggested to Harding (1976) that the folds and the
San Andreas fault were independent responses to a
diffuse coupling in the deep crust, and that the folds
propagated outward in an expanding deformational front.
The fact that younger folds, like the Kettleman Hills and
Lost Hills anticlines, are approximately parallel to the
San Andreas and not oblique to it indicates that they are
not purely a response to shear in the San Andreas
system. The basinward expansion of the Kettleman
Hills-Lost Hills part of the fold belt seems to be, in
contrast to the model proposed by Harding (1976), a
response to an eastward-advancing thrust front at depth
associated with the emplacement of a Franciscan wedge
at the base of the Great Valley sequence (Wentworth and
others, 1983) (pl. 1, sections A, B).

Extension in the Basin and Range province again had
an effect on the San Joaquin basin in the Miocene. Basin
and Range faulting began in the late Miocene, probably
about 10 Ma (Zoback and others, 1981), and probably
related left-lateral movement on the Garlock fault is
assumed to have begun at about the same time. This

MAJOR CONTROLS 0N SEDIMENTATION 15

extension resulted in the westward movement of the
Sierra Nevada block, carrying the San Joaquin basin with
it, and the consequent formation of the bend in the San
Andreas fault (Davis and Burchﬁel, 1973; Hill, 1982;
Bohannon and Howell, 1982). The space problem arising
from this westward movement probably caused compres-
sion normal to the San Andreas at the west side of the
Sierran block (Wentworth and Zoback, 1986).

The last major uplift of the Sierra Nevada is also
believed to have begun after 10 Ma (Christensen, 1966;
Huber, 1981), but the uplift and the westward movement
may not have been directly related. It has been sug-
gested that the late Cenozoic uplift of the Sierra Nevada
was caused by thermal thinning of the lithosphere after
northward passage of the Mendocino triple junction
(Crough and Thompson, 1977; Mavko and Thompson,
1983). The cold subducting slab north of the triple
junction insulated the overlying continental lithosphere,
whereas subduction had stopped south of the triple
junction which allowed the base of lithosphere to be
heated and converted to less-dense asthenosphere.

Acceleration in the slip rate on the San Andreas fault
in latest Miocene and Pliocene time correlates with an
increase in deformation in the fold belt adjacent to the
fault (Harding, 1976); this acceleration may have contrib-
uted to the rapid subsidence of the southern San Joaquin
basin (Dickinson and Snyder, 1979; Davis, 1983). Fault-
normal compression in Pliocene and Pleistocene time,
resulting from changes in relative plate motion at about
5 Ma, produced fault-parallel folds and reverse faults
(Zoback and others, 1987). This compression probably
caused uplift of the Temblor and Diablo Ranges (Enge-
bretson and others, 1985) and, together with the devel-
oping bend in the fault, was probably the principal factor
leading to northward-directed thrusting at the south end
of the basin.

The Cenozoic subsidence history of the San Joaquin
basin in relation to regional tectonics is not well known.
Inferences about subsidence have been made from esti-
mates of paleobathymetry and from the present depth
and basinward-thickening trends of individual strati-
graphic units, but because of the lack of precision in both
paleoecology and absolute age of the commonly used
benthic foraminiferal faunas, some margin of error exists
in reconstructions of subsidence history. Preliminary
attempts at geohistory analysis have been made by
Dickinson and others (1987), Moxon (1986), and Olson and
others (1986). Collectively, these studies suggest periods
of rapid subsidence in the (1) late Paleocene and earliest
Eocene, (2) middle Eocene, (3) latest Oligocene and early
Miocene, and (4) middle and late Miocene. Uplifts are
suggested in the Oligocene and near the early Miocene-
middle Miocene boundary. These generalizations are
based on preliminary geohistory analyses of widely

 

separated parts of the basin and, consequently, an event
identiﬁed from any one locality is not necessarily a
basinwide event. It should be noted that a rapid rise in
relative sea level can also result from steady basin
subsidence in combination with a eustatic rise in sea
level. Nevertheless, geohistory analysis is a promising
technique and will ultimately provide valuable informa-
tion on basin evolution. At the present time, however, a
detailed subsidence history has not been established for
the San Joaquin basin, nor is it possible to identify with
much certainty the speciﬁc causes of tectonic subsidence.

SEA-LEVEL CHANGE

Eustatic sea-level change can result from changes in
the volume of ocean water, changes in the volume of the
ocean basin, or possibly both. The processes that can
contribute to sea-level change were reviewed by Pittman
(1978) and Donovan and Jones (1979). Those that are
potentially most signiﬁcant in terms of rate and magni-
tude are (1) ﬂuctuations in continental ice sheets, (2)
changes in volume of the mid-ocean ridge system, and (3)
desiccation and ﬂooding of isolated ocean basins. An
additional process, suggested by Schlanger and others
(1981), is mid-plate thermal uplift and volcanism in ocean
basins. Ultimately, all processes affecting sea level can
be traced back, directly or indirectly, to global tectonics
(ﬁg. 2).

Sea-level change controls sedimentation by controlling
the environment of deposition. This is most obvious near
the strandline where change in sea level produces lateral
shifts of nonmarine and shallow-marine environments.
The stratigraphic record of sea-level change along an
active continental margin will probably be obscure be-
cause the effects of the prevailing tectonism will tend to
mask the relatively minor effects of sea-level change.
Without a global standard sea—level curve to which local
sections may be compared, it would be virtually impos-
sible 'to identify any particular event in the sedimentary
record as sea-level controlled. The coastal-onlap curves of
Vail and coworkers (Vail and others, 1977; Vail and
Hardenbol, 1979) and the eustatic sea-level curve of Haq
and others (1987) provide a standard for comparison. The
eustatic sea-level curve (Haq and others, 1987) used in
this report (pl. 2) has been adjusted to ﬁt the Cenozoic
chronology of Berggren and others (1985). Although the
basic concept of eustatic ﬂuctuation of sea level through
geologic time is generally accepted, questions have been
raised about the actual inﬂuence of sea level on the
stratigraphic record and about the validity of the Vail
sea—level model (Watts, 1982; Parkinson and Summer-
hayes, 1985; Miall, 1986). The Vail model is still being
tested and revised, but some results to date suggest that
it may be accurate enough, if used with caution, to serve

16 THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

as a useful tool in interpreting continental margin history
(May and others, 1984; Poag, 1984; Poag and Schlee,
1984; Aubry, 1985; Poag and Ward, 1987).

The coastal-onlap curve of Vail and others (1977) was
originally equated directly with relative change of sea
level. The asymmetric shape of the curve, representing
long periods of sea-level rise followed by apparently
instantaneous fall, provoked some controversy. It is now
recognized that the coastal-onlap curve does not equate
directly with global change in sea level, but the lows in
the sawtooth pattern, representing downward shifts of
coastal onlap, mark times of global unconformities. These
unconformities correlate with inﬂection points on the
eustatic sea—level curve during periods of falling sea
level, so that the lowstand of sea level actually comes
shortly after the downward shift in coastal onlap (Vail
and others, 1984). Unconformity—bounded sequences of
local sections can then be compared with the eustatic
curve by correlating unconformities with the inﬂection
points on the curve. Correlation of unconformities be-
tween the global standard and local sections is evidence
that supports sea-level control, but it does not prove
exclusive sea-level control because global tectonism can
affect sea level as well as local tectonic activity.

CLIMATE

With the exception of glacial and periglacial environ-
ments of the Ice Ages, climate is the least inﬂuential of

 

/ GLOBAL TECTONICS

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

/
// l
/ / / ¢ Epei rogeny
/ / II \
// / I Subduction \\
/ / l V \
/ I, ' M.O.R. \
/ l “ volume 0 rogeny, \\
I l . isostasy \
I \ Geoudal |
l \ va riation 4 l
I \ , I
l‘ ‘ l/ I,
\ Basin Ocean-basin / /
\\ desiccation volume // ll
\
/
/ Ma ri ne
Glaciation // sedimentation

 

 

 

 

\ i //

EUSTATIC SEA LEVEL

 

 

 

 

FIGURE 2. —Interrelation of factors causing eustatic ﬂuctuations of sea
level. The major controls on volume changes of the ocean basins and
of the global hydrosphere are enclosed in boxes. Ultimately, global
tectonics is the overriding inﬂuence, either directly (solid lines) or
indirectly (dashed lines). M.O.R., mid-ocean ridge. From May and
others (1984).

 

the controls on sedimentation. The climatic inﬂuence is
more likely to be reﬂected in the character of the
sediments, that is, facies or mineral composition, than it
is in the control of transgressions or regressions.

The Tertiary was a period of general climatic deterio-
ration from the very warm global climates of the Late
Cretaceous to the cool glacial climate of the Quaternary.
Climates were warmest during the Eocene, when a low
latitudinal temperature gradient and high precipitation
prevailed (Frakes, 1979; Wolfe, 1978). A rapid deterio-
ration near the Eocene-Oligocene boundary led to glacial
conditions in Antarctica during the Oligocene (Frakes,
1979; Mathews, 1984) and generally cooler, dryer global
climates. Temperatures warmed somewhat during the
late Oligocene, early to middle Miocene, and latest
Miocene, but at no time were temperatures as warm as
those of the Eocene (Wolfe and Hopkins, 1967; Addicott,
1970). The climate ﬂuctuated in the Neogene, but the
overall trend was toward cooler temperatures. Sea-level
glaciation in the Arctic and alpine glaciation in the Sierra
Nevada both date from the late Pliocene (Frakes, 1979).

In the San Joaquin basin, the warm, wet tropical
climate of the Eocene is reﬂected in the quartz-kaolinite
sandstone and lignite, along with the associated laterites,
of the Eocene deposits. Outwash from alpine glaciers in
the Sierra Nevada contributed to the San Joaquin basin
alluvial sedimentation beginning in late Pliocene time;
the pluvial climate of the Pleistocene resulted in a series
of large lakes in the San Joaquin Valley.

THE SEDIMENTARY RECORD

The early Cenozoic sedimentary history of the San
Joaquin basin is fundamentally different from that of the
later Cenozoic, as can be seen on plate 2. The early
Cenozoic is characterized by a few long-lasting basinwide
depositional sequences, whereas the later Cenozoic is
characterized by shorter sequences of more local extent.
The change took place during the Oligocene at about the
beginning of the transition from a convergent continental
margin to a transform margin. This fundamental change
in sedimentation patterns seems to show a clear correla-
tion to the change in continental margin tectonics, but
other more localized sedimentary events are not always
easily analyzed. .

PALEOGENE
GREAT VALLEY SEQUENCE

The Upper Cretaceous and lower Tertiary part of the
Great Valley sequence represents the last phase of
deposition in the late Mesozoic to early Tertiary forearc
basin. It consists principally of deep-sea fan deposits and
associated facies on the west side of the basin and

THE SEDIMENTARY RECORD 17

shallow-marine to deltaic deposits on the east side
(Ingersoll, 1979; Cherven, 1983). The uppermost or lower
Tertiary part of the sequence consists mostly of west-
ward-prograding slope, shelf, and deltaic facies (Cher-
ven, 1983). The Great Valley sequence is apparently
absent south of the Bakersﬁeld arch, as noted earlier.
This absence may be due to uplift associated with the
oroclinal bending of the southernmost Sierra Nevada in
the early Tertiary.

The Great Valley sequence is separated from the
overlying upper Paleocene and lower Eocene sequence,
consisting principally of the Tesla and Lodo Formations,
by an unconformity (pl. 2). In the northern Diablo Range
at the west end of the Stockton arch, there is a slight
angular discordance at the unconformity, whereas far-
ther south on the west ﬂank of the northern Diablo
Range, the Moreno and Tesla Formations are concord-
ant. The contact between the Moreno and Laguna Seca
Formations appears to be gradational in exposures just
south of Los Banos (Briggs, 1953), but still farther south
at the type area of the Lodo Formation, there is
paleontologic evidence of a hiatus (Berggren and Aubert,
1983); in the Vallecitos syncline there is evidence of an
unconformity with westward overlap of the Lodo For-
mation on the Great Valley sequence (White, 1940;
Dibblee, 1979a). Despite apparent conformity locally,
there is probably at least a disconformity at the contact
throughout the basin, and the base of the overlying
sequence is transgressive.

An angular unconformity at the top of the Great Valley
sequence was probably produced in the northern part of
the basin by uplift of the Stockton arch during the
Paleocene. In the central and possibly southern Diablo
Range, the unconformity was probably a result of either
concurrent en echelon folding associated with movement

on the pr0t0~San Andreas fault (Harding, 1976) or.

thrusting associated with emplacement of a Franciscan
wedge. The angular discordance is always slight or
nonexistent, however, and indicates only mild deforma-
tion.

The unconformity between the Great Valley sequence
and the upper Paleocene and lower Eocene sequence
correlates with the boundary between global supercycles
“TA1” and “TAZ” in the sequence stratigraphy of Haq
and others (1987) (pl. 2). The good correlation suggests
that sea—level change was a major contributing factor to
the regression that produced the unconformity. The
widespread nature of the unconformity in the San J oa—
quin basin and the absence of evidence for strong
deformation also support this interpretation.

UPPER PALEOCENE AND LOWER EOCENE SEQUENCE

The deposits of the upper Paleocene and lower Eocene
sequence show a prominent shoaling trend northward
onto the Stockton arch (Nilsen and Clarke, 1975). The

 

Lodo Formation in the Vallecitos syncline area, and to
the southeast, consists of thin neritic deposits at the base
that are overlain by middle or lower bathyal deposits
(Berggren and Aubert, 1983). In the northern part of the
basin, the equivalent units consist of neritic or shallower
facies that include common nearshore marine and ﬂuvial-
deltaic facies ﬂanking the Stockton arch (Dickinson and
others, 1979). The amount of deepening during this rapid
transgression in the southern part of the basin was 400 m
or more (Berggren and Aubert, 1983). This is more than
could be accounted for by a sea-level rise alone and
indicates a pronounced southward or southeastward tilt
of the basin during latest Paleocene or early Eocene time.

Deposition of the upper Paleocene and lower Eocene
sequence ended after a regression that culminated in a
widespread unconformity (pl. 2). The overlying Eocene
sequence, consisting of the Domengine Sandstone, the
Kreyenhagen Shale, and their correlatives, rests on
rocks as old as Cretaceous in the southern Diablo and
northern Temblor Ranges (Dibblee, 1973a) and laps onto
Sierran basement rocks on the east side of the basin
(Bartow, 1985). Although the sequence is not in deposi-
tional contact With Franciscan rocks, Franciscan detri-
tus, in the form of red and green radiolarian chert pebbles
or glaucophane, is present in basal sandstones of the
sequence in the southern Diablo Range (Nilsen and
Clarke, 1975; Dickinson and others, 1979). The uncon-
formity seems to be present throughout the northern
part of the basin, as evidenced by the local truncation of
lower Eocene and even Paleocene rocks (Church and
Krammes, 1958; Bartow, 1985). However, from the
Kettleman Hills southeastward, the upper Paleocene and
lower Eocene and the overlying Eocene sequences ap-
pear to be conformable (Church and Krammes, 1959).
There are virtually no data on rocks older than late
Eocene in the southern Temblor Range. The offset
equivalent of these older rocks on the opposite side of the
San Andreas fault, however, now lies in the Santa Cruz
Mountains more than 300 km to the northwest (Clarke
and Nilsen, 1973). The deep-sea fan deposits of the offset
Point of Rocks Sandstone Member of the Kreyenhagen
Shale (San Joaquin basin) and Butano Formation (Santa
Cruz Mountains) span the time represented by the
regression, but indicate continued bathyal sedimentation
in the southwestern part of the basin (Clarke, 1973;
Nilsen and Clarke, 1975) and in its western extension
(Nilsen and Clarke, 1975; Stanley, 1985).

The unconformity between the upper Paleocene and
lower Eocene sequence and the overlying Eocene se-
quence correlates with the boundary between the “TA2”
and “TA3” global supercycles of Haq and others (1987)
(pl. 2). Similarly correlative unconformities have been
recorded in several areas on the Paciﬁc coast (Berggren
and Aubert, 1983; May and others, 1984), in Libya (Barr
and Berggren, 1981), and in several areas of Europe
(Aubry, 1985). The global nature of this unconformity is

18 THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

strong evidence for lowered sea level at the end of the
early Eocene. However, an angular discordance at the
unconformity and Franciscan detritus above the uncon-
formity provide convincing evidence of concurrent tec-
tonic activity in central California. Coarse Franciscan
detritus appears to be concentrated in a belt that lies in
the southern Diablo Range as far north as the Vallecitos
syncline. Assuming 305 km of Neogene offset on the San
Andreas fault (Clarke and Nilsen, 1973) and 115 km on
the San Gregorio-Hosgri fault trend (Graham and Dick-
inson, 1978), the belt would lie adjacent to the inferred
Paleogene position of the north end of the Salinia terrane.
If 150 km of offset is assumed for the San Gregorio-
Hosgri fault (Clark and others, 1984), the belt would lie
just north of the inferred north end of the Salinia terrane.
This location of the belt suggests a relation between uplift
of the Franciscan Complex and emplacement of the
Salinia terrane. Franciscan source areas may have been
uplifted following emplacement of the Salinia terrane to
the west, as suggested by Dickinson and others (1979).
Inasmuch as emplacement of the Salinia terrane opposite
the southern San Joaquin basin in the early Tertiary must
have involved strike slip on the proto-San Andreas or
closely related faults, the uplift in the area of the Diablo
Range may, then, have been due to related wrench
tectonism that continued into the early Eocene. The role
of wrench tectonism at that time is, however, question-
able, and the uplift might better be considered as
evidence of early Tertiary thrust emplacement of a
Franciscan wedge at depth under the southern Diablo
Range, just north of the north end of the Salinia terrane.
Whatever the tectonic mechanism, the regression at the
end of the early Eocene was probably produced by a
combination of eustatic sea-level change and local uplift.

EOCENE SEQUENCE

Deposition of the Eocene sequence began during a
rapid transgression. A minor hiatus near the base of the
sequence (Milam, 1985) is associated with a condensed
section, as evidenced by abundant glauconite and low
sedimentation rates in the basal Kreyenhagen Shale.
Milam (1985) calculated a sedimentation rate of less than
1 cm/1,000 years for the lower part of the Kreyenhagen.
The hiatus and condensed section are results of rapid rise
of relative sea level that produced starved—basin-type
sedimentation and was probably caused by rising global
sea level combined with basin subsidence. Deposition of
the bathyal shale that makes up most of the sequence
corresponds to a highstand of sea level on the eustatic
sea-level curve of Haq and others (1987).

A regression at the end of the Eocene is apparent
mostly on the ﬂanks of the Stockton arch in the northern

 

part of the basin and in the San Emigdio Mountains at the
south end of the basin (pl. 2). A minor reversal of the
overall Paleogene transgressive trend along the south-
eastern margin of the basin is evidenced by intertonguing
nonmarine and shallow-marine deposits (Bartow and
McDougall, 1984). Elsewhere in the southern part of the
basin, deep-marine sedimentation prevailed into the
Oligocene with only slight to moderate shallowing; an
unconformity in the Temblor Range area (pre-Temblor
Formation) was interpreted by Carter (1985) as a result
of submarine erosion at bathyal depths. Evidence of
tectonic activity near the Eocene-Oligocene boundary
appears in the Poverty Flat Sandstone along the east
ﬂank of the Diablo Range in the northwest, and in the
Tecuya and Pleito Formations in the south. The Poverty
Flat contains a conglomerate that consists principally of
red radiolarian-chert pebbles derived from the Fran-
ciscan Complex and a few other pebbles of Franciscan
and ophiolite lithologies (Bartow and others, 1985). This
conglomerate represents the earliest appearance of ap-
preciable coarse Franciscan detritus in the northern
Diablo Range. Lenses of granitic breccia at the base of
the Tecuya (Nilsen and others, 1973) and in the lower
part of the Pleito in the San Emigdio Mountains were
interpreted by DeCelles (1986) as a result of a seismically
triggered rockslide and associated submarine mass move-
ments. Syndepositional deformation structures in coeval
sediments were ascribed to the results of seismic shak-
ing. This inferred seismic event may mark the beginning
of Oligocene tectonic activity at the south end of the
basin.

Uplift of a Franciscan source area in the northern
Diablo Range at the end of the Eocene was probably, as
was the earlier uplift in the southern Diablo Range, a
result of thrust emplacement of a Franciscan wedge at
depth. This is, in turn, inferred to be a subduction-related
process. A change from oblique to normal subduction is
believed to have taken place near the end of the Eocene,
although there was a net decrease in the normal compo-
nent of plate convergence through the Oligocene and a
consequent steepening of the angle of subduction (Page
and Engebretson, 1984). The approach of the spreading
ridge to the continental margin (Engebretson and others,
1985) would have resulted in the subduction of young,
buoyant lithosphere. No deﬁnite cause and effect rela-
tions can be established, but it seems probable that
regional tectonic activity from the latest Eocene into the
Oligocene was related to these plate tectonic events.
Furthermore, Paleogene deformation in the northern
Diablo Range seems to be consistent with subduction-
related tectonics.

The late Eocene regression is, coincidentally, approxi-
mately correlative with an interval of lowered sea level in
the lower part of supercycle “TA4” of Haq and others

THE SEDIMENTARY RECORD 19

(1987) (pl. 2). The pattern of the regression—that is,
strongly developed in the north and south, weakly
developed on the southeast and west margins, and
apparently absent in the deeper parts of the basin—sug-
gests that the local effects of tectonism may have been
augmented by a fall in sea level. Tectonism was dominant
in the north, in the area of the Diablo Range, and possibly
at the south end of the basin. A minor regression is
recorded elsewhere along the basin margins where the
fall of sea level outpaced basin subsidence, but the record
in much of the deeper parts of the basin is not clear.

LOWER OLIGOCENE SEQUENCE

The lower Oligocene sequence is present only in the
southern San Joaquin basin. Along the southwest side of
the basin, deep-marine deposition continued from the
Eocene into the Oligocene with only minor shallowing
and the two sequences are not clearly separated. At the
south end of the basin, deep-marine deposition resumed
after the late Eocene regression, while in the north, an
extensive hiatus indicates that the Stockton arch re-
mained a positive area through most of the Oligocene (pl.
2). Eocene and Paleocene units are truncated over the
crest of the arch, and later Tertiary nonmarine strata
rest directly on Cretaceous rocks (Church and Krammes,
1958). The sea also withdrew from most of the Sacra-
mento basin at the end of the Eocene, leaving only a
narrow embayment occupying the former Markeley sub-
marine canyon (Almgren, 1978). Extensive marine de-
position was restricted to the southern part of the San
Joaquin basin from the Oligocene through the Pliocene.
Continued mild uplift of the northern San Joaquin basin
and concurrent reverse movement on the Stockton fault
during the Oligocene, while subsidence continued in the
southern part of San Joaquin basin, was probably a
continuation of the subduction-related tectonism that
began in the late Eocene.

Oligocene deep-marine deposition in the southern San
Joaquin basin was interrupted at mid-Oligocene time by
a regression (pl. 2). Evidence for the regression is best
developed in the western San Emigdio Mountains at the
extreme southwest end of the basin where upper Oligo-
cene rocks of the Temblor Formation overlap Eocene
rocks to lie directly on the basement (Nilsen and others,
1973; Lagoe, 1986). Along the west-side fold belt, a thin
shallow—water sandstone unit, the Wygal Sandstone
Member of the Temblor Formation, occurs within a
deep-water shale section (Addicott, 1973; Carter, 1985).
A slight angular unconformity is present at the base of
the Wygal locally in the fold belt, but the units may be
conformable farther east and south (Harding, 1976;
Carter, 1985). Evidence for a mid-Oligocene regression

 

elsewhere in the basin is somewhat questionable, al-
though a major mid-Oligocene unconformity is inferred
on the margins of the La Honda basin (Stanley, 1985), the
offset western continuation of the southern San Joaquin
basin.

The mid-Oligocene regression appears to be approxi-
mately coincident with a major lowering of sea level at
about 29 to 30 Ma (Haq and others, 1987) (pl. 2). The
amount of relative sea-level change required in the San
Joaquin basin, however—a shallowing from middle bath-
yal to inner neritic depths (1,500—2,000 m)—-is far too
much to be accounted for by eustatic sea-level change
alone. The encounter of the Paciﬁc-Farallon ridge with
the North American continental margin was also at about
29—30 Ma (Atwater, 1970; Atwater and Molnar, 1973).
This is, as Stanley (1985, p. 11) termed it, a “cruel
coincidence between major eustatic and tectonic events.”
Uplift of southern California in response to the approach
of the spreading ridge, as proposed by Nilsen (1984), may
have been felt as far north as the southernmost end of the
basin, where it would have caused the uplift of the San
Emigdio Mountains area. Farther north, the ridge itself
would not have had an effect, but the subduction of
young, buoyant lithosphere in the east ﬂank of the
approaching ridge might have. The regional tectonic
framework in which the uplift took place includes, in
addition to the approach of the spreading ridge, a
steepening of the angle of subduction of the Farallon
plate and consequent southwestward migration of the arc
volcanism into Nevada (Snyder and others, 1976; Cross
and Pilger, 1978), as well as the onset of extensional
deformation in the Basin and Range (Zoback and others,
1981). Broad uplift of the region between the trench and
the westward-advancing volcanism and Basin and Range
extension might, then, be considered a subduction-
related event, but the exact cause of the abrupt uplift at
the west margin of the San Joaquin basin is not known.
It may have been a continuation of uplifts that started at
the end of the Eocene, which were probably related to
the subduction of young lithosphere and to the change
from oblique to normal subduction. If the mid-Oligocene
regression was basinwide, as it appears, and not just
restricted to the south and west margins, it may be due
to a combination of eustatic sea-level change and local
tectonism.

UPPER OLIGOCENE SEQUENCE

The upper Oligocene sequence is also restricted to the
southern San Joaquin basin, although alluvial sedimen-
tation of the Valley Springs Formation began in the
northern part of the basin in the late Oligocene. The
sequence in the Kettleman Hills area is atypical (pl. 2) in

20 THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

that sedimentation continued into the Miocene after a
minor reversal of the regressive trend at the Oligocene-
Miocene boundary.

A regression near the end of the Oligocene produced
unconformities around the northwest, north, and east
margins of the basin and a shoaling at the south end. Part
of the Coalinga-Kettleman Hills area remained high well
into the early Miocene (Kuespert, 1983, 1985). An uncon-
formity at the base of the Agua Sandstone Bed of the
Santos Shale Member of the Temblor Formation3 trun-
cates older units in the northern Temblor Range (Heik-
kila and MacLeod, 1951; Carter, 1985; Pence, 1985), but
the units become conformable farther southeast and the
Agua pinches out. Where present, the Agua indicates a
shallowing, similar to that of the Wygal Sandstone
Member, from bathyal to neritic depths (Carter, 1985).
In the Bakersﬁeld arch area, the deep-water Vedder
Sand is disconformably overlain by shallow-water basal
deposits of the Jewett Sand (Bartow and McDougall,
1984).

There is apparently no major eustatic sea-level event
that correlates with this ﬁnal Paleogene regression,
although a lowstand at about 25 Ma (Haq and others,
1987) is close (pl. 2). The amount of relative sea-level
change that took place would rule out eustasy as a
primary cause anyway. The encounter of the Paciﬁc-
Farallon ridge with the North American margin had, by
late Oligocene time, resulted in the formation of a triple
junction at the continental margin. However, this triple
junction was still located off southern California (Atwa-
ter, 1970), and triple-junction tectonism should not have
affected the San Joaquin basin at that time. Uplift near
the end of the Oligocene at the south end of the basin was
probably a continuation of the southern California uplift;
elsewhere, it may have been a continuation of seemingly
subduction related tectonism that began earlier in the
Oligocene. This tectonism seems to have had the most
pronounced uplift effects on the west or northwest side of
the basin and lesser effects elsewhere.

An indirect effect of the regional tectonism and asso-
ciated volcanism in central Nevada that began in the
Oligocene was the deposition of rhyolitic ash-ﬂow and
air-fall tuffs in the northern Sierra Nevada and adjacent
northern San Joaquin basin (Slemmons, 1966). Rare tuffs
date from before 30 Ma, but more widespread pyroclastic
deposition in the latest Oligocene (Dalrymple, 1964) was
concurrent with the beginning of alluvial sedimentation
(Valley Springs Formation) that continued into the
Miocene and may indicate uplift east of the basin.

3This awkward name was introduced into the literature by Maher and others (1975); it is little
used.

 

NEOGENE AND QUATERNARY

LOWER AND MIDDLE MIOCENE SEQUENCE

The regression near the end of the Oligocene was
followed by a marine transgression in the southern part
of the basin that began near the Oligocene-Miocene
boundary and brought about a rapid return to the bathyal
or abyssal depths of the Oligocene (Bandy and Arnal,
1969). Lavas, mostly of basaltic and andesitic composition
but including minor dacitic lava, erupted at the southeast
end of the basin in the Tehachapi-San Emigdio Mountains
area at about 22—23 Ma and ﬂowed westward or north-
westward across the strandline (Nilsen and others, 1973).
Volcanism and extensional tectonism also characterized
the late Oligocene-early Miocene history of the western
extension of the San Joaquin basin in the Santa Cruz
Mountains (Stanley, 1985). In the northern part of the
San Joaquin basin, extensive alluvial sedimentation of
the Valley Springs Formation continued into the Mio-
cene.

The apparent coincidence between extensional tecton-
ism, as shown by basin subsidence and volcanism, and
passage of the Mendocino triple junction has been noted
in the section on “Plate Tectonics.” The northward
progression of volcanism and its association with triple-
junction migration, of which the volcanic rocks in the San
Emigdio Mountains are one element, has been well
documented by Johnson and O’Neil (1984). The case is not
so clear for basin subsidence, however, because the data
of Bandy and Arnal (1969) indicate that the south end of
the basin was as deep during the Oligocene, before
passage of the triple junction, as it was in the Miocene,
after passage of the triple junction. Nevertheless, geo-
history analysis (Olson and others, 1986) indicates that
the basin subsided rapidly after the brief regression that
intervened between the late Oligocene and early Miocene
deep-basin intervals. Extensional tectonism is demon-
strated by a latest Oligocene and early Miocene episode
of normal faulting recorded in the San Emigdio Moun-
tains (Davis, 1986; Hirst, 1986). The most likely expla-
nation for extensional tectonism and basin subsidence
beginning about 24 Ma at the s0uth end of the basin is the
passage of the Mendocino triple junction that created an
extensional stress regime in its wake (Dickinson and
Snyder, 1979; Ingersoll, 1982). Alluvial sedimentation of
the Valley Springs Formation, however, is too old and
too far north to have been related to triple-junction
passage and is more probably a result of increased
sediment supply due to uplift in the source area.

The Valley Springs Formation in the northern San
Joaquin Valley, and its associated rhyolitic pyroclastic
deposits, was succeeded in the middle Miocene by exten-
sive andesitic volcaniclastic sediments of the Mehrten

THE SEDIMENTAR‘Y RECORD 21

Formation. There is no apparent angular discordance
between the two units (Marchand, 1977; Grant and
others, 1977), although there is an unconformity with as
much as 120 m of erosional relief in the eastern part of the
outcrop area (Gale and others, 1939). Although there was
obviously no major tilting of the northwest edge of the
basin at this time, there may have been regional uplift of
the Sierra Nevada and the San Joaquin basin without a
southwest tilt. A minor regressive pulse, evidenced by a
nonmarine tongue (the lower variegated unit) in the
Temblor Formation, can be seen on the west side of the
basin as far south as Kettleman Hills north dome (pl. 2).
This regressive pulse suggests a regional uplift of the
northern part of the basin.

Deposition of the lower and middle Miocene sequence
ended after a regression that is apparent along the
southeast margin of the basin and in the northern
Temblor Range segment of the west-side fold belt (pl. 2).
The Olcese Sand forms a clastic wedge of shallow-marine
and nonmarine sandstone between the deeper water
Freeman and Round Mountain Silts in the Bakersﬁeld
arch area (Addicott, 1970; Bartow and McDougall, 1984).
Abrupt changes in foraminiferal faunas suggest the
presence of a disconformity, at least locally, within the
Olcese near Bakersﬁeld (Bartow and McDougall, 1984).
A nonmarine conglomerate unconformably overlies older
Miocene rocks in the Tehachapi and San Emigdio Moun-
tains area (Nilsen and others, 1973; Bartow and McDou-
gall, 1984). In the northern Temblor Range area, an
unconformity truncates early Miocene and older rocks
over fold crests and is overlain by a shallow-marine
sandstone, the Relizian-age Buttonbed Sandstone Mem-
ber of the Temblor Formation (Dibblee, 1973b; Harding,
1976). The Buttonbed is absent farther southeast, but an
unconformity is present between Saucesian and Relizian-
age units in structures along the west side of the San
Joaquin Valley (Harding, 1976). Still farther south in the
southern Temblor Range, the lower Miocene is generally
too deeply buried to determine the presence or absence of
an unconformity with any conﬁdence; however, in the
offset western extension of the southwestern San Joa-
quin basin in the Santa Cruz Mountains, there is a
widespread unconformity at the top of the Saucesian
(Stanley, 1985).

The unconformity at the top of the lower and middle
Miocene sequence provides conclusive evidence of the
earliest en echelon folding along the southwest margin of
the basin (Harding, 1976) that can be stratigraphically
dated at a minimum age of 16—17 Ma (pl. 2). The west-side
folding is probably a result of the initial stages of San
Andreas wrench faulting. The unconformity in the Santa
Cruz Mountains area, together with the regressive
elastic wedge on the east side of the San Joaquin basin
and the unconformity in the Tehachapi-San Emigdio

 

area, provides evidence of regional tectonism (Olson and
others, 1986; Dickinson and others, 1987) that probably
cannot all be ascribed to wrench tectonism. It has been
suggested that this early middle Miocene uplift was an
isostatic response to the northward movement of the
Mendocino fracture zone, which marked the northern
edge of a slab of young, buoyant lithosphere, under the
basin (Loomis and Glazner, 1986). This is an appealing
idea, but it requires a shallow subduction angle so that
the subducted plate is in contact with the overlying
lithosphere under the basin and can, therefore, affect the
overlying plate isostatically. The distribution and timing
of magmatism in the western Cordillera indicates much
steeper subduction by the middle Miocene (Coney and
Reynolds, 1977; Keith, 1978). Paleomagnetic data indi-
cate that much of the clockwise rotation of the Tehachapi
Mountains occurred in the early Miocene (Plescia and
Calderone, 1986), presumably under right-lateral shear
stress after passage of the Mendocino triple junction. To
what extent this may have inﬂuenced uplift near the end
of the early Miocene is not really known. The problem
remains unresolved.

MIDDLE AND UPPER MIOCENE SEQUENCE

Deposition of the middle and upper Miocene sequence
began during a rapid subsidence. Shallow-marine trans-
gressive sandstone was deposited at the base (Buttonbed
Sandstone Member of the Temblor Formation on the
west side and the upper part of the Olcese Sand on the
east side). Sporadic inﬂuxes of coarse clastic sediments at
the southeast, south, and west basin margins fed deep-
sea fan systems (Stevens sandstone of local usage) in the
deep-basin areas (MacPherson, 1978; Webb, 1981). The
sequence is characterized, however, by its thick accumu-
lation of ﬁne-grained siliceous sediment (mostly included
in the Monterey Formation). The transgression reached
its greatest areal extent at about mid-Mohnian time,
resulting in broad, shallow shelves along the north and
east basin margins (Graham and others, 1982). A thin
marine unit in the subsurface at the Chowchilla gas ﬁeld
(about 70 km northwest of Fresno) has yielded late
middle Miocene(?) diatoms (J .A. Barron, written com-
mun., 1986). This unit is the northernmost marine
Miocene deposit in the San Joaquin basin and very
probably represents the approximate northern limit of
the middle Miocene transgression. The siliceous lithology
of the Monterey-type sediments and their high compo-
nent of pelagic organisms are due to a combination of
broad shelves that caused terrigenous sediment to be
trapped in shallow estuaries or lagoons (Graham and
others, 1982) and proliﬁc middle Miocene diatom produc-
tivity caused by changes in oceanic circulation (Ingle,
1981; Barron, 1986).

22 THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

The middle Miocene transgression approximately co-
incides with a eustatic highstand of sea level shown by
Haq and others (1987). High sea level may well be the
primary factor in the widespread transgression, but the
basin also appears to have reached its maximum depth at
progressively later times northward (pl. 2). This north-
ward progression suggests that tectonism also played a
part and that the northward migration of the Mendocino
triple junction may have induced a northward—moving
wave of basin subsidence in the San Joaquin basin.

Beginning near the end of the middle Miocene, nonma-
rine, coarse clastic sediments derived entirely from Coast
Range sources were deposited along the northwest side
of the San Joaquin basin. These alluvial fan deposits are
included in the Oro Lorna Formation (Briggs, 1953) and
the Carbona unit of Raymond (1969); the latter was
mapped as late Miocene and early Pliocene(?) fanglomer-
ate by Bartow and others (1985) (pl. 2). The base of the
Carbona unit is dated by vertebrate fossils at 10 Ma
(Raymond, 1969; Bartow and others, 1985), but Lettis
(1982) suggested that the Oro Loma might be older
because it apparently does not contain detritus from the
nearby Quien Sabe Volcanics, which range in age from
10.7 to 7.5 Ma (Prowell, 1974; Drinkwater, 1983). Nev-
ertheless, because the Oro Loma and the Carbona have
very similar lithologies and are in the same stratigraphic
position on the east ﬂank of the Diablo Range, they
probably record the same tectonic event. This event must
have begun at, or shortly before, 10 Ma and indicates the
beginning of Neogene uplift of the Diablo Range and
exposure of the Franciscan core. The uplift seems to have
begun in advance of the northward-migrating Mendocino
triple junction, using either the global-circuit reconstruc—
tions of Atwater and Molnar (1973) or the hot-spot
method of Engebretson and others (1985). The eruption
of the Quien Sabe Volcanics, on the other hand, was
approximately coincident with the passage of the triple
junction and is one further element in the northward
progression of volcanic events (Dickinson and Snyder,
1979; Johnson and O’Neil, 1984).

Deposition of the middle and upper Miocene sequence

A was interrupted along the southeast margin of the basin
by a brief regression near the middle Miocene-late
Miocene boundary. An unconformity at the base of the
nonmarine Chanac Formation, or its marine equivalent
the Santa Margarita Formation, truncates older Miocene
units (Bartow and McDougall, 1984). This unconformity
indicates a local uplift of the southern Sierra Nevada and
Tehachapi Mountains that may have been the ﬁrst pulse
of the gradually accelerating late Cenozoic uplift of the
Sierra Nevada. Hay (1976) ﬁrst proposed that late
Cenozoic uplift of the Sierra Nevada began earlier in the
south and he also pointed out the relation between the
timing of that uplift and the evolution of the San Andreas

 

transform, but it was Crough and Thompson (1977) who
proposed what is probably the most likely mechanism for
the uplift. That mechanism was thermal thinning of the
lithosphere following the northward migration of the
Mendocino triple junction and the end of subduction (see
section on “Regional Tectonics”).

The regressive phase of the middle and upper Miocene
sequence was characterized by an increase in terrigenous
material, along with a shallowing of the basin. Shallow-
marine sandstone (Santa Margarita Formation) was
deposited on the north and east side of the restricted
marine basin and nonmarine sediments (Chanac Forma-
tion) were deposited on the southeast side. The middle
and upper Miocene sequence is separated from the
overlying upper Miocene, Pliocene, and P1eistoCene se-
quence by an unconformity around the margins of the
southern San Joaquin basin, but deposition was appar-
ently continuous in the center of the basin. A local
unconformity Within the nonmarine Mehrten Formation
along the northeast margin of the basin (Wagner, 1981)
may correlate with the unconformity in the marine
section farther south.

The late Miocene regression in the San Joaquin basin
was approximately synchronous with tectonic events in
the surrounding regions that are not precisely dated, but
that seem to cluster at about 10 Ma. The regression also
correlates with a pronounced fall in sea level at 10 to 11
Ma on the Haq and others’ (1987) eustatic sea-level curve
(pl. 2). In the Basin and Range province, there was a
clockwise change in the direction of least-principal stress
(from west-southwest—east-northeast to west-north-
west—east-southeast) at about 10 Ma, which is consistent
with the superposition of right-lateral shear in that
region (Zoback and others, 1981). An apparently syn—
chronous event was the accelerating uplift of the Sierra
Nevada (Christensen, 1966; Huber, 1981). Folding along
the southwest side of the San Joaquin basin, on the other
hand, has been apparently continuous since the early
middle Miocene; there has been no particular change in
rate, merely a shift from one structure to another in a
general basinward progression (Harding, 1976).

The cause of tectonic activity at about 10 Ma is unclear,
partly because the possible changes in plate motion at
that time are also unclear, as discussed in the section on
“Plate Tectonics.” An acceleration in Paciﬁc—North
American relative plate motion might well have been an
inﬂuence, but the imprecision in the timing of that change
makes it difﬁcult to relate it to speciﬁc tectonic events in
central California. Regardless of plate motions, however,
there was an acceleration of slip rates on the San Andreas
fault through the Miocene, particularly the late Miocene
(Huffman, 1972; Graham, 1978), probably due to in-
creased coupling between the Salinia terrane west of the
fault and the Paciﬁc plate. A resulting gradual increase in

THE SEDIMENTARY RECORD 23

tectonic activity throughout the region may have initi-
ated a regression, but it is probable that the fall in sea
level at 10 to 11 Ma was a strong contributing factor.

UPPER MIOCENE, PLIOCENE, AND
PLEISTOCENE SEQUENCE

Deposition of the upper Miocene, Pliocene, and Pleis—
tocene sequence in the southern San Joaquin basin was
started with transgression of the Etchegoin Formation
over older Miocene units. Alluvial fans and deltas pro-
graded basinward as abundant coarse detritus was de—
livered to the basin from the rising Sierra Nevada on the
east, the San Emigdio Mountains on the south, and
eventually from the Coast Ranges on the west. Alluvial
fan deposition along the southeast margin of the basin
(the Kern River Formation) began about 8 Ma, although
there was no comparable event at that time in the
northeast. The coarse elastic sediments of the Kern River
may, then, be evidence that the accelerating late Ceno-
zoic uplift of the Sierra Nevada began earlier at the south
end of the range. In the northeastern part of the basin,
some ﬁne arkosic alluvium of late Pliocene age from the
Sierra Nevada (Laguna Formation) is Virtually unweath-
ered and resembles modern rock ﬂour produced by
glaciers eroding granitic rocks (Marchand, 1977). The
presence of this rock-ﬂour—like alluvium suggests a late
Pliocene onset of alpine glaciation in the Sierra Nevada.

The ﬁnal regression, which began in the latest Mio-
cene, was greatly accelerated through the Pliocene as
progradation of coarse clastic sediments continued from
all sides of the basin, and it culminated with the ﬁnal
retreat of the sea by about the end of the Pliocene. The
stratigraphic sequence in the center of the southern San
Joaquin basin records a gradual shallowing from shallow-
marine shelf (Etchegoin Formation) through restrictive
marine to brackish facies (San Joaquin Formation) and
ﬁnally to freshwater ﬂuvial and lacustrine facies (Tulare
Formation) in the late Pliocene to middle Pleistocene.
This shallowing took place even as the basin continued to
subside, rapidly in the western part of the Maricopa-
Tejon subbasin. The San Joaquin and Tulare are conform-
able in the center of the basin and both interﬁnger
eastward with Kern River alluvial fan deposits, but an
unconformity is present at the base of the Tulare along
the west and south margins of the basin. The San
J oaquin-Tulare conformable contact at Elk Hills, and
presumably elsewhere at the south end of the basin, is
dated at 2.5 to 3.0 Ma (C.A. Repenning, written com—
mun., 1980), which is somewhat older than the equivalent
boundary farther north.

A general decline in sea level after a highstand at about
5 Ma (Haq and others, 1987) may have contributed to the
regression during the Pliocene, but it is more likely that

 

the principal cause was tectonism. Increasing tectonic
activity and uplift to the east, west, and south, especially
following the change in relative plate motion and the
consequent increase in compression normal to the San
Andreas fault at about 5 Ma, resulted in greatly increased
sedimentation that outpaced basin subsidence. Contrib-
uting factors in the transition to a nonmarine basin in the
late Pliocene were the progressive closing off of the
basin’s western outlet by continued northwestward mi-
gration of the Salinia terrane, and folding and uplift in the
Temblor and Diablo Ranges. The unconformity at the
base of the Tulare Formation is due to continued defor-
mation of the western and southern basin margins in
response to compression across the San Andreas fault. In
the San Emigdio Mountains, major north—directed
thrusting on the Pleito fault system produced the coarse
alluvial sediments of the Tulare Formation (Davis, 1983,
1986). The rapid subsidence of the western part of the
Maricopa-Tejon subbasin during the Pliocene and early
Pleistocene, which was concurrent with the shallowing
trend, was probably due to tectonic loading by thrust
plates at the south margin of the basin.

UPPER PLEISTOCENE AND HOLOCENE DEPOSITS

Late Quaternary sedimentation in the San Joaquin
basin consisted of episodic deposition of alluvial sedi-
ments at the valley margins (Marchand and Allwardt,
1981), which grade basinward into a more continuous
section containing a series of lacustrine deposits (Croft,
1972; Marchand, 1977). By about the middle of the
Pleistocene, the San Joaquin basin drainage outlet was
closed or nearly closed, and the impounded drainage
created a. large lake, evidenced by a widespread lacus-
trine clay—the Corcoran Clay Member of the Tulare
(Frink and Kues, 1954) and Turlock Lake (Marchand and
Allwardt, 1981) Formations (ﬁg. 3). Disappearance of the
Corcoran lake was approximately coincident with, and
was probably caused by, the establishment of the present
Central Valley drainage outlet through the Carquinez
Strait and San Francisco Bay at about 0.6 Ma (Sarna-
Wojcicki and others, 1985). The sedimentary record of
the latest Quaternary is similar to that of the middle
Pleistocene, but it contains a succession of smaller pluvial
lakes.

Tectonism has played an important role in the Quater-
nary history of the San Joaquin basin. Closing of the
valley’s drainage outlet and continued uplift of the
surrounding ranges that supply sediment to the alluvial
basin are the consequences of tectonism, but the Quater—
nary sedimentary record reﬂects climatic controls more
than tectonic. Cycles of alluviation, soil formation, and
channel incision in the Quaternary deposits of the north-
eastern San Joaquin basin can be correlated with climatic

24

ﬂuctuations and the resultant glacial stages in the Sierra
Nevada (Bateman and Wahrhaftig, 1966; Marchand,
1977). A similar cyclical pattern is apparent in the alluvial
fan deposits on the west side of the basin that were
derived from the unglaciated Diablo Range (Lettis, 1982,
1985). This pattern suggests that the inferred climatic
control of sedimentation is more complex than a simple
correlation of alluviation event With glacial outwash
event; it involves, as well, the effect of climate on rates
of weathering and on changing vegetation patterns, and
how these factors, in turn, inﬂuence sediment supply
(Marchand, 1977; Lettis, 1982, 1985). Creation or peri-
odic enlargement of pluvial lakes in the center of the
basin can also be correlated with the cyclic alluvial
deposits and with Sierran glaciations (Atwater and
others, 1986). The appearance and disappearance of these
lakes has been dependent on the balance between the
inﬂow volume on one hand, and basin subsidence and the
growth of alluvial fan dams on the other.

100

San Joaquin River

 

, \ . l. .
100— l f , ::~","o::"' _
We -
/\ I , > \ . ‘ _, _ - Lacustnneclay
_ ”,3 \_\—_,:,_ \\-,.'.o ._..D.
\x‘L"\,’," IsL" X’ /\,°'°.° D ,, °.'.°l
\ , \_: \ , \ « / “(,4 'o- o‘ o 0 ° _F|TIT1T Contact—Hachures denote
a 8 _ \e . . . ..
’ - I ~/- a 0 buried 50115 or oxxdized

 

 

   
   
 

THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

BASIN EVOLUTION

The Great Valley of California, which is today an
alluvial plain surrounded by mountains, was, at the
beginning of the Cenozoic, occupied by a marine shelf and
slope system that was part of an extensive forearc basin
at the west edge of North America (Dickinson and Seely,
1979; Ingersoll, 1979). The forearc basin, which origi-
nated in the Mesozoic in a convergent-margin setting,
evolved during the Cenozoic into the hybrid intermon-
tane basin in a transform-margin setting that exists
today. Each of the ﬁve structural regions of the basin,
which show differences in style of deformation, have
somewhat different tectonic histories as a result of the
changes in stress regimes through the Cenozoic. This
evolution of Cenozoic stress, discussed separately for
each region of the basin in the section on “Geology,” is
integrated into the generalized summary diagrams of
ﬁgure 4. Changes in regional stress are closely related in

 

 

' f ~ ' EXPLANATION

o ,

‘ . 1' ,' °~ .° . . i
- . ‘ ." Sierra Nevada detritus

Coast Range detritus

 

horizons

W Cut and ﬁll—Diagrammatic

— _q/ Approximate location of

30 KILOMETERS
| /_ _ facies change

 

 

FIGURE 3. —Diagrammatic cross section of the northern San Joaquin Valley showing stratigraphic relations of Quaternary alluvial deposits in
the valley subsurface. Modiﬁed from Lettis (1982).

BASIN EVOLUTION 25

space and time to the northwestward migration of the
Mendocino triple junction and the consequent evolution
of the San Andreas transform. The changing stress
regimes and the resulting tectonic changes are reﬂected
in the varying patterns of sedimentation and changing
paleogeography during the Cenozoic evolution of the
basin.

The discussion that follows, which also serves as a brief
summary of some of the main points from foregoing
sections of this report, will make reference to a series of
nine paleogeographic maps of the San Joaquin basin and
surrounding regions. These maps were constructed for
fairly narrow time slices of the Tertiary, shown on plate
2 as stippled bands, rather than for broader intervals
such as stages (Zemorrian, Saucesian, and so forth) or
even epochs, as has commonly been done in the past. This
procedure greatly reduces the problem of trying to show
too many, often conﬂicting events on one map, but it has
the disadvantage of implying greater precision in corre-
lation, particularly in the earlier Tertiary, than it is
presently possible to achieve. The maps were con-
structed on palinspastic bases that assume 305 km of
Neogene right-lateral slip on the San Andreas fault
(Clarke and Nilsen, 1973; Graham, 1978) and an unspec—
iﬁed amount of early Paleogene right slip on the proto-
San Andreas fault, together with 150 km of Neogene
right slip on the San Gregorio fault (Clark and others,
1984). It is also assumed that the original trace of the San
Andreas fault was more or less straight and that the
present big bend near the south end of the San Joaquin
Valley was acquired in the late Neogene and Quaternary.
The paleogeography was compiled from published maps
and modiﬁed to accord with more recent stratigraphic
and sedimentologic data. In many cases where data are
sparse or even nonexistent, the maps represent the
interpretations and biases of the author.

PALEOCENE

The forearc basin that existed through the late Meso-
zoic and into the earliest Tertiary began to change in the
Paleocene, although the basin geometry of the old
arc-trench system persisted. The principal factors inﬂu-
encing Paleocene paleogeography were right slip on the
proto—San Andreas fault and uplift of the Stockton arch,
which together reﬂect an overall north-south compres-
sive stress (ﬁg. 4A). This stress was apparently a
consequence of oblique subduction with a northerly to
northeasterly convergence direction (Engebretson and
others, 1985). The proto-San Andreas fault, which orig-
inated in the Late Cretaceous and had, by Paleocene
time, produced a continental borderland of small wrench-
fault basins in the Salinia terrane (Nilsen and Clarke,
1975), was active through the Paleocene. Shear stress

 

associated with right-lateral faulting may have been
responsible for en echelon folding in the northern part of
the west-side fold belt (Harding, 1976) and for oroclinal
bending or rotation of the southern Sierra Nevada
(Kanter and McWilliams, 1982; McWilliams and Li, 1985).
At least part of the folding and rotation may have taken
place during the Paleocene. Uplift of the Stockton arch
and concurrent up-to-the-south movement on the Stock-
ton fault began in the Paleocene and strongly inﬂuenced
sedimentation patterns in the northern part of the basin.

Figure 5 shows the paleogeography of the San Joaquin
basin in the late Paleocene during the transgressive
phase of the upper Paleocene and lower Eocene deposi-
tional sequence. The northern and eastern parts of the
basin were occupied by a marine shelf; deeper marine
slope and basinal facies were restricted to the southwest.
An upland, probably of low relief, lay to the northeast,
and the shelf and slope were largely open to the ocean on
the west. Uplift of the Stockton arch produced a broad
west-trending peninsula between the San Joaquin and
Sacramento basins. It is not known how much of the arch
was exposed at this time because erosion has removed
Paleocene strata, but it is assumed that some shallow or
nearshore marine deposition took place over the west end
of the arch. The extent of the Diablo uplift is also not
known, but it was probably not large and represented
only a portion of the structural high in the subduction
complex at the southwest side of the forearc basin that
was uplifted with the Stockton arch. At the south end of
the basin, the ﬁrst stages of the oroclinal bending of the
southern Sierra Nevada had produced a westward devi-
ation of the southeast-trending shoreline and left the
future Bakersﬁeld arch and the Maricopa—Tejon subbasin
area emergent.

EOCENE

The Eocene, as a result of recent revisions in Cenozoic
geochronology (Berggren and others, 1985), is the long-
est of the Cenozoic epochs. Although the Paleogene was
generally “quieter” than the Neogene, a number of
events, including a major regression separating two
basinwide depositional sequences, affected the basin
during the Eocene. The combined effects of tectonism
and eustatic sea-level change resulted in broad ﬂuctua-
tions in the shoreline and produced major changes in
paleogeography. Tectonic activity included uplifts in the
Diablo Range area in late early and late Eocene, and
possibly in the San Emigdio Range area near the end of
the Eocene. In the north, uplift was probably in response
to northeast-southwest subduction-related compression
that might be considered evidence of Paleogene emplace—
ment of a deep-seated Franciscan wedge, whereas in the
south, it was probably in response to regional north-
south compression (ﬁg. 4A).

26 THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

 

 

Stockton

_...o
a'
"a;
Z
‘7
z
3’:
33
(D
“a;
C
‘2‘

Stockton
. . D. .

        
 

 

 

 

 

 

 

 

 

 

Stockton Stockton
. . . D. . . . .0. .

  

0

Fresno

 

 

 

 

 

 

 

FIGURE 4. —Evolution of regional stress patterns in the San Joaquin Valley area. A, Early Paleogene. B, Late Oligocene-early Miocene. C , Late
Miocene-early Pliocene. D, Late Pliocene-Quaternary. Edge of present valley alluvium and present major faults generalized from ﬁgure 1 for
location. Small arrows indicate inferred local stress, large arrows indicate inferred overall compressive stress, paired half-arrows indicate
shear couple; queried where uncertain.

BASIN EVOLUTION 27

A major regression occurred at the end of the upper
Paleocene-lower Eocene depositional sequence. This re-
gression was largely a result of eustatic lowering of sea
level that left the marine basin greatly restricted (ﬁg. 6).
Although their presence somewhat speculative, large
deltas are inferred to have prograded westward across
the basin from the central Sierra Nevada, which was the
principal source of Eocene sediment. Parts of the sub—
duction complex at the west side of the basin were
emergent and contributed some sediment, and Salinia
terrane highs were emergent to the southwest (Nilsen
and Clarke, 1975). At the beginning of the ensuing
transgression, basal sands containing Franciscan detri-
tus onlapped the formerly emergent areas on the west
side of the basin. Much of the formerly emergent area at
the south end of the basin, including the Bakersﬁeld arch,
was also inundated. A condensed section in the lower part
of the Kreyenhagen Shale (Milam, 1985) immediately
above the basal sand of the Eocene depositional sequence
is an indication of a rapid transgression.

The point of maximum transgression for the entire
Tertiary was reached at middle Eocene time (ﬁg. 7).

EXPLANATION ‘

Marine deposits—Solid line indicates inferred
shoreline; hachures indicate inferred shelf edge;
queried where uncertain

 

Nonmarine deposition—Dotted line indicates in-
ferred extent; queried where uncertain

 

 

 

Lacustrine deposition

 

 

Emergent area—Queried where uncertain

 

 

AAAAAA
AAAAAAA
AAAAAA

 

 

Volcanic center

 

Faults—Queried where uncertain. Arrows indicate
direction of relative movement

-—:—— Probably active

Possibly active

Future trace of Neogene San Andreas fault
Thrust—Sawteeth on upper plate

Explanation for ﬁgures 5 through 13.

 

Widespread pelagic sediments indicate that most of the
present-day San Joaquin Valley was covered by deep-
marine waters and the basin was largely open to the
west, as it had been in the Paleocene. A large deep-sea
fan was constructed in the southwestern part of the basin
that had its source and proximal part on the Salinia
terrane west of the present San Andreas fault (Clarke,
1973; Clarke and Nilsen, 1973). The east side of the basin
was fringed by a belt of ﬂuvial and deltaic deposits.

The regressive phase of the Eocene depositional se—
quence near the end of the Eocene is recorded in the
northern part of the basin and at the south end, whereas
the record indicates continued deep-water deposition in
the central part. Uplifts at this time represent the
beginning of tectonic activity that continued into the
Oligocene, and probably reﬂect the approach of the
Paciﬁc-Farallon spreading ridge and the transition from
oblique to normal subduction (Engebretson and others,
1985).

 

Mount 3 “393503 {-

Diamon’Canyon r.
4’" H Isn‘t},

’Imm’lm

 

4/

/

0 10 20 30 40 50 KILOMETERS
LLl—I_I_l

o
Bakersﬁeld

 

 

“In”
min“ ”m
\\I\“‘"” ’IIII/
I,"
1,,"
Hm”
‘

 

FIGURE 5. —Late Paleocene (about 59 Ma) paleogeography of the San
Joaquin basin area. Based on data from Repenning (1960), Clarke and
others (1975), Dickinson and others (1979), Nilsen and McKee (1979),
Clark and others (1984), Fischer (1984), and Stanley (1985).

28

OLIGOCENE

The Oligocene marks the onset of change throughout
the western United States. Following the change in
convergence direction near the end of the Eocene, the
angle of subduction steepened during the Oligocene
(Page and Engebretson, 1984), leading to changes in the
patterns of volcanism (Snyder and others, 1976) and to
the initiation of extension in the Basin and Range
province (Zoback and others, 1981). The most important
event affecting California during the Oligocene was the
ridge-trench encounter off southern California that initi-
ated the San Andreas transform system (Atwater, 1970;
Atwater and Molnar, 1973). All these events had an
effect, whether directly or indirectly, on the San Joaquin
basin during the Oligocene and, augmented by eustatic
sea-level change, produced major changes in the central
California paleogeography.
Continued uplift of the Stockton arch in the northern
part of the basin, concurrent with movement on the

  
 
 

  
 

4 ;~\ -
”4"... "S"? - .
"Jim"

: ﬁdiim} 7,,
, § ' Santa

Km i\\"<“"“um¥’
\
“II V

0 lo 20 30 4O 50 KILOMETERS
..__.._;_._._a

 

FIGURE 6. —-Early to middle Eocene (about 52 Ma) paleogeography of
the San Joaquin basin area. See ﬁgure 5 for explanation. Based on
data from Repenning (1960), Clarke and others (1975), Graham
(1978), Graham and Berry (1979), Nilsen (1979), Nilsen and McKee

 

THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

Stockton reverse fault, and of the San Emigdio area in
the south are both indications of a general north-south
compression. In the north, the tectonism was presumably
subduction related. At the south end of the basin, at least
during the later Oligocene, the tectonism was part of a
major uplift throughout southern California in response
to the ridge-trench encounter (Nilsen, 1984).

The entire northern part of the basin was emergent
through the Oligocene, while deep-water sedimentation
continued from the Eocene into the Oligocene in the
southwestern part. At the time of maximum regression
in the middle Oligocene (ﬁg. 8), the Stockton fault was
active, the Stockton arch was being eroded, and alluvium
was deposited to the north and south of the arch. The
earliest rhyolitic tephra were also deposited in the
northern Sierra Nevada at about this time. The Diablo
Range and probably the northern Temblor Range areas
were emergent, so deep-marine deposition was restricted
to the southwestern part of the formerly extensive basin.
In addition to the nonmarine deposits south of the

4/

/

0 10 20 30 40 50 KILOMETERS

 

 

FIGURE 7.—Middle Eocene (about 44—45 Ma) paleogeography of the

San Joaquin basin area. See ﬁgure 5 for explanation. Based on data

from Repenning (1960), Clarke (1973), Clarke and others (1975),

 

(1979), Slagle (1979), and Stanley (1985).

Nilsen (1979), Nilsen and McKee (1979), and Palmer and Merrill
(1982).

BASIN EVOLUTION 29

Stockton arch, a narrow fringe accumulated along the
east side of the marine embayment, while coarser alluvial
fan deposits derived from uplifts to the south accumu-
lated along the south and southeast margins of the basin.
A north-south axial proﬁle of the basin ﬂoor at this time
would start above sea level in the north (probably not
more than a few tens of meters) and reach depths of over
1,800 m in the south, illustrating the marked southward
tilt of the basin in the Oligocene.

MIOCENE

The evolution of the San Joaquin basin was accelerated
by tectonic events of the Miocene, most of which were a
result of the northwestward migration of the Mendocino
triple junction. The interval during which the triple
junction was located opposite the basin nearly coincides
with the Miocene epoch (Snyder and others, 1976;
Johnson and O’Neil, 1984) (pl. 2). Paleogeographic

 

      
  

we on.

STOCKTON ‘40,}

v

 
 

-" 7
' \EDISON FAULT

0 IO 20 30 40 60 KILOMETERS
_._._._._.

 

 

FIGURE 8. —Oligocene (about 30 Ma) paleogeography of the San
Joaquin basin area. See ﬁgure 5 for explanation. Based on data from
Repenning (1960), Addicott (1968), Bandy and Arnal (1969), Greene
and Clark (1979), Nilsen and McKee (1979), Nilsen (1984), and Pence
(1985).

 

changes took place at a faster pace, particularly adjacent
to the developing San Andreas fault system. Marine
deposition was restricted to the southern part of the
basin, but extensive nonmarine deposition began in the
north. There was also a major change in regional volcan-
ism, as the rhyolitic pyroclastic deposits of the late
Oligocene and earliest Miocene were replaced near the
end of the early Miocene by reestablished andesitic arc
volcanism in the northern Sierra Nevada (Slemmons,
1966; Stewart and Carlson, 1976). The are volcanism was
then progressively shut off from south to north as the
Mendocino triple junction migrated northward and cut off
subduction (Snyder and others, 1976).

The patterns of regional stress, which had begun to
change in the Oligocene, continued to change through the
Miocene. This was ﬁrst apparent at the south end of the
basin where east-west-oriented normal faulting and sub-
sidence, probably beginning in the latest Oligocene, and
volcanism in the early Miocene indicate north—south
extension (Davis, 1986; Hirst, 1986) (ﬁg. 43). En echelon
folding in the west-side fold belt, beginning near the end
of the early Miocene and continuing through the Miocene,
is a manifestation of a newly established northwest-
southeast-oriented shear couple centered on the San
Andreas fault system (Harding, 1976). The accelerating
uplift of the Sierra Nevada during the late Miocene
(Huber, 1981) was accompanied by north-south-oriented
normal faulting that indicates minor east-west extension
in the southern Sierran block. The Diablo uplift and the
eruption of the Quien Sabe Volcanics in the late Miocene
are closely associated with the passage of the Mendocino
triple junction (Johnson and O’Neil, 1984). The uplift
seems to indicate compression, presumably oriented
northeast-southwest, and was followed immediately by
minor local extension in a developing northwest-south-
east shear regime.

The early Miocene marine embayment (ﬁg. 9) was not
very different from the Oligocene embayment. The
northern Temblor Range area that was brieﬂy exposed at
mid-Oligocene time was inundated in the early Miocene,
as it had been in the late Oligocene, and the early Miocene
strandline advanced even farther eastward onto the
southern Sierran block. Nonmarine deposition expanded
northwestward as tuffaceous alluvial plain deposits cov-
ered the northern basin and Stockton arch areas. The last
stages of the previously extensive coarse alluvial fan
deposition took place at the south end of the basin.

The paleogeography changed considerably near the
early Miocene-middle Miocene boundary (ﬁg. 10). Uplift
of the southern part of the basin produced a regression
there as alluvial fan and fan-delta deposits prograded
basinward. Farther north, however, there was a trans-
gression, as the strandline advanced northwestward onto
the Diablo uplift in the Coalinga area and in the Vallecitos

3O

syncline. There was probably a shallow seaway trending
northwest through the Vallecitos syncline. The initiation
of wrench tectonism on the southwest side of the basin
resulted in uplifts in the adjacent Salinia terrane and
nonmarine deposition in the southern Diablo Range area.
The beginning of andesitic volcanism in the northern
Sierra Nevada is reﬂected in the nonmarine volcaniclastic
deposits of the northern San Joaquin and Sacramento

THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA.

fan deposits were derived from the east, the south, and
the southwest.

By late Miocene time (ﬁg. 11) the northward move-
ment of the Salinia terrane, composed of isolated highs
surrounded by shallow seas (Graham, 1978), west of the
San Andreas fault was beginning to close off the San

basins.

After the early to middle Miocene regression,

marine embayment expanded to its Neogene maximum

Joaquin basin on the west. A new seaway had opened
through the Priest Valley area west and northwest of
Coalinga, but there was no longer a deep-marine outlet to
the the Paciﬁc Ocean. The deep-marine embayment was

becoming more restricted as shallow-marine shelf depos-
extent, approximately coincident with a middle Miocene its and nonmarine deposits prograded basinward along
highstand of sea level (pl. 2). Marine deposits of late the east side of the basin. Sedimentation in the northern
middle Miocene age reach as far north as Chowchilla (70 part of the basin was still dominated by volcaniclastic
km northwest of Fresno). The basin axis at that time sediments derived from the extensive andesitic volcan-
seems to have been considerably farther east than the ism in the Sierra Nevada, but the inception of coarse
present axis, probably because uplift of the southern alluvial fan deposition of sediments derived from the
Diablo Range and consequent sediment inﬂux from the

west forced the northern basin axis to the east. In the
deep southernmost part of the basin, extensive deep—sea

   

Mount
Diablo

 

- O} Fresno

 

hi11thimxmyhhhh”mummm
n ,

 

,:
*2
'5;
c.
,
‘3
F;
"a

\.

“a“
1
’Ihmm : m mamas“

.
/«
’2'

’9’."

55,6. 1 "”411,
. 4."

 

.1”

(a
. .

,mk -

Santa Cruz I

Q
4L
*4,

a e
. a \
‘. 1”"! ma mh‘w‘

1m ,1. mamhmmmlﬂm

4 0 IO 20 30 IO 50 KlLOMETERS
14, __._.__

 

nnnh

   
 
      

Diablo Range on the northwest indicates uplift of that
range. Volcanic centers were active in the Sierra Nevada
(Moore and Dodge, 1980) and the central Diablo Range.

 

      
  
  
  

3
s ' ' Relief Peak
5: volcanic

  

'A. Mqhnl biablo .

 

"gFresno

f/
O S

 

01
’ Hum,,,,m\ o 10 20 so 40 so KILOMETERS
._.__._‘_._.

 

-€‘%‘i-. -.; . I
\7
DOV/5

“unkind”Mimi/m,

 

FIGURE 9. — Early Miocene (about 20—21 Ma) paleogeography of the San
Joaquin basin area. See ﬁgure 5 for explanation. Based on data from
Gale and others (1939), Repenning (1960), Addicott (1968), Bandy and

Arnal (1969), Graham (1978), Kuespert (1983), Pence (1985), and
Stanley (1985).

 

FIGURE 10. —Middle Miocene (about 16 Ma) paleogeography of the San
Joaquin basin area. See ﬁgure 5 for explanation. Based on data from
Repenning (1960), Addicott (1968), Bandy and Arnal (1969), Fritsche

 

(1977), Graham (1978), Cooley (1982), Kuespert (1983), Bate (1984),
Bent (1985), Pence (1985), and Stanley (1985).

BASIN EVOLUTION 31

Flows from the Sierra Nevada centers reached at least to
the basin margin along the ancestral Stanislaus and San
Joaquin Rivers.

A Widespread unconformity in the southern part of the
basin marks the culmination of the late Miocene regres-
sion (pl. 2). Coarse alluvial fan sedimentation along the
southeast margin of the basin in the latest Miocene marks
the beginning of the accelerated late Neogene uplift of
the Sierra Nevada.

PLIOCENE

Neogene tectonic activity around the San Joaquin
basin increased in intensity during the Pliocene, leading
to the elimination of the marine embayment by the close
of the Pliocene. The San Andreas fault had become the
principal element of the transform system by the begin-
ning of the Pliocene (Graham, 1978), and the consequent
increase in slip rates caused the Salinia terrane to move

 

 

Slanlslaus
volcanlc
center

   
   
   

San Joaquin
volcanic

0‘. Fresno
' ’ center

 

4/

/

0 ‘0 20 30 to SD KILOMETERS

V i EDISON FAULT

WHITE WOLF FAULT

 

GARLOCK _' FAULT

E

 

FIGURE 11. —Late Miocene (about 9—10 Ma) paleogeography of the San
Joaquin basin area. See ﬁgure 5 for explanation. Based on data from
Repenning (1960), Addicott (1968), Bandy and Arnal (1969), Fritsche
(1977), MacPherson (1978), Phillips (1981), and Graham and others
(1982).

 

rapidly northward, cutting off the southwestern marine
connection with the Paciﬁc Ocean.

The regional stress pattern that originated in the
Miocene remained in effect into the Pliocene, with only
moderate change (ﬁg. 4C). The northwest-southeast
shear couple associated with the San Andreas fault
extended farther northwest, and there was an increase in
compressive stress normal to the San Andreas as a result
of changes in plate convergence direction near the
Miocene-Pliocene boundary (Page and Engebretson,
1984; Engebretson and others, 1985). This change in
plate motions, together with the westward movement of
the Sierran block as a result of extension in the Basin and
Range province (Eaton, 1979), caused northeast-south-
west compressive stress along the west side of the
Sierran block and was largely responsible for increased
late Neogene deformation in the fold belt, including
deep-seated thrust faults (Wentworth and Zoback, 1986;
Zoback and others, 1987). Strong north-south compres-
sion at the south end of the basin, probably due in part to
the developing bend in the San Andreas fault, was
responsible for the onset of northward-directed thrusting
in the late Pliocene. The increased loading of the south
end of the basin by thrust plates was, in turn, probably
responsible for the accelerated subsidence of the western
part of the Maricopa—Tejon subbasin in the latest Plio-
cene.

The paleogeography of the Pliocene (ﬁg. 12) differed
signiﬁcantly from that of the late Miocene. The embay-
ment was much smaller and the basin, mostly brackish by
this time, was enclosed on the south and southwest.
Nonmarine deposits prograded into the shallowing em-
bayment from all sides and the emergent Salinia terrane
was transported northwestward to completely close the
marine outlet by about the end of the Pliocene. A
developing uplift lay south of the basin, While the'western
part of the Maricopa-Tejon subbasin was subsiding rap-
idly. A shallow seaway west of Coalinga connected the
rapidly shallowing embayment with the Paciﬁc Ocean. In
the northern part of the basin, there was an increasing
sediment supply to the west-side alluvial fans from the
rising Diablo Range. On the northeast side of the basin,
a change from volcaniclastic to arkosic alluvium indicates
that major Sierran rivers had cut down through the
blanket of volcanic rocks. Although alpine glaciers in the
Sierra Nevada probably appeared before the end of the
Pliocene, there is no direct evidence of glaciation prior to
about 2.5 Ma.

PLEISTOCENE

By the beginning of the Pleistocene, the San Joaquin
basin Was entirely emergent and was largely enclosed by
the ice-capped Sierra Nevada on the east and by low hills

32 THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

of the Coast Ranges on the west. The valley itself
differed from the present valley principally in having its
outlet somewhere in the southwest rather than through
the Carquinez Straits to San Francisco Bay, as at
present.

Uplift and westward tilting of the Sierra Nevada
continued through the Pleistocene, while major deforma-
tion and uplift of the Coast Ranges, begun during the late
Pliocene, also continued. Regional patterns of stress
changed little during the Pleistocene (ﬁg. 4D). North-
east-southwest compression normal to the San Andreas
fault resulted in the southern part of the west-side fold
belt being uplifted as the Temblor Range. The Diablo
Range was also uplifted, largely as a whole, with little
internal differential vertical movement (Page, 1981).
Right slip on the Ortigalita fault is evidence that the
northwest—southeast shear was also imposed on the
Diablo Range (Lettis, 1982, 1985). A very strong north-
south compressive stress at the south end of the valley
resulted in overthrusting on the faults of the Pleito

 

 
   

3
o
E
g'
gi
an
E
a

  
 

    

VNlAVH '

‘ '_ oS‘tockt‘o'h:

nnva swimwear;

onus/um FAULT

 
  

 

  
  

'- . 0' .1
’.Fresno'-._

WHITE OLF FAULT 0 W 10 30 w 50 KILOMETERS

FAULT

 

GARLOCK _'

 

FIGURE 12.—Pliocene (about 3—4 Ma) paleogeography of the San
Joaquin basin area. See ﬁgure 5 for explanation. Based on data from
Repenning (1960), Galehouse (1967), Foss (1972), Cole and Armen-
trout (1979), and Greene and Clarke (1979).

 

system and reverse movement on the White Wolf fault
(Davis, 1986).

A major feature of the Pleistocene paleogeography was
the large lake, the Corcoran lake, that occupied nearly
the whole valley for a brief interval near the middle of the
Pleistocene (ﬁg. 13). This lake was the largest and
perhaps the earliest of a succession of lakes that occupied
the valley during the Quaternary. Alpine glaciers in the
Sierra Nevada fed rivers as far south as the Kern River,
which deposited an apron of outwash along the east side
of the valley and built deltas into the lake. After the
withdrawal of the sea from the marine embayment at the
end of the Pliocene, the drainage outlet of the valley
probably remained along the old Priest Valley seaway for
a short time. It seems likely, though, that continued
uplift and deformation along the San Andreas fault zone
between the the Diablo Range and Gabilan Range to the
west would have soon closed this outlet. A possible
alternative outlet lay farther south at Bitterwater Val-
ley, where the valley drainage could have crossed the
northern Temblor Range (pl. 2) to ﬂow down the now

 

 
   

  
 

CALAvsnAs FAULT

ORTIGALITA FAULY

  
    

"i

‘ '3 : WHITE WOLF FAULT 0 1D 10 30 w 50 KILOMETERS
"e ‘—‘——‘—‘—'
1_;..

PLEIYO \FAULT
GARLOCK -' FAULT

 

 

FIGURE 13. —Pleistocene (about 0.6—0.7 Ma) paleogeography of the San
Joaquin Valley area. See ﬁgure 5 for explanation. Based on data from
Wahrhaftig and Birman (1965), Croft (1972), and Page (1986).

CONCLUSIONS 33

underﬁt Cholame and Estrella Creeks to join the Salinas
River north of Paso Robles. This alternative was chosen
for the mid—Pleistocene paleogeography of ﬁgure 13. In
either case, the drainage was at least partly impounded
at that time to form the Corcoran lake. The disappear-
ance of the lake (at about 0.6 Ma) is probably a result of
the opening of the present Central Valley drainage outlet
through the Carquinez Straits (Sarna-Wojcicki and oth-
ers, 1985).

HOLOCENE

It is commonly assumed that Quaternary tectonism
continues unabated to the present day, although the
speciﬁc evidence is limited. The Corcoran Clay Member
of the Tulare Formation today lies at depths of 60 m to

. more than 200 m below sea level along the synclinal axis
of the San Joaquin basin (Miller and others, 1971; Croft,
1972), indicating the magnitude of subsidence in the
valley in the past 600,000 years. The sharp upbending of
the Corcoran along the west side of the syncline, further-
more, is evidence of Coast Ranges uplift during the same
period. Historical subsidence due to ﬂuid withdrawal
(Poland and Evenson, 1966), however, tends to mask
recent tectonic subsidence. Geodetic measurements also
seem to indicate continuing deformation (Burford, 1965;
Stein and Thatcher, 1981), although the probable errors
inherent in surveying often approach in magnitude the
deformations being measured.

The most unambiguous evidence of continuing tecton-
ism is seismicity. Although historical seismicity in Cali-
fornia has been dominated by the San Andreas fault
system, there have been a few moderate to large
earthquakes within the San Joaquin Valley, most notably
the 1952 magnitude 7.2 Arvin-Tehachapi and the 1983
magnitude 6.5 Coalinga earthquakes. The 1952 Arvin-
Tehachapi earthquake was centered on the White Wolf
fault (Oakeshott, 1955; Stein and Thatcher, 1981), and
the oblique slip during that event, reverse slip plus a
left-lateral component, is evidence of an existing north-
south to northeast-southwest compressive stress at the
south end of the valley. The 1983 Coalinga earthquake
occurred on a northeast-verging thrust fault under the
Coalinga anticline (Eaton, 1985b) and may be related to
folding and thrusting along the entire west margin of the
Central Valley (Wentworth and Zoback, 1986). Lower
level seismicity has been recorded from several areas in
the San Joaquin Valley (La Forge and Lee, 1982; Eaton,
1985a; Wong and Ely, 1983; Wong and Savage, 1983) and
collectively indicates north-south to northeast-southwest
compression. Coseismic uplift of as much as 45 cm
associated with the Coalinga earthquake (Stein, 1985)
demonstrates the continuing growth of young anticlines
at the west side of the valley.

 

CONCLUSIONS

The Paleogene history of the San Joaquin basin was
dominated by a tectonic regime resulting from the
presence of a subduction zone lying along the continental
margin to the west. Oblique convergence in the early
Paleogene (Page and Engebretson, 1984) produced a
north-south compressive stress and a right-lateral shear
couple in the western part of the continent. Right-lateral
slip on the proto-San Andreas fault and the northwest—
ward movement of the Salinia terrane into position
opposite the south end of the basin (Nilsen and Clarke,
1975; Graham, 1978; Dickinson and others, 1979), large
en echelon folds in the southern Diablo Range (Harding,
1976), and clockwise rotation of the southernmost part of
the Sierra Nevada (Kanter and McWilliams, 1982;
McWilliams and Li, 1985) are all consequences of the
early Paleogene stress regime. Although this tectonism
shaped the underlying structural framework and
strongly inﬂuenced Paleogene geography, eustatic sea-
level change also had a major inﬂuence on the Paleogene
sedimentary record and geography. A eustatic fall in sea
level was probably the principal cause for the regression
at the end of the upper Paleocene and lower Eocene
depositional sequence; it was also a contributing factor
for each of the other Paleogene regressions, except the
ﬁnal one at the end of the upper Oligocene sequence.

Neogene tectonism and basin evolution were controlled
at ﬁrst by the tectonic effects of the northwestward
migration of the Mendocino triple junction along the
California continental margin, and they were later con-
trolled by wrench tectonism associated with the San
Andreas fault system (Dickinson and Snyder, 1979; Page
and Engebretson, 1984). The ﬁrst effects of Mendocino
triple-junction passage, felt at the south end of the basin
beginning at about 23—24 Ma, were extension-induced
subsidence and volcanism. These events were followed at
16—17 Ma by regional uplift in the southern part of the
basin; the uplift seems to have been associated with
passage of the triple junction and may have been related,
in some way, to the presence of the subducted fracture
zone under the basin (Loomis and Glazner, 1986). Con-
tinued subsidence after the uplift, accompanied by
wrench tectonism in the fold belt, may have been
augmented by thermal decay of the subducted plate. The
transgression resulting from middle Miocene subsidence
was augmented by a eustatic highstand of sea level
(Graham and others, 1982). Two lines of evidence from
the San Joaquin basin, the inception of en echelon folding
near the end of the Saucesian (Harding, 1976) and the
distribution and inferred western provenance of Temblor
Formation detritus (Graham and others, 1986), indicate
that San Andreas fault movement may have begun as
early as late early Miocene time along the Temblor Range

34 THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

segment. Folding continued through the later Cenozoic
and deformation increased in intensity near the San
Andreas fault in the Pliocene and Pleistocene as a result
of increased fault-normal compression. However, ongo-
ing deformation at the east edge of the Coast Ranges, and
by implication, the basinward expansion of the fold belt,
may be considered evidence for deep-seated, eastward-
directed thrusting (Wentworth and Zoback, 1986).

REFERENCES CITED

Addicott, W.O., 1968, Mid-Tertiary zoogeographic and paleogeographic
discontinuities across the San Andreas fault, in Dickinson, W.R.,
and Grantz, Arthur, eds., Proceedings of conference on geologic
problems of the San Andreas fault system: Stanford University
Publications, Geological Sciences, v. 11, p. 144—165.

— 1970, Tertiary paleoclimatic trends in the San Joaquin basin,
California: US Geological Survey Professional Paper 644-D, 19 p.

—- 1973, Oligocene molluscan biostratigraphy and paleontology of
the lower part of the type Temblor Formation, California: US.
Geological Survey Professional Paper 791, 48 p.

Almgren, A.A., 1978, Timing of Tertiary submarine canyons and
marine cycles of deposition in the southern Sacramento Valley,
California, chap. 19 of Stanley, D.J., and Kelling, G., eds.,
Sedimentation in submarine canyons, fans, and trenches: Strouds-
burg, Penn., Dowden, Hutchinson & Ross, p. 276—291.

Almgren, A.A., Filewicz, M.V., and Heitman, H.L., 1988, Lower
Tertiary foraminiferal and calcareous nannofossil zonation of
California: an overview and recommendation, in Filewicz, M.V.,
and Squires, R.L., eds., Paleogene stratigraphy, west coast of
North America: Society of Economic Paleontologists and Miner-
alogists, Paciﬁc Section, Annual meeting, 1988, Santa Barbara,
Calif., p. 83—105.

Anderson, L.W., Anders, M.H., and Ostenaa, D.A., 1982, Late
Quaternary faulting and seismic hazard potential, eastern Diablo
Range, California, in Hart, E.W., Hirschfeld, SE, and Schulz,
S. S., eds., Proceedings of conference on earthquake hazards in the
eastern San Francisco Bay region: California Division of Mines and
Geology Special Publication 62, p. 197—206.

Antonnen, G.J., Danehy, E.A., and Fallin, J .A.T., 1974, The Kings
Canyon lineament: a cross-grain ERTS—l lineament in central
California, in Proceedings of First International Conference on
the new basement tectonics: Utah Geological Association Publica-
tion 5, p. 81—93.

Atwater,/BF” Adam, D.P., Bradbury, J.P., Forester, R.M., Mark,
R.K., Lettis, W.R., Fisher, G.R., Gobalet, K.W., and Robinson,
S.W., 1986, A fan dam for Tulare Lake, California, and implica-
tions for the Wisconsin glacial history of the Sierra Nevada:
Geological Society of America Bulletin, v. 97, p. 97—109.

Atwater, Tanya, 1970, Implications of plate tectonics for the Cenozoic
tectonic evolution of western North America: Geological Society of
America Bulletin, v. 81, p. 3513—3536.

Atwater, Tanya, and Molnar, Peter, 1973, Relative motion of the
Paciﬁc and North American plates deduced from sea-ﬂoor spread-
ing in the Atlantic, Indian, and South Paciﬁc oceans, in Kovach,
R.L., and Nur, Amos, eds., Proceedings of conference on the
tectonic problems of San Andreas fault system: Stanford Univer-
sity Publications, Geological Sciences, v. 13, p. 136—148.

Aubry, Marie-Pierre, 1985, Northwestern European Paleogene mag-
netostratigraphy, biostratigraphy, and paleogeography: calcare—
ous nannofossil evidence: Geology, v. 13, p. 198—202.

Bailey, E.H., Irwin, W.P., and Jones, D.L., 1964, Franciscan and

 

related rocks, and their signiﬁcance in the geology of western
California: California Division of Mines and Geology Bulletin 183,
177 p.

Bandy, O.L., and Arnal, RE, 1969, Middle Tertiary basin develop-
ment, San Joaquin Valley, California: Geological Society of Amer-
ica Bulletin, v. 80, p. 783—819.

Barr, F.T., and Berggren, W.A., 1981, Lower Tertiary biostratigraphy
and tectonics of northeastern Libya, in Salem, M., and Busrewil,
M-T., eds, The geology of Libya, v. 1: London, Academic Press,
p. 163—192.

Barron, J .A., 1986, Paleoceanographic and tectonic controls on depo-
sition of the Monterey Formation and related siliceous rocks in
California: Paleogeography, Paleoclimatogy, Paleoecology, v. 53,
p. 27—45.

Bartow, J.A., 1980, Constraints on the latest movements on the

Melones fault zone, Sierra Nevada foothills, California: US.

Geological Survey Professional Paper 1126—J, 4 p.

1983, Map showing conﬁguration of the basement surface,
northern San Joaquin Valley, California: US. Geological Survey
Miscellaneous Field Studies Map MF—1430, scale 1:250,000.

— 1984, Geologic map and cross sections of the southeastern margin
of the San Joaquin Valley, California: US Geological Survey
Miscellaneous Investigations Series Map I—l496, scale 1:125,000.

— 1985, Map and cross sections showing Tertiary stratigraphy and

structure of the northern San Joaquin Valley, California: US.

Geological Survey Miscellaneous Field Studies Map MF—1761,

scale 1:250,000.

1990, Cenozoic stratigraphy and geologic history of the Coalinga
region, central California, chap. 1 of Rymer, M.J., and Ellsworth,
W.L., eds., The Coalinga, California earthquake of May 2, 1983:
US. Geological Survey Professional Paper 1487, p. 3—13.

Bartow, J .A., Lettis, W.R., Sonneman, H.S., and Switzer, J .R., Jr.,
1985, Geologic map of the east ﬂank of the Diablo Range from
Hospital Creek to Poverty Flat, San Joaquin, Stanislaus, and
Merced Counties, California: US Geological Survey Miscella-
neous Investigations Series Map I—1656, scale 1262,500.

Bartow, J .A., and McDougall, Kristin, 1984, Tertiary stratigraphy of
the southeastern San Joaquin Valley, California: US. Geological
Survey Bulletin 1529—J, 41 p.

Bate, M.A., 1984, Sedimentary facies and depositional environments of
the Big Blue Formation between Anticline Ridge and the Do-
mengine Ranch, in Carlson, Christine, ed., Depositional facies of
sedimentary serpentinite: selected examples from the Coast
Ranges, California: Society of Economic Paleontologists and
Mineralogists, Midyear Meeting, 1984, San Jose, Calif., Field Trip
Guidebook no. 3, p. 81—85. ‘

Bateman, RC, and Wahrhaftig, Clyde, 1966, Geology of the Sierra
Nevada, in Bailey, E.H., ed., Geology of northern California:
California Division of Mines and Geology Bulletin 190, p. 107—172.

Bent, J .V., 1985, Provenance of upper Oligocene—middle Miocene
sandstones of the San Joaquin basin, California, in Graham, S.A.,
ed., Geology of the Temblor Formation, western San Joaquin
basin, California: Society of Economic Paleontologists and Miner-
alogists, Paciﬁc Section, Fall 1985 Field trip, Coalinga, Calif.,
Guidebook, p. 97—120.

Berggren, W.A., and Aubert, Jane, 1983, Paleogene benthic forami-
niferal biostratigraphy and paleobathymetry of the central ranges
of California, in Brabb, E.E., ed., Studies in Tertiary stratigraphy
of the California Coast Ranges: US Geological Survey Profes-
sional Paper 1213, p. 4.21.

Berggren, W.A., Kent, D.V., Flynn, J.J., and Van Couvering, J.A.,
1985, Cenozoic geochronology: Geological Society of America
Bulletin, v. 96, p. 1407—1418.

Blake, M.C., Jr., Campbell, R.H., Dibblee, T.W., Jr., Howell, D.G.,
Nilsen, T.H., Normark, W.R., Vedder, J.C., and Silver, E.A.,

 

 

REFERENCES CITED 35

1978, Neogene basin formation in relation to plate—tectonic evolu-
tion of San Andreas fault system, California: American Associa-
tion of Petroleum Geologists Bulletin, v. 62, p. 344—372.

Bohannon, R.G., and Howell, D.G., 1982, Kinematic evolution of the
junction of the San Andreas, Garlock, and Big Pine faults,
California: Geology, v. 10, p. 358—363.

Briggs, L.I., Jr., 1953, Geology of the Ortigalita Peak quadrangle,
California: California Division of Mines and Geology Bulletin 167,
61 p.

Burford, R.O., 1965, Strain analysis across the San Andreas fault and
Coast Ranges of California, in Proceedings 2nd International
Symposium on Recent Crustal Movements, Aulanko, Finland, p.
100-110.

Cady, J.W., 1975, Magnetic and gravity anomalies in the Great Valley
and western Sierra Nevada metamorphic belt, California: Geolog-
ical Society of America Special Paper 168, 56 p.

Carter, J.B., 1985, Depositional environments of the type Temblor
Formation, Chico Martinez Creek, Kern County, California, in
Graham, S.A., ed., Geology of the Temblor Formation, western
San Joaquin basin, California: Society of Economic Paleontologists
and Mineralogists, Paciﬁc Section, Fall 1985 Field trip, Coalinga,
Calif, Guidebook, p. 5—18.

Castle, R.O., Church, J.P., Yerkes, R.F., and Manning, J.C., 1983,
Historical surface deformation near Oildale, California: US.
Geological Survey Professional Paper 1245, 42 p.

Cherven, V.B., 1983, A delta—slope-submarine fan model for Maes—
trichtian part of Great Valley sequence, Sacramento and San
Joaquin basins, California: American Association of Petroleum
Geologists Bulletin, v. 67, p. 772—816.

Christensen, MN, 1966, Late Cenozoic crustal movements in Sierra
Nevada: Geological Society of America Bulletin, v. 77, p. 163—182.

Church, H.V., Jr., and Krammes, Kenneth, co—chairmen, 1957, Corre-
lation section across central San Joaquin Valley from San Andreas
fault to Sierra Nevada foothills: American Association of Petro-
leum Geologists, Paciﬁc Section.

—— 1958, Correlation sections longitudinally north-south thru [sic]
central San Joaquin Valley from Rio Vista thru [sic] Riverdale (10
North) and Riverdale thru [sic] Tejon Ranch area (10 South),
California: American Association of Petroleum Geologists, Paciﬁc
Section.

— 1959, Correlation section longitudinally north-south thru [sic]
westside San Joaquin Valley from Coalinga to Midway Sunset and
across San Andreas fault into southeast Cuyama Valley: American
Association of Petroleum Geologists, Paciﬁc Section.

Clark, J.C., Brabb, E.E., Greene, H.G., and Ross, D.G., 1984, Geology
of Point Reyes Peninsula and implications for San Gregorio fault
history, in Crouch, J .K., and Bachman, S.B., eds., Tectonics and
sedimentation along the California margin: Society of Economic
Paleontologists and Mineralogists, Paciﬁc Section, Annual Meet-
ing, 1984, San Diego, Calif, p. 67—86.

Clarke, S.H., Jr., 1973, The Eocene Point of Rocks Sandstone—Prov-
enance, mode of deposition, and implications for the history of
offset along the San Andreas fault in central California: Berkeley,
University of California, Ph.D. thesis, 302 p. ’

Clarke, S.H., Jr., Howell, D.G., and Nilsen, T.H., 1975, Paleogene
geography of California, in Weaver, D.W., Hornaday, G.R., and
Tipton, Ann, eds., Paleogene symposium and selected technical
papers: American Association of Petroleum Geologists-Society of
Economic Paleontologists and Mineralogists—Society of Explora-
tion Geophysicists, Paciﬁc Section, Annual Meeting, 1975, Long
Beach, Calif., p. 121—154.

Clarke, S.H., Jr., and Nilsen, T.H., 1973, Displacement of Eocene
strata and implications for the history of offset along the San
Andreas fault, central and northern California, in Kovach, R.L.,
and Nur, Amos, eds., Proceedings of the conference on tectonic

 

problems of the San Andreas fault: Stanford University Publica-
tions, Geological Sciences, v. 13, p. 358—367.

Cole, M.R., and Armentrout, J .M., 1979, Neogene paleogeography of
the western United States, in Armentrout, J .M., Cole, M.R., and
TerBest, H., Jr., eds., Cenozoic paleogeography of the western
United States: Society of Economic Paleontologists and Mineral-
ogists, Paciﬁc Section, Annual Meeting 1979, Anaheim, Calif,
Paciﬁc Coast Paleogeography Symposium 3, p. 297—323.

Coney, P.J., and Reynolds., S.J., 1977, Cordilleran Benioff zones:
Nature, v. 270, p. 403—406.

Cooley, S.A., 1982, Depositional environments of the lower and middle
Miocene Temblor Formation of Reef Ridge, Fresno and Kings
Counties, California, in Williams, L.A., and Graham, S.A. eds.,
Monterey Formation and associated coarse elastic rocks, central
San Joaquin basin, California: Society of Economic Paleontologists
and Mineralogists, Paciﬁc Section, Fall 1982 Field trip, Coalinga,
Ca1if., Guidebook, p. 55—72.

Cox, Allan, and Engebretson, David, 1985, Change in motion of Paciﬁc
plate at 5 Myr BP: Nature, v. 313, p. 472—474.

Croft, MG, 1972, Subsurface geology of the late Tertiary and
Quaternary water-bearing deposits of the southern part of the San
Joaquin Valley, California: US. Geological Survey Water-Supply
Paper 1999—H, 29 p., 16 plates.

Croft, MG, and Gordon, G.V., 1968, Geology, hydrology, and quality
of water in the Hanford—Visalia area, San Joaquin Valley, Califor—
nia: US Geological Survey open-ﬁle report, 63 p.

Cross, T.A., and Pilger, R.H., Jr., 1978, Constraints on absolute
motion and plate interaction: American Journal of Science, v. 278,
p. 865—902.

Crough, ST, and Thompson, G.A., 1977, Upper mantle origin of the
Sierra Nevada uplift: Geology, v. 5, p. 396—399.

Crowell, J .C., 1987, Late Cenozoic basins of onshore southern Califor-
nia: complexity is the hallmark of their tectonic history, in Ernst,
W.G., and Ingersoll, R.V., eds., Cenozoic basin development of
coastal California (Rubey volume 6): Englewood Cliffs, N.J.,
Prentice-Hall, p. 207—241.

Dalrymple, GB, 1964, Cenozoic chronology of the Sierra Nevada,
California: University of California, Publications in Geological
Sciences, V. 47, 41 p.

— 1979, Critical tables for conversion of K-Ar ages from old to new
constants: Geology, v. 7, p. 558—560.

Davis, G.A., and Burchﬁel, B.C., 1973, Garlock fault: An intraconti-
nental transform structure, southern California: Geological Soci-
ety of America Bulletin, v. 84, p. 1407-1422.

Davis, T.L., 1983, Late Cenozoic structure and tectonic history of the
western “Big Bend” of the San Andreas fault and adjacent San
Emigdio Mountains: Santa Barbara, University of California,
Ph.D. thesis, 580 p.

— 1986, A structural outline of the San Emigdio Mountains, in
Davis, T.L., and Namson, J .S., eds., Geologic transect across the
western Transverse Ranges: Society of Economic Paleontologists
and Mineralogists, Paciﬁc Section, Annual Meeting, 1986, Bakers-
ﬁeld, Calif., Field Trip Guidebook, p. 23—32.

DeCelles, P.G., 1986, Middle Tertiary depositional systems of the San
Emidgio Range, southern California: Society of Economic Paleon-
tologists and Mineralogists, Paciﬁc Section, Annual Meeting,
1986, Bakersﬁeld, Calif, Field Trip Guidebook, 32 p.

Dibblee, T.W., Jr., 1973a, Regional geologic map of San Andreas and
related faults in Carrizo Plain, Temblor, Caliente, and La Panza
Ranges and vicinity, California: US. Geological Survey Miscella-
neous Geologic Investigations Map 1—757, scale 1:125,000.

— 1973b, Stratigraphy of the southern Coast Ranges near the San
Andreas fault from Cholame to Maricopa, California: US. Geo-
logical Survey Professional Paper 764, 45 p.

— 1979a Geologic map of the central Diablo Range between Hollis-

36 THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

ter and New Idria, San Benito, Merced, and Fresno Counties,

California: U.S. Geological Survey Open-File Report 79-358, scale

1262,500.

1979b, Middle and late Cenozoic deposition and tectonics of the
central Diablo Range between Hollister and New Idria, in Nilsen,

T.H., and Dibblee, T.W., Jr., eds., Geology of the central Diablo

Range between Hollister and New Idria, California: Geological

Society of America, Cordilleran Section, April 9—11, 1979, Field
Trip Guidebook, p. 56—65.

Dibblee, T.W., Jr., and Chesterman, C.W., 1953, Geology of the
Breckenridge Mountain quadrangle: California Division of Mines
and Geology Bulletin 168, 56 p.

Dickinson, W.R., Armin, R.A., Bechvar, Nancy, Goodlin, T.C.,
Janecke, S.U., Mark, R.A., Norris, R.D., Radel, Gregg, and
Wortman, A.A., 1987, Geohistory analysis of rates of sediment
accumulation and subsidence for selected California basins, in
Ernst, W.G., and Ingersoll, R.V., eds., Cenozoic basin develop-
ment of coastal California (Rubey volume 6): Englewood Cliffs,
N.J., Prentice-Hall, p. 1—23.

Dickinson, W.R., Ingersoll, R.V., and Graham, S.A., 1979, Paleogene
sediment dispersal and paleotectonics in northern California:
Geological Society of America Bulletin, pt. 2, v. 90, p. 1458—1528.

Dickinson, W.R., and Seely, DR, 1979, Structure and stratigraphy of
forearc regions: American Association of Petroleum Geologists
Bulletin, v. 63, p. 2—31.

Dickinson, W.R., and Snyder, W.S., 1979, Geometry of triple junctions
related to San Andreas transform: Journal of Geophysical Re-
search, v. 84, no. B2, p. 561—572.

Donovan, D.T., and Jones, E.J.W., 1979, Causes of world-wide changes
in sea level: Journal of Geological Society of London, v. 136, p.
187—192.

Drinkwater, J .L., 1983, Geology of the northeastern part of the Quien
Sabe Volcanics, Merced County, California: San Jose, Calif., San
Jose State University, MS. thesis, 102 p.

Dunwoody, J .A., 1969, Correlation section 17, San Joaquin Valley,
Kingsburg to Tejon Hills: American Association of Petroleum
Geologists, Paciﬁc Section.

Eaton, GR, 1979, Regional geophysics, Cenozoic tectonics, and
geologic resources of the Basin and Range Province and adjoining
regions, in Newman, G.W., and Goode, H.D., eds., Basin and
Range Symposium and Great Basin Field Conference: Rocky
Mountain Association of Geologists and Utah Geological Associa-
tion, p. 11—39.

Eaton, J .P., 1985a, Regional seismic background of the May 2, 1983

Coalinga earthquake, in Rymer, M.J., and Ellsworth, W.L., eds.,

Mechanics of the May 2, 1983 Coalinga earthquake: U.S. Geolog-

ical Survey Open-File Report 85—44, p. 44—60.

1985b, The May 2, 1983 Coalinga earthquake and its aftershocks:

A detailed study of the hypocenter distribution and of the focal

mechanisms of the larger aftershocks, in Rymer, M.J., and

Ellsworth, W.L., eds., Mechanics of the May 2, 1983 Coalinga

earthquake: U.S. Geological Survey Open-File Report 85—44, p.

132—201.

Engebretson, D.C., Cox, Allan, and Gordon, R.G., 1985, Relative
motions between oceanic and continental plates in the Paciﬁc
basin: Geological Society of America Special Paper 206, 59 p.

Evernden, J.F., Savage, D.E., Curtis, G.H., and James, G.T., 1964,
Potassium-argon dates and the Cenozoic Mammalian chronology of
North America: American Journal of Science, v. 262, p. 145—198.

Fischer, P.J., 1984, An early Paleogene submarine canyon and fan
system: The Meganos Formation, southern Sacramento basin,
California, in Cherven, Victor, and Fischer, P.J., eds., Latest
Cretaceous and early Tertiary depositional systems of the north-
ern Diablo Range, California: Society of Economic Paleontologists
and Mineralogists, Midyear Meeting, 1984, San Jose, Calif., Field

 

 

 

Trip Guidebook No. 3, p. 53—64.

Foss, C.D., 1972, A preliminary sketch of the San Joaquin Valley
stratigraphic framework, in Rennie, E.W., ed., Geology and oil
ﬁelds—west side central San Joaquin Valley: American Associa-
tion of Petroleum Geologists-Society of Economic Paleontologists
and Mineralogists—Society of Exploration Geophysicists, Paciﬁc
Sections, 1972 Guidebook, p. 40-50.

Fox, K.E. Jr., Fleck, R.J., Curtis, G.H., and Meyer, C.H., 1985,
Implications of the northwestwardly younger age of the volcanic
rocks of west-central California: Geological Society of America
Bulletin, v. 96, p. 647—654.

Frakes, L.A., 1979, Climates throughout geologic time: Amsterdam,
Elsevier, 310 p.

Frink, J.W., and Kues, H.A., 1954, Corcoran clay—a Pleistocene
lacustrine deposit in San Joaquin Valley, California: American
Association of Petroleum Geologists Bulletin, v. 38, p. 2357—2371.

Fritsche, A.E., 1977, Miocene paleogeography of California, in Late
Mesozoic and Cenozoic sedimentation and tectonics in California:
San Joaquin Geological Society, Short Course, Bakersﬁeld, Feb.
1977, p. 109—119.

Gale, H.S. Piper, AM, and Thomas, HE, 1939, Geology, in Piper,
A.M., Gale, H.S., Thomas, HE, and Robinson, T.W., Geology
and ground-water hydrology of the Mokelumne area, California:
U.S. Geological Survey Water-Supply Paper 780, p. 14—101.

Galehouse, J.S., 1967, Provenance and paleocurrents of the Paso
Robles Formation, California: Geological Society of America
Bulletin, v. 78, p. 951—978.

Garfunkel, Zvi, 1973, History of the San Andreas fault as a plate
boundary: Geological Society of America Bulletin, v. 84, p.
2035—2042.

Graham, S.A., 1978, Role of Salinian block in evolution of San Andreas
fault system: American Association of Petroleum Geologists Bul-
letin, v. 62, p. 2214—2231.

Graham, S.A., and Berry, K.D., 1979, Early Eocene paleogeography of
the central San Joaquin Valley: Origin of the Cantua Sandstone, in
Armentrout, J.M., Cole, MR, and TerBest, Harry, Jr., eds.,
Cenozoic paleogeography of the western United States: Society of
Economic Paleontologists and Mineralogists, Paciﬁc Section, An-
nual Meeting, 1979, Anaheim, Calif., Paciﬁc Coast Paleogeogra—
phy Symposium 3, p. 119—127.

Graham, S.A., and Dickinson, W.R., 1978, Evidence for 115 kilometers
of right slip on the San Gregorio-Hosgri fault trend: Science, v.
199, p. 179—181.

Graham, S.A., Stanley, R.G., Bent, J.V., and Carter, J.B., 1986,
Record of initial tectonism and total right-slip along the San
Andreas fault from Oligo-Miocene strata, Santa Cruz Mountains
and Temblor Range [abs]: Geological Society of America, Cordil-
leran Section, Abstracts with Programs, v. 18, no. 2, p. 110.

Graham, S.A., Williams, L.A., Bate, MA, and Weber, LS, 1982,
Stratigraphic and depositional framework of the Monterey For-
mation and associated coarse clastics of the central San Joaquin
basin, in Williams, L.A., and Graham, S.A., eds., Monterey
Formation and associated coarse clastic rocks, central San Joaquin
basin, California: Society of Economic Paleontologists and Miner-
alogists, Paciﬁc Section, Fall 1982 Field Trip Coalinga, Calif.,
Guidebook, p. 3—16.

Grant, T.A., McCleary, J .R., and Blum, R.L., 1977, Correlation and
dating of geomorphic and bedding surfaces on the east side of the
San Joaquin Valley using dip, in Singer, M.J., ed., Soil develop-
ment, geomorphology and Cenozoic history of the northeastern
San Joaquin Valley and adjacent areas, California: American
Society of Agronomy, Geological Society of America, Soil Science
Society of America, Joint ﬁeld section, 1977, Guidebook, Davis,
Univ. of Calif. Press, p. 312-218.

Greene, H.G., and Clarke, J .C., 1979, N eogene paleogeography of the

REFERENCES CITED 37

Monterey Bay area, California, in Armentrout, J .M., Cole, M.R.,
and TerBest, H., J r., eds., Cenozoic paleogeography of the
western United States: Society of Economic Paleontologists and
Mineralogists, Paciﬁc Section, Annual Meeting, 1979, Anaheim,
Calif, Paciﬁc Coast Paleogeography Symposium 3, p. 277—296.

Hackel, Otto, 1966, Summary of the geology of the Great Valley, in
Bailey, E.H., ed., Geology of Northern California: California
Division of Mines and Geology Bulletin 190, p. 217—238.

Haq, B.U., Hardenbol, Jan, and Vail, P.R., 1987, The new chrono-
stratigraphic basis of Cenozoic and Mesozoic sea level cycles, in
Ross, C.A., and Haman, Drew, eds., Timing and depositional
history of eustatic sequences: constraints on seismic stratigraphy:
Cushman Foundation for Foraminiferal Research, Special Publi-
cation 24, p. 7—13. '

Harding, TR, 1976, Tectonic signiﬁcance and hydrocarbon trapping
consequences of sequential folding synchronous with San Andreas
faulting, San Joaquin Valley, California: American Association of
Petroleum Geologists Bulletin, v. 60, p. 356—378.

Hay, E.A., 1976, Cenozoic uplifting of the Sierra Nevada in isostatic
response to the North American and Paciﬁc plate interactions:
Geology, v. 4, p. 763—766.

Heikkila, H.H., and MacLeod, C.M., 1951, Geology of Bitterwater
Creek area, Kern County, California: California Division of Mines
and Geology Special Report 6, 21 p.

Herd, D.G., 1979a, Neotectonics of the San Francisco Bay region,
California, in Summaries of technical reports, v. 8, prepared by
participants in National earthquake hazards reduction program:
Menlo Park, US. Geological Survey, p. 29—31.

— 1979b, The San Joaquin fault zone—evidence for late Quaternary
faulting along the west side of the northern San Joaquin Valley,
California [abs.]: Geological Society of America, Cordilleran Sec—
tion, Abstracts with Programs, v. 11, no. 3, p. 83.

Hill, DR, 1982, Contemporary block tectonics: California and Nevada:
Journal of Geophysical Research, v. 87, no. B7, p. 5433-5450.
Hirst, Brian, 1986, Tectonic development of the Tejon and adjacent

areas, Kern County, California, in Structure and stratigraphy of

the east side San Joaquin Valley: American Association of Petro-

leum Geologists, Paciﬁc Section, Southeast San Joaquin Valley
, Field Trip, April 18—19, 1986, Guidebook, p. 2—8.

Hluza, A.G. , 1960, Northeast area of Wheeler Ridge oil ﬁeld: California
Division of Oil and Gas, Summary of Operations—California Oil
Fields, v. 46, no. 2, p. 61—67.

Hodges, C.A., 1979, Preliminary maps of photolineaments along parts
of the western Sierra Nevada foothills and eastern Coast Range
foothills, California, based on Landsat images and U—2 aircraft
photographs: US. Geological Survey Open-File Report 79—1470, 8

p.

Hoffman, R.D., 1964, Geology of the northern San Joaquin: San Joaquin
Geological Society Selected Papers, v. 2, p. 30—45.

Holzer, T.L., 1980, Faulting caused by groundwater declines, San
Joaquin Valley, California: Water Resources Research, v. 16, no.
6, p. 1065—1070.

Hoots, H.W., Bear, T.L., and Kleinpell, W.D., 1954, Geological
summary of the San Joaquin Valley, California, in J ahns, R.H.,
ed., Geology of southern California: California Division of Mines
and Geology Bulletin 170, p. 113-129.

Howell, D.G., Champion, DE, and Vedder, J .G., 1987, Terrane
accretion, crustal kinematics, and basin evolution, southern Cali-
fornia, in Ingersoll, R.V., and Ernst, W.G., eds., Cenozoic basin
development of coastal California (Rubey volume 6): Englewood
Cliffs, N.J., Prentice-Hall, p. 242—258.

Howell, D.G., Crouch, J .K., Greene, H.G., McCulloch, BS, and
Vedder, J .G., 1980, Basin development along the late Mesozoic
and Cenozoic California margin: a plate tectonic margin of sub-
duction, oblique subduction and transform tectonics, in Ballance,

 

P.F., and Reading, H.G., eds., Sedimentation in oblique-slip
mobile zones: International Association of Sedimentologists Spe-
cial Publication 4, p.43-62.

Huber, N.K., 1981, Amount and timing of late Cenozoic uplift and tilt
of the central Sierra Nevada, California—evidence from the upper
San Joaquin basin: US. Geological Survey Professional Paper
1197, 28 p.

Huffman, 0.F. , 1972, Lateral displacement of upper Miocene rocks and
the Neogene history of offset along the San Andreas fault in
central California: Geological Society of America Bulletin, v. 83, p.
2913—2946.

Ingersoll, R.V., 1979, Evolution of the Late Cretaceous forearc basin,
northern and central California: Geological Society of America
Bulletin, pt. 1, v. 90, p. 813—826.

-—— 1982, Triple-junction instability as cause for late Cenozoic exten-
sion and fragmentation of the western United States: Geology, v.
10, p. 621-624.

Ingle, J .C., 1981, Origin of Neogene diatomites around the north Paciﬁc
rim, in Garrison, RE, and Douglas, R.G., eds., The Monterey
Formation and related siliceous rocks of California: Society of
Economic Paleontologists and Mineralogists, Paciﬁc Section, An-
nual Meeting, 1981, San Francisco, Calif, p. 159—179.

Jennings, C.W., 1977, Geologic map of California: California Division of
Mines and Geology, scale 1:750,000.

Johnson, C.M., and O’Neil, J .R., 1984, Triple junction magmatism: a
geochemical study of Neogene volcanic rocks in western Califor-
nia: Earth and Planetary Science Letters, v. 71, p. 241—262.

Kanter, L.R., and McWilliams, MO, 1982, Rotation of the southern-
most Sierra Nevada, California: Journal of Geophysical Research,
v. 87, no. B5, p. 3819—3830.

Keith, SB, 1978, Paleosubduction geometries inferred from Creta-
ceous and Tertiary magmatic patterns in southwestern North
America: Geology, v. 6, p. 516—521.

Kuespert, J .G., 1983, The depositional environment of the Miocene
Temblor Formation and associated Oligo-Miocene units in the
vicinity of Kettleman North Dome, San Joaquin Valley, California:
Stanford, Calif, Stanford University, MS thesis, 105 p.

— 1985, Depositional environments and sedimentary history of the
Miocene Temblor Formation and associated Oligo-Miocene units in
the vicinity of Kettleman North Dome, San Joaquin Valley,
California, in Graham, S.A., ed., Geology of the Temblor Forma-
tion, western San Joaquin basin, California: Society of Economic
Paleontologists and Mineralogists, Paciﬁc Section, Fall 1985 Field
Trip, Coalinga, Calif, Guidebook, p. 53—67.

LaForge, R., and Lee, W.H.K., 1982, Seismicity and tectonics of the
Ortigalita fault and southeast Diablo Range, California, in Hart,
E.W., Hirschfeld, SE, and Schulz, S.S., eds., Proceedings of
conference on earthquake hazards in the eastern San Francisco
Bay region: California Division of Mines and Geology Special
Publication 62, p. 93—101.

Lagoe, MB, 1986, Stratigraphic nomenclature and time-rock relation-
ships in the Paleogene rocks of the San Emigdio Mountains, in
Davis, T.L., and Namson, J .S., eds., Geologic transect across the
western Transverse Ranges: Society of Economic Paleontologists
and Mineralogists, Paciﬁc Section, Annual Meeting, 1986, Bakers-
ﬁeld, Calif., Field Trip Guidebook, p. 11—21.

Lettis, W.R., 1982, Late Cenozoic stratigraphy and structure of the
western margin of the central San Joaquin Valley, California: U. S.
Geological Survey Open-File Report 82—526, 26 plates, scale
1:24,000.

— 1985, Late Cenozoic stratigraphy and structure of the west
margin of the central San Joaquin Valley, California, in Weide,
D.L. ed., Soils and Quaternary geology of the southwestern
United States: Geological Society of America Special Paper 203, p.
97—114.

38 THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

Lipman, P.W., Prostka, H.J., and Christiansen, R.L., 1972, Cenozoic
volcanism and plate-tectonic evolution of the western United
States. 1. Early and middle Cenozoic: Philosophical Transactions
of the Royal Society of London, Series A, v. 271, p. 217—248.

Loomis, DR, and Glazner, A.F., 1986, Middle Miocene tectonic uplift
of southern San Joaquin basin, California: American Association of
Petroleum Geologists Bulletin, v. 70, p. 1003—1007.

Los Angeles (City) Department of Water and Power, 1975, San Joaquin
Nuclear Project, Early Site Review Report: Los Angeles, Depart-
ment of Water and Power, report submitted to the U.S. Nuclear
Regulatory Commission, 5 vols.

MacPherson, B.A., 1978, Sedimentation and trapping mechanism in
upper Miocene Stevens and older turbidite fans of southeastern
San Joaquin Valley, California: American Association of Petro—
leum Geologists Bulletin, v. 62, p. 2243—2274.

McWilliams, Michael, and Li, Yianping, 1985, Oroclinal bending of the
southern Sierra Nevada batholith: Science, v. 230, p. 172—175.

Maddock, M.E., 1964, Geology of the Mt. Boardman quadrangle, Santa
Clara and Stanislaus Counties, California: California Division of
Mines and Geology Map Sheet 3, scale 1:62,500.

Maher, J .C., Carter, RD, and Lantz, R.J., 1975, Petroleum geology
of Naval Petroleum Reserve No. 1, Elk Hills, Kern County,
California: U.S. Geological Survey Professional Paper 912, 109 p.

Marchand, DE, 1977, The Cenozoic history of the San Joaquin Valley
and adjacent Sierra Nevada as inferred from the geology and soils
of the eastern San Joaquin Valley, in Singer, M.J., ed., Soil
development, geomorphology and Cenozoic history of the north-
eastern San Joaquin Valley and adjacent areas, California: Amer-
ican Society of Agronomy, Geological Society of America, Soil
Sciences Society of America, Joint ﬁeld section 1977, Guidebook,
Davis, Univ. of Calif. Press, p. 39—50.

Marchand, DE, and Allwardt, Alan, 1978, Preliminary geologic map
showing Quaternary deposits of the northeastern San Joaquin
Valley in parts of Stanislaus, Merced, Madera, Fresno, and
Mariposa Counties, California: U.S. Geological Survey Miscella-
neous Field Studies Map MF-945, scale 1:125,000.

— 1981, Late Cenozoic stratigraphic units, northeastern San Joa-
quin Valley, California: U.S. Geological Survey Bulletin 1470, 70

p.

Mathews, R.K., 1984, Oxygen isotope record of ice-volume history: 100
million years of glacio-eustatic sea-level ﬂuctuation, in Schlee,
J .S., ed., Interregional unconformities and hydrocarbon accumu—
lation: American Association of Petroleum Geologists Memoir 36,
p. 97—107.

Mavko, B.B., and Thompson, GA, 1983, Crustal and upper mantle
structure of the northern and central Sierra Nevada: Journal of
Geophysical Research, v. 88, no. B7, p. 5874—5892.

May, J.A., Yeo, R.K., and Warme, J .E., 1984, Eustatic control on
synchronous stratigraphic development: Cretaceous and Eocene
coastal basins along an active margin: Sedimentary Geology, v. 40,
p. 131—149.

Medwedeff, D.A., and Suppe, John, 1986, Growth fault—bend folding—
precise determination of kinematics, timing, and rates of folding
and faulting from syntectonic sediments [abs]: Geological Society
of America Abstracts with Programs, v. 18, no. 6, p. 692.

Miall, AD, 1986, Eustatic sea level changes interpreted from seismic
stratigraphy: a critique of the methodology with particular refer-
ence to the North Sea Jurassic record: American Association of
Petroleum Geologists Bulletin, v. 70, p. 131—137.

Milam, R.W., 1985, Biostratigraphy and sedimentation of the Eocene
and Oligocene Kreyenhagen Formation, central California: Stan-
ford, Calif., Stanford University, Ph.D. thesis, 240 p.

Miller, R.E., Green, J.H., and Davis, G.H., 1971, Geology of the
compacting deposits in the Los Banos-Kettleman City subsidence
area, California: U.S. Geological Survey Professional Paper

 

497—13, 46 p.

Miller, RH, and Bloom, C.V., 1937, Mountain View oil ﬁeld: California
Division of Oil and Gas, Summary of Operations—California Oil
Fields, v. 22, no. 4, p. 5—36.

Miller, T.L., 1987, Early Neogene coccolith biostratigraphy of R.M.
Kleinpell’s original stratotype section—Reliz Canyon, Monterey
County, California [abs]: Abstract volume, International Con-
gress on Neogene Stratigraphy, 4th, July 29—31, 1987, Berkeley,
Calif, p. 77-78.

Minster, J .B., and Jordan, T.H., 1984, Vector constraints on Quater-
nary deformation of the western United States east and west of
the San Andreas fault, in Crouch, J .K., and Bachman, S.B., eds.,
Tectonics and sedimentation along the California margin: Society
of Economic Paleontologists and Mineralogists, Paciﬁc Section,
Annual Meeting, 1984, San Diego, Calif, p. 1—16.

Moore, D.G., and Curray, J .R., 1982, Geological and tectonic history of
the Gulf of California, in Curray, J.R., and Moore, D.G., and
others, Initial Reports of the Deep Sea Drilling Project, v. 64, pt.
2: Washington, U.S. Government Printing Ofﬁce, p. 1279—1294.

Moore, J .G., and Dodge, F.C.W., 1980, Late Cenozoic volcanic rocks of
the southern Sierra Nevada, California: I. Geology and petrology:
Summary: Geological Society of America Bulletin, pt. 1, v. 91, p.
515—518.

Moxon, I.W., 1986, Subsidence history of the San J oaquin-Sacramento
Valleys: Forearc evolution coupled to convergent margin tectonics
[abs]: International Sedimentological Congress, 12th, 24—30 Au-

, gust, 1986, Canberra, Australia, Abstracts, p. 220.

Namson, J.S., and Davis, T.L., 1984, Deformational history, thrust-
belt structural styles and plate tectonic origin of Coast Range
structures along the San Joaquin Valley, California [abs]: Geo-
logical Society of America Abstracts with Programs, v. 16, no. 6,
p. 607.

Nilsen, T.H., 1979, Early Cenozoic stratigraphy, tectonics and sedi-
mentation of the central Diablo Range between Hollister and New
Idria, in Nilsen, T.H., and Dibblee, J .W., J r., eds., Geology of the
central Diablo Range between Hollister and New Idria, California:
Geological Society of America, Cordilleran Section, April 9—11,
1979, Field Trip Guidebook, p. 31—55.

— 1984, Oligocene tectonics and sedimentation, California: Sedi-
mentary Geology, v. 37, p. 305—336.

—— 1987, Paleogene tectonics and sedimentation of coastal California,
in Ingersoll, R.V., and Ernst, W.G., eds., Cenozoic basin devel-
opment of coastal California (Rubey volume 6): Englewood Cliffs,
N.J., Prentice—Hall, p. 81—123.

Nilsen, T.H., and Clarke, S.H., Jr., 1975, Sedimentation and tectonics
in the early Tertiary continental borderland of central California:
U.S. Geological Survey Professional Paper 925, 64 p.

Nilsen, T.H., Dibblee, T.W., Jr., and Addicott, W.O., 1973, Lower and
middle Tertiary stratigraphic units of the San Emidgio and
western Tehachapi Mountains, California: U.S. Geological Survey
Bulletin 1372—H, 23 p.

Nilsen, T.H., and McKee, E.H., 1979, Paleogene paleogeography of the
western United States, in Armentrout, J.M., Cole, MR, and
TerBest, H., Jr., eds., Cenozoic paleogeography of the western
United States: Society of Economic Paleontologists and Mineral-
ogists, Paciﬁc Section, Annual Meeting, 1979, Anaheim, Calif,
Paciﬁc Coast Paleogeography Symposium 3, p. 257—276.

Oakeshott, G.B., 1955, The Kern County earthquakes in California’s
geologic history, in Oakeshott, G.B., ed., Earthquakes in Kern
County, California during 1952: California Division of Mines and
Geology Bulletin 171, p. 15—22.

Obradovich, J .D., and Naeser, C.W., 1981, Geochronology bearing on
the age of the Monterey Formation and siliceous rocks in Califor-
nia, in Garrison, RE, and Douglas, R.G., eds., The Monterey
Formation and related siliceous rocks in California: Society of

REFERENCES CITED 39

Economic Paleontologists and Mineralogists, Paciﬁc Section, An-
nual Meeting, 1981, San Francisco, Calif, p. 87—95.

Olson, H.C., Miller, G.E., and Bartow, J.A., 1986, Stratigraphy,
paleoenvironment and depositional setting of Tertiary sediments,
southeastern San Joaquin basin, in Structure and stratigraphy of
the east side San Joaquin Valley: American Association of Petro-
leum Geologists Paciﬁc Section, Southeast San Joaquin Valley
Field Trip, April 18-19, 1986, Guidebook p. 18—56.

Page, B.M., 1981, The southern Coast Ranges, in Ernst, W.G., ed.,
The geotectonic development of California (Rubey volume 1):
Englewood Cliffs, N.J., Prentice-Hall, p. 329—417.

Page, B.M., and Engebretson, DC, 1984, Correlation between the
geologic record and computed plate motions for central California:
Tectonics, v. 3, p. 133-155.

Page, R.W., 1986, Geology of the fresh ground-water basin of the
Central Valley, California: U.S. Geological Survey Professional
Paper 1401—0, 54 p.

Page, R.W., and LeBlanc, R.A., 1969, Geology, hydrology, and water
quality in the Fresno area, California: U.S. Geological Survey
Open-File Report, 193 p.

Palmer, GM, and Merrill, R.D., 1982, Braided-stream and alluvial-fan
depositional environments in the lower to middle Eocene Ione
Formation, Madera County, California, in Ingersoll, R.V., and
Woodburne, M.O., eds. , Cenozoic nonmarine deposits of California
and Arizona: Society of Economic Paleontologists and Mineralo-
gists, Paciﬁc Section, Annual Meeting, 1982, Anaheim, Calif, p.
1—10.

Parkinson, Neil, and Summerhayes, Colin, 1985, Synchronous global
sequence boundaries: American Association of Petroleum Geolo—
gists Bulletin, v. 69, p. 685—687.

Pence, J .J ., 1985, Sedimentology of the Temblor Formation in the
northern Temblor Range, California, in Graham, S.A., ed.,
Geology of the Temblor Formation, western San Joaquin basin,
California: Society of Economic Paleontologists and Mineralogists,
Pacific Section, Fall 1985 Field Trip, Coalinga, Calif, Guidebook,
p. 19—34.

Phillips, F.J., Tipton, Ann, and Watkins, Rodney, 1974, Outcrop
studies of the Eo—Oligocene Tumey Formation, Monocline Ridge,
Fresno County, California, in Homaday, G.R., ed., The Paleo—
gene of the Panoche Creek-Cantua Creek area, central California:
Society of Economic Paleontologists and Mineralogists, Paciﬁc
Section, Fall 1974 Field Trip Guidebook, p. 38-68.

Phillips, R.L., 1981, Depositional environments of the Santa Margarita
Formation in the Santa Cruz Mountains, California: Santa Cruz,
University of California, Ph.D. thesis, 358 p.

Pittman, W.C., III, 1978, Relationship between eustacy and strati-
graphic v. 89, p. 1389—1403.

Plescia, J .B., and Calderone, G.J., 1986, Paleomagnetic constraints on
the timing and extent of rotation of the Tehachapi Mountains,
California [abs]: Geological Society of America, Cordilleran Sec-
tion, Abstracts with Programs, v. 18, no. 2, p. 171.

Poag, C.W., 1984, Unconformities and sedimentary sequences on
passive continental margins [abs]: testing the Vail depositional
model: Geological Society of America Abstracts with Programs,
1984, v. 16, no. 6, p. 624.

Poag, C.W., and Schlee, J .S., 1984, Depositional sequences and
stratigraphic gaps on submerged United States Atlantic margin,
in Schlee, J .S. , ed. , Interregional unconformities and hydrocarbon
accumulation: American Association of Petroleum Geologists
Memoir 36, p. 165—182.

Poag, C.W., and Ward, L.W., 1987, Cenozoic unconformities and
depositional supersequences of the North Atlantic continental
margins: testing the Vail model: Geology, v. 15, p. 159—162.

Poland, J .F., and Evenson, RH, 1966, Hydrogeology and land
subsidence, Great Central Valley, California, in Bailey, E.H., ed.,

 

Geology of northern California: California Division of Mines and
Geology Bulletion 190, p. 239—247.

Poore, R.Z., 1980, Age and correlation of California Paleogene benthic
foraminiferal stages: U.S. Geological Survey Professional Paper
1162—0, 8 p.

Poore, R.Z., Barron, J .A., and Addicott, W.O., 1981, Biochronology of
the northern Paciﬁc Miocene: Proceedings of IGCP 114 Interna-
tional Workshop on Paciﬁc Neogene Biostratigraphy, 6th Inter-
national Working Group Meeting, November 25—29, 1981, Osaka
Museum of Natural History, Osaka, p. 91—98.

Prowell, DO, 1974, The geology of selected Tertiary volcanics in the
central coast mountains of California and their bearing on the
Calaveras and Hayward fault problems: Santa Cruz, University of
California, Ph.D. thesis, 182 p.

Raymond, L.A., 1969, The stratigraphic and structural geology of the
northern Lone Tree Creek and southern Tracy quadrangles: San
Jose, Calif, San Jose State College, M.S. thesis, 143 p.

Reed, R.D., 1933, Geology of California: Tulsa, Oklahoma, American
Association of Petroleum Geologists, 355 p.

Reed, R.D., and Hollister, J .S., 1936, Structural evolution of southern
California: Tulsa, Oklahoma, American Association of Petroleum
Geologists, 157 p.

Rentschler, MS, 1985, Structurally controlled Neogene sedimentation
in the Vallecitos syncline, in Graham, S.A. ed., Geology of the
Temblor Formation, western San Joaquin basin, California: Soci-
ety of Economic Paleontologists and Mineralogists, Paciﬁc Sec-
tion, Fall 1985 Field Trip, Coalinga, Calif, Guidebook, p. 87—96.

Repenning, C.A., 1960, Geologic summary of the Central Valley of
California with reference to the disposal of liquid radioactive
waste: U.S. Geological Survey Trace Elements Investigation
Report 769, 69 p.

Sarna—Wojcicki, A.M., Meyer, C.E., Bowman, H.R., Hall, N.T.,
Russell, P.C. Woodward, M.J., and Slate, J .L., 1985, Correlation
of the Rockland ash bed, a 400,000-year-old stratigraphic marker
in northern California and western Nevada, and implications for
middle Pleistocene paleogeography of central California: Quater-
nary Research, v. 23, p. 236—257.

Schlanger, S.O., Jenkyns, HO, and Premoli—Silva, Isabella, 1981,
Volcanism and vertical tectonics in the Paciﬁc basin related to
global Cretaceous transgressions: Earth and Planetary Science
Letters, v. 52, p. 435—449.

Schwartz, D.P., Swan, F.H., III, Harpster, R.E., Rogers, T.H., and
Hitchcock, D.E., 1977, Surface faulting potential, in Earthquake
evaluation studies of the Auburn Dam area: San Francisco,
Woodward-Clyde Consultants report for U.S. Bureau of Recla-
mation, 135 p.

Shedd, Solon, 1932, Bibliography of the geology and mineral resources
of California to December 31, 1930: California Division of Mines
Bulletin 104, 376 p.

Slagle, L.P., 1979, Depositional systems and structures of the middle
Eocene Domengine-Yokut Sandstone, Vallecitos, California: Stan-
ford, Calif., Stanford University, MS. thesis, 59 p.

Slemmons, DB, 1966, Cenozoic volcanism of the central Sierra
Nevada, California, in Bailey, E.H., ed., Geology of northern
California: California Division of Mines and Geology Bulletin 190,
p. 199—208.

Snyder, W.S., Dickinson, W.R., and Silberman, M.L., 1976, Tectonic
implications of space-time patterns of Cenozoic magmatism in the
western United States: Earth and Planetary Science Letters, v.
32, p. 91—106.

Stanley, R. G., 1985, Middle Tertiary sedimentation and tectonics of the
La Honda basin, central California: U.S. Geological Survey
Open-File Report 85—596, 263 p.

Stein, R.S., 1985, Evidence for surface folding and subsurface fault slip
from geodetic elevation changes associated with the 1983 Coalin-

40 THE CENOZOIC EVOLUTION OF THE SAN JOAQUIN VALLEY, CALIFORNIA

ga, California, earthquake, in Rymer, M.J., and Ellsworth, W.L.,
eds., Mechanics of the May 2, 1983 Coalinga earthquake: U.S.
Geological Survey Open-File Report 85—44, p. 225—253.

Stein, R.S., and Thatcher, Wayne, 1981, Seismic and aseismic defor-
mation.associated with the 1952 Kern County, California earth-
quake and relationship to the Quaternary history of the White
Wolf fault: Journal of Geophysical Research, v. 86, no. B6, p.
4913-4928.

Stewart, J .H., and Carlson, J.E., 1976, Cenozoic rocks of Nevada:
Nevada Bureau of Mines and Geology, Map 52, 4 sheets, scale
1:1,000,000.

Sullivan, J .C., 1966, Kettleman North Dome oil ﬁeld: California
Division of Oil and Gas, Summary of Operations—California Oil
Fields, v. 52, no. 1, p. 5—21.

Sullivan, J .C., and Weddle, J .R., 1960, Rio Bravo oil ﬁeld: California
Division of Oil and Gas, Summary of Operations—California Oil
Fields, p. 46, no. 1, p. 27—39.

Teitsworth, R.A., 1964, Geology and development of the Lathrop gas
ﬁeld, San Joaquin County, California: San Joaquin Geological
Society Selected Papers, V. 2, p. 19—29.

Turner, D.L., 1970, Potassium-argon dating of Paciﬁc Coast Miocene
foraminiferal stages, in Bandy, O.L., ed., Radiometric dating and
paleontologic zonation: Geological Society of America Special
Paper 124, p. 91-129.

Vail, RR, and Hardenbol, Jan, 1979, Sea-level changes during the
Tertiary: Oceanus, v. 22, no. 3, p. 71—79.

Vail, P.R., Hardenbol, J ., and Todd, R.G., 1984, Jurassic unconform-
ities, chronostratigraphy, and sea-level changes from seismic
stratigraphy and biostratigraphy, in Schlee, J .S., ed., Interre-
gional unconformities and hydrocarbon accumulation: American
Association of Petroleum Geologists Memoir 36, p. 129—144.

Vail, P.R., Mitchum, R.M., Jr., and Thompson, S., III, 1977, Seismic
stratigraphy and global changes of sea level, Part 4: Global cycles
of relative changes of sea level, in Payton, C.E., ed., Seismic
stratigraphy—applications to hydrocarbon exploration: American
Association of Petroleum Geologists Memoir 26, p. 83—97.

Vedder, J .G., 1970, Geologic map of the Wells Ranch and Elkhorn Hills
quadrangles, San Luis Obispo and Kern Counties, California: U.S.
Geological Survey Miscellaneous Geologic Investigations Map
I—585, scale 1:24,000.

Wagner, H.M., 1981, Geochronology of the Mehrten Formation in
Stanislaus County, California: Riverside, University of California,
Ph.D. thesis, 385 p.

Wahrhaftig, Clyde, and Birman, J.H., 1965, The Quaternary of the
Paciﬁc mountain system in California, in Wright, H.E., J r., and
Frey, D.G., eds., The Quaternary of the United States, 7th
INQUA Congress Review Volume: Princeton, Princeton Univer-
sity Press, p. 299-340.

Warne, A.H., 1955, Ground fracture patterns in the southern San
Joaquin Valley resulting from the Arvin-Tehachapi earthquake, in
Oakeshott, D.B., ed., Earthquakes in Kern County, California
during 1952: California Division of Mines and Geology Bulletin 171,
p. 57—66.

Watts, A.B., 1982, Tectonic subsidence, ﬂexure and global changes of
sea level: Nature, v. 297, p. 469—474.

Webb, G.W., 1977, Stevens and earlier Miocene turbidite sands, San
Joaquin Valley, California, in Late Miocene geology and new oil

 

ﬁelds of the southern San Joaquin Valley: American Association of
Petroleum Geologists, Paciﬁc Section, Field Trip Guidebook, p.
37—55.

— 1981, Stevens and earlier Miocene turbidite sandstones, southern
San Joaquin Valley, California: American Association of Petro—
leum Geologists Bulletin, v. 65, p. 438—465.

Wentworth, C.M., 1985, Central California deep crustal study: U.S.
Geological Survey Open-File Report 86—31, p. 98—102.

Wentworth, C.M., Blake, M.C., Jr., Jones, D.L., Walter, A.W., and
Zoback, M.D., 1984, Tectonic wedging associated with emplace-
ment of the Franciscan assemblage, California Coast Ranges, in
Blake, M.C., ed., Franciscan geology of northern California: Los
Angeles, Society of Economic Paleontologists and Mineralogists,
Paciﬁc Section, p. 163—173.

Wentworth, C.M., Walter, A.W., Bartow, J .A., and Zoback, M.D.,
1983, Evidence on the tectonic setting of the 1983 Coalinga
earthquakes from deep reﬂection and refraction proﬁles across the
southeastern end of Kettleman Hills, in Bennett, J .H., and
Sherburne, R.W., eds., The 1983 Coalinga, California earth-
quakes, California Division of Mines and Geology Special Paper 66,
p. 113—126.

Wentworth, C.M., and Zoback, M.D., 1986, An integrated faulting
model for the Coalinga earthquake: A guide to thrust deformation
along the Coast Range-Great Valley boundary [abs.]: Eos, v. 67,
no. 44, p. 1222.

White, RT, 1940, Eocene Yokut Sandstone north of Coalinga,
California: American Association of Petroleum Geologists Bulle-
tin, v. 24, p. 1722—1751.

Wolfe, J .A., 1978, A paleobotanical interpretation of Tertiary climates
in the northern hemisphere: American Scientist, v. 66, p. 694—703.

Wolfe, J .A., and Hopkins, D.M., 1967, Climatic changes recorded by
land ﬂoras of northwestern North America, in Hatai, T., ed.,
Tertiary correlations and climatic change in the Paciﬁc: Proceed-
ings, Paciﬁc Science Congress, 11th, Symposium 25, p. 67—76.

Wong, LG, and Ely, R.W., 1983, Historical seismicity and tectonics of
the Coast Ranges-Sierran block boundary: implications to the 1983
Coalinga, California earthquake, in Bennett, J .H., and Sher-
burne, R.W., eds., The 1983 Coalinga, California earthquakes:
California Division of Mines and Geology Special Publication 66, p.
89—104.

Wong, LG, and Savage, W.U., 1983, Deep intraplate seismicity in the
western Sierra Nevada, central California: Bulletin of Seismolog-
ical Society of America, v. 73, p. 797—812.

Woodring, W.P., Stewart, Ralph, and Richards, R.W., 1940, Geology
of the Kettleman Hills oil ﬁeld, California: U.S. Geological Survey
Professional Paper 195, 170 p.

Zoback, M.L., Anderson, RE, and Thompson, G.A., 1981, Cenozoic
evolution of the state of stress and style of tectonism of the Basin
and Range province of the western United States: Philosophical
Transactions of the Royal Society of London, Series A, v. 300, no.
1454, p. 407—434.

Zoback, M.D., Zoback, M.L., Mount, V.S., Suppe, John, Eaton, J .P.,
Healy, J .H., Oppenheimer, David, Reasenburg, Paul, Jones,
Lucille, Raleigh, C.B., Wong, I.G., Scotti, Oona, and Wentworth,
Carl, 1987, New evidence on the state of stress of the San Andreas
fault system: Science, v. 238, p. 1105—1111.

 

 

AVAILABILITY OF BOOKS AND MAPS OF THE U.S. GEOLOGICAL SURVE

Instructions on ordering publications of the U.S. Geological Survey, along with prices of the last offerings, are given in the t
rent-year issues of the monthly catalog "New Publications of the U.S. Geological Survey." Prices of available U.S. Geological E
vey publications released prior to_the current year are listed in the most recent annual "Price and Availability List." Publicati
that are listed in various U.S. Geological Survey catalogs (see back inside cover) but not listed in the most recent annual "Price

Availability List" are no longer available.

Prices of reports released to the open ﬁles are given in the listing "U.S. Geological Survey Open-File Reports," updated mor
1y, which is for sale in microﬁche from the U.S. Geological Survey, Books and Open-File Reports Section, Federal Center, 1
25425, Denver, CO 80225. Reports released through the NTIS may be obtained by writing to the National Technical Informal
Service, U.S. Department of Commerce, Springﬁeld, VA 22161; please include NTIS report number with inquiry.

Order U.S. Geological Survey publications by mail or over the counter from the ofﬁces given below.

BY MAIL
BOOKS

Professional Papers, Bulletins, Water-Supply Papers, Techniques
of Water-Resources Investigations. Circulars. publications of general in-
terest (such as leaflets, pamphlets, booklets), single copies of Earthquakes
& Volcanoes. Preliminary Determination of Epicenters, and some mis-
cellaneous reports, including some of the foregoing series that have gone
out of print at the Superintendent of Documents, are obtainable by mail
from

U.S. Geological Survey, Books and Open-File Reports
Federal Center, Box 25425
Denver, CO 80225

Subscriptions to periodicals (Earthquakes & Volcanoes and
Preliminary Determination of Epicenters) can be obtained ONLY from
the

Superintendent of Documents
Government Printing Ofﬁce
Washington, D.C. 20402

(Check or money order must be payable to Superintendent of Docu-
merits.)

Maps
For maps. address mail orders to
U.S. Geological Survey, Map Distribution

Federal Center, Box 25286
Denver, CO 80225

Residents of Alaska may order maps from
Alaska Distribution Section, U.S. Geological Survey,

New Federal Building - Box 12
101 Twelfth Ave., Fairbanks, AK 99701

OVER THE COUNTER
Books

Books of the U.S. Geological Survey are available over
counter at the following Geological Survey Public Inquiries Ofﬁces
of which are authorized agents of the Superintendent of Document

- WASHINGTON, D.C.--Main Interior Bldg, 2600 corridor.
18th and C Sts., NW.

- DENVER, Colorado--Federal Bldg, Rm. 169, 1961 Stout S

0 LOS ANGELES, California--Federal Bldg.. Rm. 7638, 300
Los Angeles St.

- MENLO PARK, California-Bldg. 3 (Stop 533). Rm. 3128,
345 Middleﬁeld Rd.

- RESTON, Virginia-603 National Center, Rm. 1C402. 122C
Surn'ise Valley Dr.

0 SALT LAKE CITY, Utah--Federal Bldg., Rm. 8105. 125
South State St.

0 SAN FRANCISCO, California——Customhouse, Rm. 504, 55
Battery St.

- SPOKANE, Washington--U.S. Courthouse, Rm. 678, West
920 Riverside Ave.

- ANCHORAGE, Alaska--Rm. 101, 4230 University Dr.

- ANCHORAGE, Alaska--Federal Bldg, Rm. 13-146. 701 C S

Maps

Maps may be purchased over the counter at the U.S. Geol
cal Survey offices where books are sold (all addresses in above list)
at the following Geological Survey ofﬁces:

- ROLLA, MIssourI--l400 Independence Rd.

- DENVER, Colorado-~Map Distribution. Bldg. 810, Federal
Center

0 FAIRBANKS, AlaskauNew Federal Bldg.. 101 Twelfth Av

Microbanded Manganese Formations:
Protoliths in the Franciscan Complex,
California

By J STEPHEN HUEBNER and MARTA J.K. FLOHR

 

U.S. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1502

Petrographic features, chemistry, and origin of protoliths, which included gel-like
materials, of the manganese-rich carbonate, silicates, and oxides that form lenses
in cherts of the Franciscan Complex

 

 

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON: 1990

US. DEPARTMENT OF THE INTERIOR

MANUEL LUJAN, JR., Secretary

U.S. GEOLOGICAL SURVEY

Dallas L. Peck, Director

Any use of trade, product, or ﬁrm names in this publication is for
descriptive purposes only and does not imply endorsement by the
US. Government

 

Library of Congress Cataloging in Publication Data

Huebner, J .S.

Microbanded manganese formations: protoliths in the Franciscan Complex, California / by J. Stephen Huebner and Marta J .K.
Flohr.
p. cm. — (US. Geological Survey professional paper ; 1502)

Bibliography: p.

Supt. of Docs. no.: I 19.161502

1. Manganese ores—Geology—California—Buckeye Mine Region. I. Flohr, Marta J .K. 11. Title. 111. Title: Protoliths in
the Franciscan Complex. IV. Title: Franciscan Complex. V. Series: Geological Survey professional paper ; 1502.

TN490.M3H78 1990

553.4’629’097941 — dc20 89—600191

CIP

 

For sale by the Books and Open-File Reports Section, US. Geological Survey,
Federal Center, Box 25425, Denver, CO 80225

CONTENTS

 

 

  
 
 
 

 

  

 

 
 
  
  

   
 

   

 

 

 

 

Page Page
Abstract ........................................................................... 1 Mineral characterization and chemistry—Continued
Introduction ...................... 1 Caryopilite and chlorite ................................................ 39
Acknowledgments ........ 2 Taneyamalite and phase A 42
Previous work ................... 4 Santaclaraite ...................... 44
Methodology ..................................................................... 9 Rhodochrosite ... ........... 44
Regional geology ............................................................... 11 Braunite .................... 45
Regional lithologies ...... 12 Hausmannite . .. ........... 45
Provenance ................................................................. 19 Discussion .................................................................. 45
Regional metamorphism ................................................ 19 Vein minerals .............................................................. 46
Protoliths ............................ 21 Origin .............................................................................. 47
Chert ............................ 21 Modern processes of marine manganese deposition ............ 47
Rhodochrosite protolith 25 Modern hydrothermal analogues ............................ 53
Caryopilite protolith ..................................................... 25 Some possible ancient analogues ..................................... 56
Taneyamalite .............................................................. 30 Projected metamorphic path .................................... 56
Chlorite protolith ......... 30 Origins of ancient deposits ............. 58
Gageite ...................................................................... 32 Origin of Buckeye deposit ............................... 61
Hausmannite protolith .................................................. 32 Environment near Buckeye deposit ............ 61
Braunite protolith ............. 33 Manganese deposition at Buckeye deposit.... 62
Diopsidic acmite protolith... 33 Diagenesis and metamorphism ...... 64
Summary of protoliths ................................ 33 Regional framework ....................................... 64
Mineral characterization and chemistry ................................. 34 Possible environments .................................................. 64
Gel-like materials ......................................................... 34 Summary and future work .................................................. 67
Gageite ...................................................................... 37 References ....................................................................... 67
ILLUSTRATIONS
Page
FIGURE 1. Geologic map of the Buckeye mine area and sample location maps ................................................................................ 3
2. Photographs of outcrops and hand specimen of chert ..................... 18
3. Pressure-temperature grid for metagraywacke ......................................................................................................... 21
4. Photographs of slabbed hand specimens of different protoliths ..................................................................................... 22
5. Photomicrographs of characteristic features of protoliths ............. 26
6. Photomicrographs of gel-like and partly recrystallized materials .................................................................................. 28
7. Graphs showing rare-earth element patterns of country rock and ore ........................................................................... 30
8. Photomicrographs of minerals and textures discussed in the text ................................................... 31
9. Ternary diagrams showing summary of chemical analyses plotted in the system Mn-Si—Z Others ..................................... 40
10—14. Diagrams showing:
10. Compositions of caryopilite and chlorites ...................................................................................................... 41
11. Nickel and zinc concentrations in caryopilite and chlorites .............................................................................. 43
12. Compositions of taneyamalite and Phase A ....................................... 44
13. Rare-earth element patterns, exclusive of the Buckeye deposit ....................................................................... 50
14. Predicted greenschist-facies assemblages for Buckeye protoliths projected onto the composition planes MnO—SiOZ—
2 Others and MnO—Si02—02 ............................................................................................................... 57
15. Sketches showing a possible mechanism for the creation of the paleodepositional environment for the Buckeye protoliths ..... 66

III

IV

TABLE 1.
2.
3.
4.

11.

12.

13.

CONTENTS

TABLES

  

Page

Known and hypothetical manganese minerals and common analogues discussed in text ..................................................... 6
Standards used for electron microprobe analysis of samples from the Buckeye deposit, California Coast Ranges ..... 10
Constituents of graywackes from the Buckeye deposit, California Coast Ranges ............................................. 13
Characterization of metasandstones, metacarbonates, and metavolcanic rocks from the Buckeye deposit, California Coast

Ranges, by X—ray powder diffractometry ........................................................................................................... 14
Compositions of minerals from nonsiliceous host rock, in weight percent, from the Buckeye deposit, California Coast

Ranges, determined by electron microprobe analysis ........................................................................................... 15
Chemistry and mineralogy of selected protoliths from the Buckeye deposit, California Coast Ranges .................................. 16
Characterization of siliceous samples from the Buckeye deposit, California Coast Ranges, by X-ray powder diffractometry.... 20
Compositions of gel-like materials, in weight percent, from the Buckeye deposit, California Coast Ranges, determined by

electron microprobe analysis ............................................................................................................................ 35
Compositions of silicates from the Buckeye deposit, California Coast Ranges, determined by electron microprobe analysis 36
Compositions of vein minerals from the Buckeye deposit, California Coast Ranges, determined by electron microprobe

analysis ........................................................................................................................................................ 37
Compositions of rhodochrosite from the Buckeye deposit, California Coast Ranges, determined by electron microprobe

analysis ........................................................................................................................................................ 38
Compositions of the oxides braunite and hausmannite from the Buckeye deposit, California Coast Ranges, determined by

electron microprobe analysis ............................................................................................................................ 39

Comparison of the Buckeye deposit, California, and some modern deep marine Mn-rich deposits ........................................ 49

MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE
FRANCISCAN COMPLEX, CALIFORNIA

By]. STEPHEN HUEBNER and MARTA J.K. FLOHR

ABSTRACT

The Buckeye manganese deposit, 93 km southeast of San Francisco
in the California Coast Ranges, preserves a geologic history that
provides clues to the origin of numerous lenses of manganese carbon-
ate, oxides, and silicates that occur with interbedded radiolarian chert
and metashale of the Franciscan Complex. Compositionally and miner-
alogically laminated Mn—rich protoliths were deformed and dismem-
bered, in a manner that mimics in smaller scale the deformation of the
host complex, and then were incipiently metamorphosed at blueschist-
facies conditions. Eight phases occur as almost monomineralic proto-
liths and mixtures: rhodochrosite, caryopilite, chlorite, gageite, taney—
amalite, braunite, hausmannite, and laminated chert (quartz).
Braunite, gageite, and some chlorite and caryopilite layers were
deposited as gel-like materials; rhodochrosite, most caryopilite, and at
least some hausmannite layers as lutites; and the chert as turbidites of
radiolarian sand. Some gel-like materials are now preserved as trans-
parent, sensibly isotropic relics of materials that fractured or shattered
when deformed, creating curved surfaces. In contrast, the micrites
ﬂowed between the fragments of gel-like materials.

The orebody and most of its constituent minerals have unusually
Mn-rich compositions that are described by the system MnO—SiOz—
Oz—COZ-HZO. High values of Mn/Fe and U/Th, and low concentra-
tions of Co, Cu, and Ni, distinguish the Buckeye deposit from many
high-temperature hydrothermal deposits and hydrogenous or diage-
netic manganese and ferromanganese nodules and pavements. This
chemical signature suggests that ore deposition was related to ﬂuids
from the sediment column and seawater. Tungsten is associated
exclusively with gageite, in concentrations as high as 80 parts per
million. The source of the manganese is unknown; because basalts do
not occur near the deposit, it was probably manganese leached from the
sediment column by reducing solutions. Low concentrations of calcium
(0310 approximately 0.6 weight percent) suggest that the host sedi-
ments formed beneath the carbonate-compensation depth.

The most probable cause of the microbanding is changing proportions
of chemical ﬂuxes supplied to the sediment-seawater interface. The
principal ﬂuxes were biogenic silica from the water column, carbon
dioxide from organic matter in the sediment column, 02 and other
seawater constituents, and Mn+2-bearing ﬂuid. The presence of A1203
and TiO2 (supplied by a detrital ﬂux) in the metashale but not the ore
lens suggests rapid ore deposition. Material supply-rate changes were
probably due to a complex combination of episodic variations in the
hydrothermal flux and periodic flows of radiolarian sand (silica and C02
ﬂuxes) that may be related to climate variations.

Manuscript approved for publication, April 4, 1989.

 

The processes that form recent marine hydrothermal mounds may be
the same as processes that formed the Buckeye deposit. Features
common to both include the presence of Mn-oxyhydroxide crusts
(corresponding to the Buckeye orebody), a large Mn/Fe ratio, low
abundances of most minor elements, and small size. The most impor-
tant differences are the absence of rhodochrosite and manganese
silicates, interlayered with oxide, and the absence of adjacent chert in
the contemporary deposits. These differences may be due to an absence
of the debris of siliceous pelagic organisms, which accumulated in the
Buckeye paleoenvironment. Periodic turbidity ﬂows of chert-forming
radiolarian sand could provide the changes in the ﬂuxes of silica and
organic matter necessary to form manganese carbonate and silicates.
Turbidity ﬂows of graywacke indicate proximity to an environment
with high relief. A possible paleodepositional environment is an oceanic
spreading center approaching a continental margin at which subduction
occurred.

INTRODUCTION

This report documents our initial effort to understand
the origin and metamorphism of banded sedimentary
manganese formations. The term “manganese forma-
tion” (Huebner, 1976) was speciﬁcally chosen to empha-
size the similarities in mineralogy and texture that are
shared by iron formations, such as the Biwabik of
Minnesota and the Hammersley of Australia, and some
manganese deposits of the Paciﬁc coast region of the
United States. In this paper we will describe the petro-
graphic, mineralogical, and chemical characteristics by
which a family of sedimentary protoliths can be recog-
nized and will speculate upon their depositional environ-
ment, thereby creating a deposit model.

Hundreds of supergene manganese-oxide pods in
chert-graywacke terrains of California were prospected
during the First and Second World Wars. At ﬁrst, only
supergene oxides could be used as metallurgical-grade
oxide ore, so mining of a deposit ceased when the deposit
was completely mined or when carbonate and silicates,
rather than oxides, were encountered. When it became
economic in the 1940’s to use rhodochrosite, mining in the
larger deposits exposed manganese carbonate, silicates,
and primary oxides, all precursors of the supergene
oxides. It is these precursors that are the subject of this

paper.

2 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

As a group and as individual deposits, the manganese
orebodies of California are well documented. Trask and
others (1943), Trask (1950), and Trengrove (1960)
described the location, occurrence, and production of
hundreds of deposits. Taliaferro and Hudson (1943),
Crerar and others (1982), and Hein and others (1987)
proposed origins for these deposits. We chose to inves—
tigate the Buckeye mine, the second largest producer of
manganese in California, because its primary (sedimen-
tary or metamorphic) ore is exposed, the deposit is
believed to be representative of many deposits in the
California Coast Ranges, and an extensive collection of
specimens was available as a result of the work of
Huebner (1967). Although most ore had been removed
(the walls of the enclosing chert are exposed), the north
orebody and a nearby prospect-adit were sampled in 1963
(ﬁg. 1). Samples were also obtained from piles of ore
during the summer of 1963 and from surface outcrops
during 1965.

The nomenclature of manganese-rich phases, particu-
larly the sheet silicates, is arcane but need not be so
because many of the Mn phases have compositions anal-
ogous to those of better known Mg or Fe minerals.
Because a major purpose of this report is to present
criteria for recognition of similar deposits, we summarize
in table 1 the Mn minerals that are mentioned below or
that we sought and list the corresponding Mg, Fe, and Al
endmembers. In compiling this list, we used mineralog-
ical summaries (Fleischer, 1981, 1987; Fleischer and
Schafer, 1981, 1983) and their cited references. A
detailed discussion of the crystal chemistry of the phases
that we observed follows the discussion of the lithologies.

The term “gel-like” is used in this paper to describe
materials that resemble the less-than-O.1-p.m ﬁne—clay
fraction from Lake Abert, Oregon (separated from sam-
ple 23de of Jones and Weir, 1983), which we have been
able to examine. This material is pale brown, transpar-
ent, flows if deformed slowly, and fractures if deformed
rapidly or shocked by striking a plastic container that
contains the gel—like material upon a hard surface.
Because these gel-like materials are largely crystalline,
albeit extremely ﬁne grained, they would be expected to
have compositions similar to those of their mineral
constituent and to recrystallize (coarsen) to form monom—
ineralic or polymineralic layers, depending upon their
initial mineralogy.

ACKNOWLEDGMENTS

This study is an outgrowth of a chapter in the ﬁrst
author’s Ph.D. dissertation, written while he was a
graduate student at The Johns Hopkins University.
Partial support was provided by the University and by
NSF Graduate Fellowships. Some laboratory and ﬁeld

 

expenses were defrayed by NSF grants to HP. Eugster;
unfortunately, he is no longer alive to see What became of
the samples. The Buckeye deposits are located on the
Willow Springs Ranch; access was made possible by G.
Purinton, owner in 1963, and by L. Douglas and M. Hall,
owner and foreman, respectively, in 1965. Field assis-
tance was provided by C.E. Beverley (1963) and E.Z.
Huebner (1963, 1965). At the time of these visits, D.F.
Hewett enthusiastically shared his knowledge of, and
stories about, California manganese deposits. J .J . Fitz-
patrick and J .J . Matzko participated in the early stages
of the project’s revival at the US. Geological Survey
(USGS). The continuing dedication of J .J . McGee of the
Reston Microprobe Laboratory made feasible the collec-
tion of the many microprobe analyses by Flohr. B.F.
Jones of the USGS taught us much about seawater-
sediment interactions and made us aware of the less—
than-0.2—um smectite gel reported in Jones and Weir
(1983). R.A. Koski (USGS) loaned us two thin sections,
one of which in shown in ﬁgure SD. We greatly appreci-
ate the enthusiastic cooperation of the following staff of
the USGS chemical analysis laboratories who turned our
“ﬁshing expedition” into a study of the matrix effects in
samples that are both highly manganiferous and distinct
from the marine nodules: P.A. Aruscavage, M.W.
Doughten, J. Fletcher, J .N. Grossman, R.G. Johnson,
B. Libby, and H. Kirschenbaum. We beneﬁted
immensely. Constructive reviews were provided by P. R.
Brett, J. Webster, and E. Force (all USGS) and L.
Raymond (Appalachian State University). Brett sug-
gested that serpentinization at low temperature might
drive convection through abyssal sediments far removed

FIGURE 1.—Geolog'ic map and sample locations. A, Distribution of
chert-metashale, graywacke-sandstone, and manganese deposits in
the Franciscan Complex near the Buckeye and Ladd mines, northern
Diablo Range of the central California Coast Ranges. Adapted from
Cox and others (in Trask, 1950, pl. 4) and Raymond (1973b, ﬁg. 2).
Inset adapted from Bailey and others (1964, pl. 1). Note the
association of manganese deposits with chert and the undulatory
contacts between chert and graywacke. The Carnegie fault and
northern part of the Pegleg fault are parts of the Tesla-Ortigalita
fault system. B, Geologic sketch map of the vicinity of the Buckeye
mine. The base is aerial photography. Subdivision into broken
formations follows usage of Raymond (1973a): Kstg, Sulphur Gulch
broken formation; KJfg, Grummett broken formation; KJfo, Oso
broken formation. C, Location of samples collected from surface
outcrops and dumps. Base is identical to that used for B. Samples
designated Hxac correspond to those labeled 65Hmc in text. Sample
B112 is a grab sample from the gulch below (south of) the north
orebody and is not plotted. D, Location of samples collected from the
open cut, 1837- and 1805—ft levels of the north Buckeye orebody.
Mine workings and trend of ore lens adapted from maps and sections
by Crittenden and others (pl. 18 of Trask, 1950). E, Location of
samples collected from exploration adit between the north and south
Buckeye orebodies.

INTRODUCTION

xocxéwEEEO
:mEEEO
85 :52).
250:.—

aoaﬁ

5:6 Siam
v>mxuzm £30m
mimic—am 552
nmEEan
>eco;m5_
vumg

cﬂnmm

O'JOFN
I—u—v—

v-‘NMV‘W‘DNCD

\\ 360qu ommcmmcmE

aw. o—IW

   

A2882, tcm 908392 66800 $86ch

EmeE, ncm 968892 85:53 >w__m> 690 E

ZOF<Z<J_n_Xm

£35205. m N F o

 

mugs. N _ a

XmJnZZOU Z<Um_oz<mm
/

\\ \\> \s\ .I/

\ \X\\\\\\ 5/ I’ll
\ \xkkw? M»
\\W\ \\ \\w\\\\\\\\\\\\\\aw ~

\\\~

VIA" 1‘3

mo
nu
v
n

wEEEBiow om ov om ON 3 o

JEiiLiI.

832% 9. on cm H: o

 

 

wmz_s_
w>mxuam

mum 8&9:ch cum. I

   
 

0.550 52‘

 

 

4 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

 

 

 

 

 

 

0 500 1000 1500 2000 FEET

| l | | J

l l | l I l l

O 100 200 300 400 500 600 METERS
EXPLANATION

a m as Contact between broken formations
(see text), following Raymond
(1973a); dashed where
repositioned

 

(see tables 4, 6)

Graywacke

V////A Metamorphic rocks

 

Basal conglomerate Contact between lithologies within

Grummett broken formation
Ribbon chert in KJfg
— — — — Trace of prominent white chert

E

Strike and dip of OUtCTOD (reef)
bedding
A Aragonite-bearing vein
+ Anticline
10 Diamond drill hole

8

from a spreading center. Although we did not adopt this
suggestion for the Buckeye, serpentinization may be an
important factor in the origin of laterally extensive
diagenetic deposits. Raymond challenged us to expand
our discussions of the relationship between the environ-
ment of deposition and accretionary complexes.

 

 
 
 
   
   
  
 

N
H40
H41

H12,13

\_ FCm

H28
-—\02

-.\N BuckeyeJ/n

IE §3§1€1€| lynx 8109-110,H78

H11

_H19-212

H123 12 _____
q— S _Buckeye Exploration adit _H53 54 H
LB4_5J\:< \< -62'65
11 H43-49,51
_26»128130-\_ H60,61
\— H23-25

   

_TEG:

 

 

 

 

0 500 1000 1500 2000 FEET
I I I I I
l l l l l l l
0 100 200 800 400 500 600 METERS
EXPLANATION
Sample from outcrop 10 Diamond drill hole
,_ _ _ _l —< Adit
LEEJ Loose sample of ore /\/ Road

PREVIOUS WORK

Previous investigators dealt with aspects of the min-
eralogy, petrology, chemistry, and genesis of the Buck-
eye (and similar) deposits, but none has considered all of
the different kinds of data that must be integrated to
understand deposits such as these. In the following
paragraphs we review the observations and data of our
predecessors and show that there is still ambiguity about
the fundamental nature and origin of the ore.

The most complete description of the California Coast
Ranges deposits is by Taliaferro and Hudson (1943) Who
made frequent reference to the Buckeye mine When
discussing the origin of similar deposits. Unique obser-
vations were made while ore was being removed. Most

PREVIOUS WORK 5

 

 

  
   

Open cut
~ 1970 ft

[FE—lgél

0 30 60 9012OFEET

 

 

35
B72 0 1O 20 SOMETERS
1837‘
adit
EXPLANATlON

Red chert-argillite \\_
Green chert-argillite
E White chert

Sample from outcrop

LEESJ Loose sample of ore

Trend of ore lens;
elevation in feet

_r_ Strike and dip of bedding

<—*~ Syncline showing plunge

of axis

 

 

 

30 FEET

|
| | l l |

10 METERS

O——O

EXPLANANTION

E White chert, argillite
Red chert, argillite
Sample number

E

—L Strike and dip of bedding

<—)~(— Syncline showing plunge
of axis

FIGURE 1. —Continued.

manganese deposits are associated with thick lenses of
light-colored massive chert, which in turn are enclosed in
darker, rhythmically bedded radiolarian cherts. The
converse is not true: the vast majority of massive cherts
are not associated with manganese deposits. Many cherts
are associated with volcanic ﬂows, but this is not the case
near the Buckeye mines. Taliaferro and Hudson thought
that the cherts were originally oozes rich in colloidal
silica. On becoming a gel the ooze expelled the impuri-
ties, thereby concentrating them in what are now the
argillaceous interbeds. Thus the rhythmic chert layering
was thought to be the product of diagenetic segregation.
Presumably, they thought that the precursor to the
massive cherts was an ooze so pure that there were
insufﬁcient impurities to form argillaceous layers.
Orebodies known to Taliaferro and Hudson formed
lenses less than 11 m thick and less than 240 m in
maximum dimension. Folding and faulting obscured the
shape of the orebodies in the plane parallel to the
layering, but internally they were banded, being com-

 

posed of smaller lenslike bodies that lay parallel to the
layering in the surrounding cherts. These lenses con-
sisted principally of manganese carbonate and silicates.
Unfortunately, neither the spatial distribution of these
two components within the lens nor the nature of their
contact with the surrounding chert was discussed. An
important observation, based in part on analysis of thin
sections but overlooked by some subsequent investiga—
tors, is that ﬁne banding was present in material that
Taliaferro and Hudson called manganiferous opal (1943,
p. 240). The lenses are now partly recrystallized to pink
rhombohedra of rhodochrosite and to yellow, orange, and
brown ﬁbers of manganese silicates called bementite.
Hausmannite and braunite were uncommon and, because
of association with barite and copper, were regarded as
having formed during subsequent hydrothermal action.
(Fleischer and Richmond (1943) reported the presence of
braunite, but not hausmannite, at the Buckeye mine.)
Taliaferro and Hudson thought that the deposits were
“chemical sediments formed at the same time and in the

 

6 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

TABLE 1. —Known and hypothetical manganese minerals and common analogues discussed in text
[—, not applicable; MVI/TIV, ratio of octahedrally coordinated to tetrahedrally coordinated cations]

 

 

Mn species Ideal formula MVI/TIV Fe analogue (Mg,Al,Ca) analogue
Rhodochrosite ................. MnCO3 — Siderite Calcite.
Hausmannite .................. Mn304 — Magnetite Unknown.
Bixbyite ...................... Mn203 — Hematite Unknown.
Pyrolusite ..................... Mn02 — Unknown Unknown.
Todorokite .................... Mn+2Mn3+4O¢ H20 — Unknown Unknown.
Nsutite ....................... Mn;2Mnf3x02‘2x(OH)2X — Unknown Unknown.
Braunite ...................... Mn+2Mng38i012 — Unknown Neltnerite.
Birnessite ..................... Na4Mn14027~9H20 — Unknown Unknown.
Vernadite ..................... (Mn+4,Fe +3, Ca,Mn)(O,OH)2~ nHzO — Unknown Unknown.
Pyrophanite ................... MnTiO3 — Ilmenite Geikielite.
Pennantite .................... Mn10A12(Si,Al)8020(OH)16 1.5 Chamosite Clinochlore.
Gonyerite ..................... Mn1(,Fe;3(Si,Fe+3)8020(0H)16 1.5 Unknown Unknown.
Bannisterite ................... (K,Ca)Mn10(Si,Al)16038(0H)8(H20)2_6 .625 Unknown Unknown.
Ganophyllite ................... (K, Na,Ca)6Mn24(Si,A1)40(O,0H)112(H20)21 .6 Unknown Unknown.
Caryopilite .................... MnGSi4Ow(OH)8 1.5 Greenalite Unknown.
Bementite ..................... Mn7Si6015(OH)8 1.17 Unknown Unknown.
Unknown ..................... (1/2Ca,Na)0‘66Mn6+2(Si,A1)8020(OH)4(H20)8 .75 Unknown Saponite.
Unknown ..................... (Vzca, Na)0‘66Mn4+3(Si,A1)8020(OH)4(H20)8 .50 N ontronite Unknown.
Unknown ..................... MnSSi8020(OH)4 .75 Unknown Talc.
Unknown ..................... Mn308i40096(OH)28 .75 Minnesotaite Unknown.
Ashley (1986) .................. K2Mn6(Mn+‘n5,Al)ZSi6020(OH)4 .75 Biotite Phlogopite.
Unknown ..................... K2Mnjf3(Mn+3,Al)2Si6022(OH)4 .50 Unknown Muscovite.
Akatoreite .................... Mn5,Alei8023(OH)9 .90 Unknown Unknown.
Piemontite .................... Ca2Mn;3AlSi3012(OH) 1.0 Epidote Clinozoisite.
Gageite ....................... Mn581209_x(0H)2x 2.5 Unknown Balangeroite.
Taneyamalite .................. NatMnIZSimOgomH)14 1.0 Howieite Unknown.
Unknown ..................... Mn6(Mn+3,Al)38i6020(OH)5 1.5 Deerite Unknown.
Unknown ..................... K0_1Mn13(Al,Si)18042(OH)14 .72 Zussmanite Unknown.
Parsettensite .................. K(Mn,Al)7Si8020(OH)8-2H20 .67 Unknown Unknown.
Unknown ..................... KsMn48(Si63A19)O 168' (OH) 48 .67 Ferrostilpnomelane Lennilenapeite.
Unknown ..................... K5Mn183(Si63A19)0216- 36H20 .67 Ferristilpnomelane Unknown.
Spessartine ................... Mn3A12(SiO4)3 1.67 Almandine, andradite Pyrope, etc.
Calderite ...................... Mn3+2l<‘e;3(SiO4)3 1.67 Unknown Unknown.
Unknown ..................... Mn3(Si04)(F,OH)2 3.0 Unknown Norbergite.
Alleghanyite .................. Mn5(SiO4)2(OH)2 2.5 Unknown Chondrodite.
Manganhumite ................. Mn7(SiO4)3(OH)2 2.33 Unknown Humite.
Leucophoenicite ............... Mn7(SiO4)3(OH)2 2.33 Unknown Unknown.
Sonolite ....................... Mn9(SiO4)4(OH,F)2 2.25 Unknown Clinohumite.
Tephroite ..................... MnZSiO4 2.0 Fayalite Forsterite.
Donpeacorite .................. MnMgSi206 1.0 Ferrosilite Enstatite.
Kanoite ....................... MnMgSiZO6 1.0 Clinoferrosilite Pigeonite.

J ohannsenite .................. CaMnSi206 1.0 Hedenberg'ite Diopside.
Ashley (1986) .................. NaMn’r‘D’SiZO6 1.0 Acmite J adeite.
Rhodonite ..................... CaMn4(SiO3)5 1.0 Ferrobustamite Unknown.
Pyroxmang‘ite ................. Mn7(SiO3)7 1.0 Pyroxferroite Unknown.
Inesite ........................ CaQMn7Si10028(OH)2(H20)5 .90 Unknown Unknown.
Santaclaraite .................. CaMn4Si5014(OH)(OH)- H20 1.0 Unknown Unknown.
Nambulite .................... (Li,Na)Mn4Si5014(OH) .80 Unknown Unknown.
Manganbabingtonite ............ CazMnFe +3Si5014(OH) .80 Babingtonite Unknown.
Marsturite .................... N212021.2Mn68i10028(0H)2 .80 Unknown Unknown.
Tirodite ....................... anMg5Si8022(OH)2 .88 Grunerite Cummingtonite.
Dannemorite .................. anFe58i8022(OH)2 .88 Grunerite Cummingtonite.
Kozulite ...................... NaNazMn4(Fe +3,Al)Si8022(OH)2 .88 Arfvedsonite Magnesioarfvedsonite.
Tsilaisite ...................... Na(Mn1.5A11.5)A16(BO3)3Si6018(O,0H,F)4 1.0 Schorl Elbaite.
Saneroite ..................... Na.2Mn10(Si,V)12034(OH)4 .83 Unknown Unknown.
Welinite ...................... Mn6(W,Mg)ZSi2(O,OH)14 4.0 Unknown Unknown.

PREVIOUS WORK 7

TABLE 1. —Known and hypothetical manganese minerals and common analogues discussed in text—Continued
[—, not applicable; MVl/Tw, ratio of octahedrally coordinated to tetrahedrally coordinated cations]

 

 

Mn species Ideal formula MVI/TIV Fe analogue (Mg,Al,Ca) analogue
Franciscanite .................. Mn6[V,[:] ]ZSi2(O,OH)14 4.0 Unknown Unknown.
Orebroite ..................... Mn6(Sb"5,Fe+‘°’)ZSi2(O,OH)14 4.0 Unknown Unknown.
Pumpellyite-Mn2+ .............. CazMn+2A12(Si04)(Si207)(OH)2- H20 1.67 Ferropumpellyite Unknown.
Okhotskite .................... CazMn+2(Mn+3,Al)2(Si04)(Si207)(OH)2‘ H20 1.67 J ulgoldite Pumpellyite.
Sursassite ..................... Mn2Ala(SiO4)(Si207)(OH)3 1.67 Unknown Unknown.
Mcfallite ...................... CazMn;3(Si04)(Si207)(0H)3 1.67 Unknown Unknown.
Manganaxinite ................. CazMn+2A12BSi4015(OH) 1.0 Ferroaxinite Magnesioaxinite.
Kanonaite ..................... Mn+3AlSiO5 1.0 Unknown Andalusite.
Carpholite .................... MnAIZSi206(OH)4 1.5 Ferrocarpholite Magnesiocarpholite.
Serandite ..................... NaMnZSi308(OH) .67 Unknown Pectolite.

J ohinnesite .................... NazMg4Mn12As;5Si12043(OH)6 1.33 Unknown Unknown.
Manganpyrosmalite ............ anSi12030(OH,Cl)20 1.33 Pyrosmalite Unknown.
Schallerite .................... Mnf628i12A5§ 3036(OH)17 1.33 Unknown Unknown.
Mcgillite ...................... Mn88i6015(OH)8012 1.33 Unknown Unknown.
Friedelite ..................... MnSSi6015(OH,CI) 10 1.33 Unknown Unknown.

 

same marine environment as the cherts” (p. 272). In their
interpretation, the cherts were deposited as siliceous
oozes and the manganese—rich materials as ﬁne-grained
gray manganese carbonate and thinly banded, resinous,
almost optically isotropic manganiferous opal containing
14 weight percent H20. Silica and manganese were
thought to be associated with volcanism, and the man-
ganese was supplied as carbonate and oxides. The mas-
sive cherts and the manganese orebodies represent peri—
ods of rapid deposition. After examining many deposits,
these authors found no evidence for replacement of chert
by ore (Taliaferro and Hudson, 1943, p. 273).

Trask and Pierce (in Trask, 1950, p. 211—221) focused
on the Ladd-Buckeye area. Their observations were
based in part on ﬁeldwork during 1940—42, when the
deposit was actively mined. Ore was always associated
with white, usually massive, chert. They reported two
kinds of ore: (1) massive, ﬁne-grained gray rhodochrosite
with bementite layers and minor hausmannite, rho-
donite, and braunite and (2) a mixture of rhodochrosite
and bementite, disseminated in chert. This second type
forms a thin transition zone between massive carbonate
and chert at the sides and ends of the lenses. Trask and
Pierce thought that the pinching and swelling of the
chert bodies reﬂected deposition in marine basins. Rec—
ognition that the red color of chert represents an oxidized
state of iron and that rhodochrosite represents a reduced
state of manganese led them to propose that the manga-
nese ores and white chert were formed in reducing basins
and the oxidized chert formed under shallower marine
conditions. The generally sharp contacts with the sur-
rounding chert, at both the sides and the extremities of
the manganiferous lenses, were thought to be caused by
presumed distinct mechanical consistencies of the man-
ganiferous sediments and silica gel. Hydrothermal activ-

 

ity was thought to be indicated by the presence of quartz
veins and Cinnabar, by the lack of bedding and non—
uniform crystallinity of the gray rhodochrosite, and by
the presence of hausmannite, rhodonite, and braunite.
But despite this hydrothermal activity, the manganese
did not migrate into the wall rocks.

Subsequent observers also found massive ore. Trask
(1950, p. 287—288) reported that, in 1942, massive ore
was exposed in the north and south Buckeye orebodies.
Crittenden (in Trask, 1950, p. 288—289) examined the
orebody in 1944, at which time the primary ore was well
exposed at levels subsequently accessible to Huebner
(1967) and Hein and others (1987). Crittenden reported
that along strike, the north Buckeye orebody was
defined by change in manganese content rather than
thickness, gradually changing to chalcedonic quartz and
thence to chert. This is the only report of a gradational
contact from the orebody into the chert. Crittenden
described the ore as “an exceedingly massive and inti-
mate mixture” (p. 288) of (in decreasing abundance) gray
rhodochrosite, massive braunite and possibly hausman-
nite, brown to tan manganese silicates, pink rhodochros-
ite, neotocite, rhodonite, inesite, and sulﬁdes. The older
minerals (p. 289) were gray rhodochrosite, braunite,
hausmannite (if present), and manganiferous chert (pre-
sumably neotocite). Of these minerals, gray carbonate
was regarded as “original” (p. 288) and the braunite and
hausmannite as possibly metamorphic. Younger miner—
als regarded as clearly hydrothermal in origin (p. 289)
were pink rhodochrosite, quartz, rhodonite, inesite, sul-
ﬁdes including molybdenite, brown and tan manganese
silicates, and neotocite. Again, the spatial distribution of
these minerals within the ore lens was not given.

The property was drilled by the Bureau of Mines in
1940—41. Volin and Matson (1949) reported that the ore

8 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

consisted of rhodochrosite, bementite, rhodonite, psilom-
elane, and pyrolusite. The last two minerals are undoubt-
edly supergene. The primary ore was a mixture of
carbonates and silicates, but layering was not noted. The
boundary between the primary ore and chert was dis-
tinct. Four drill holes encountered only low concentra—
tions (1-11 percent by weight) of Mn and demonstrated
that the high grade of the mined orebodies did not
continue downward or along strike. Subsequently Tren-
grove (1960) summarized this work as indicating that “a
large mineralized zone of low-grade manganiferous
chert” surrounded the north orebody.

Huebner (1967) described the mineral associations,
using X-ray powder diffractometry and thin sections to
identify the minerals. He determined that the major
ore-forming minerals were hausmannite, braunite,
rhodochrosite with lattice constants nearly identical to
those of pure MnC03, and What had been called bement-
ite. He found that bementite from the Buckeye and
elsewhere was a mixture of phases probably related to
septechlorite (7A, serpentine), chlorite (14A), and the
pyrosmalite group of minerals. He also found 9A and
6.9A reﬂections, now recognized as due to the presence
of taneyamalite (Matsubara, 1981) and gageite (Phillips,
1911), and in thin sections observed pseudomorphs
“more slender than most rhodonite crystals” (p. 66) of a
mineral now known as santaclaraite (Erd and Ohashi,
1984). Unlike other workers, he did not ﬁnd rhodonite,
inesite, or neotocite. Huebner did not explicitly address
the massiveness of the ores, but it is clear from his
descriptions and photographs that he encountered both
layered and massive ore. He recognized that the deposit
had an unusual bulk composition that could be closely
modelled by the system MnO—SiOZ—OZ—COZ—HZO. He
suggested that the Buckeye deposits and enclosing
cherts formed as chemical sediments in local depressions,
rapidly precipitated from thermal waters in the intervals
between turbidite (graywacke) units. Rapid deposition of
the manganiferous material and adjacent silica diluted
the supply of clastics, represented by the argillaceous
layers within the chert rhythmites, providing an expla-
nation for the high purity of the deposit and the absence
of shale interbeds in the massive chert adjacent to the
ore. In the absence of associated volcanic rocks and in
view of the rapid deposition, he did not favor the leaching
of extrusive ﬂows as the source of the manganese.
Rather, because of the high Mn/Fe ratio, he preferred a
hydrothermal source and thought that the carbonate and
silicates resulted from mixing of hydrothermal and oce-
anic waters. Most important, he noted that the Mn-rich
sediments at the Buckeye were relatively oxidized and
proposed that their high intrinsic oxygen fugacity values
were inherited from initially oxidized sediments. This

 

concept was later developed more clearly (Huebner,
1969, 1976).

Hewett (1972) reviewed the occurrences and signiﬁ—
cance of hausmannite and braunite, concluding that
hausmannite formed from hypogene (rising) solutions,
whereas braunite could be either hypogene or super—
gene. In discussing the manganese deposits of the Cali-
fornia Coast Ranges, Hewett stated that (p. 87) “layers
or lenses of nearly pure manganese carbonate and thin
layers of hydrous manganese silicate (neotocite or ‘man—
ganese opal’) were laid down either in or near persistent
layers of chert. Under the inﬂuence of later hydrother-
mal solutions at the Buckeye deposits, the carbonate was
thought to alter to hausmannite and the hausmannite to
bementite and neotocite; braunite has also been identi-
ﬁed in this assemblage.” Clearly, Hewett saw the lay-
ered nature of the ore and thought that carbonate was
primary, that opaline material was both primary and
formed by alteration of carbonate, and that all hausman-
nite and braunite were formed by alteration of earlier
formed minerals.

Raymond (1974) discussed the lithologies of the Mt.
Oso area of the Diablo Range, which includes the Buck-
eye mine, and advocated a depositional environment that
combined an abyssal ocean ﬂoor or trench with an ocean
ridge. The manganese deposits were not discussed, but
subsequently Raymond (1977) placed the Ladd-Buckeye
district in the context of a seaﬂoor spreading center.

Crerar and others (1982) initiated a detailed geochem-
ical study of manganese deposits of the Franciscan
Complex. They considered the 800 known deposits to be
similar and focused on two deposits in northern Califor-
nia, the Blue Jay and South Thomas mines. Crerar and
others proposed an origin involving hydrothermal pro-
cesses analogous to those forming the mounds near the
Galapagos rift (Corliss and others, 1978). Among the
principal lines of evidence cited are the linear distribu-
tion of orebodies and the fact that massive ore is devoid
of detritus, is depleted in Fe and some trace elements
relative to marine nodules, and has rare-earth-element
(REE) patterns similar to patterns thought to be of
hydrothermal origin. High heat ﬂow associated with a
spreading center was assumed to provide the energy
needed to drive the hydrothermal system. Crerar and
others favored deposition at either a midocean ridge or
back-arc basin.

Stable-isotope data for rhodochrosite from the Buck-
eye mine were provided by Hodgson (1966) and Yeh and
others (1985). Highly negative 813CPDB values of —45 to
—54 per mil (referred to Peedee belemnite) were inter-
preted as indicating that the CO2 was derived by the
oxidation of methane (Hathaway and Degens, 1969).
Because this conclusion does not depend upon partition-
ing of carbon isotopes between ﬂuid and unusual phases

METHODOLOGY 9

with unknown behavior, we accept an origin for the 003’
that involves decomposition of organic matter. From a
single SEOSMOW value (+19.5 per mil, referred to stan-
dard mean ocean water) for a rhodochrosite sample
assumed to have been deposited from seawater, Yeh and
others (1985) assigned a temperature of formation of 75
°C. From 81805M0W values of 18.7 to 26.6 for eight
rhodochrosite samples from deposits within the Fran-
ciscan Complex, Hein and Koski (1987) obtained temper-
atures of formation of 12 to 120 °C. The signiﬁcance of
these temperatures is in doubt because the fractionation
factor for rhodochrosite precipitation is uncertain (see
discussion by Okita and others, 1988, p. 2683) and
because, as we will show, all rhodochrosite from the
Buckeye deposit observed by us is admixed with other
phases of unknown oxygen isotope fraction behavior and
all samples of Buckeye rhodochrosite observed by us are
at least partially recrystallized.

Huebner and others (1986a) brieﬂy described mineral-
ogical banding and the presence of relic gels (the gel-like
materials in this report) in the Buckeye deposit. Carbon—
ate mud, interlayered with silicate and oxide gels, was
deposited at the sediment-seawater interface. The layer-
ing was thought to have been caused by relatively rapid
variations in the supply ﬂuxes and may have been
climatically induced.

Hein and others (1987) provided chemical, mineralog-
ical, and isotopic data for the Buckeye and adjacent Ladd
deposits. They cited Huebner (1967) for observing abun-
dant hausmannite and Erd and Ohashi (1984) for “rare”
santaclaraite, apparently because they observed neither
mineral in their own samples. Because we observed
these minerals in many samples collected at the Buckeye,
we believe that sampling by Hein and coworkers may not
have been as comprehensive as that of Huebner (1967).
They proposed that the Buckeye and Ladd deposits
formed in a deep basin at a continental margin by
replacement of thin turbidites of radiolarian sand by
carbonate. The massive chert, which consists of lami-
nated chert without argillaceous interbeds, was thought
to have formed from silica released by this replacement
process, but the composite nature of the massive chert
was not explained. Because opal-CT was reported to
occur in the manganese ores but not in the surrounding
rocks, this replacement process was thought to occur at
less than 80 °C and at a depth of 100—700 m in the
sediment pile. Hein and others believed that the manga-
nese was originally derived from manganese oxides and
oxyhydroxides dispersed in adjacent chert-metashale
layers, but by comparison with the compositions of
average shales, argillites, and cherts (Pettijohn, 1957),
we do not ﬁnd that the chert-argillites of the Ladd-
Buckeye district are depleted in Mn.

 

Hein and Koski (1987) summarized available stable-
isotope data for the manganese deposits of the Fran-
ciscan Complex, including the Ladd and Buckeye depos-
its. The dominant ore mineral was reported to be
rhodochrosite; some samples contained braunite and
bementite; and kutnahorite, manganiferous calcite, pen-
nantite, serandite, and sursassite were reported to be
less abundant. No mention was made of compositional
banding. They reported the “existence of replacement
textures of rhodochrosite and silica polymorphs” (p. 725)
but gave no examples. In their preferred scenario,
manganese was originally deposited within the rhythmic
chert sequences as hydrogenous oxides and oxyhydrox-
ides. During diagenesis, buried organic matter produced
CH4 that rose into the overlying sediments and was
oxidized, in situ, by the oxides and oxyhydroxides,
forming COZ. COZ, CH4, and reduced manganese moved
through ﬂuids produced by compaction and silica dehy-
dration to form rhodochrosite at or below the zone of
sulfate reduction and sulﬁde precipitation, below the
seawater-sediment interface. In their model, it is not
clear why rhodochrosite did not form at the same time
and place at which the oxides and oxyhydroxides within
the sediment column were reduced. A replacement ori-
gin was implied, but the internal structure and composi-
tional diversity of the orebody were not explained.

Most early investigators regarded the gray manganese
carbonate and manganiferous opal as chemically precip-
itated sediments, but in recent work Hein and others
(1987) and Hein and Koski (1987) concluded that the
carbonate is the product of a replacement process. Most
investigators also appeal to later hydrothermal processes
to form the diverse mineralogy, but Huebner (1967;
1986a,b) interpreted this diversity to be a feature of the
sedimentary environment. Although there are some
reports of banded or layered ore at the Ladd and
Buckeye mines, most reports are of massive ore at the
working face of the Buckeye mine. The recent papers by
Hein and his coworkers make no reference to layered
ore. If similar deposits or their metamorphosed equiva-
lents are to be recognized elsewhere, it is essential that
the internal structure, compositional diversity, and min-
eralogy be known.

METHODOLOGY

Sample locations are shown in ﬁgure 1. Data were
obtained by using transmitted and reﬂected light optical
microscopy, X—ray powder diffractometry (XPD),
electron-microprobe analysis (EMP), and chemical bulk
analyses. Because the manganese-rich materials are het-
erogeneous, it was necessary to examine a large number
of samples. We examined 77 polished thin sections, 37

10 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

covered thin sections, and 24 polished blocks. We
crushed 87 samples for bulk chemical or X-ray diffraction
analyses. Each crushed sample represents a single litho-
logic type, and most were immediately adjacent to
material used for a thin section. Because much of the ore
consists of thin layers, many of these crushed samples
were less than 2 g in weight. Approximately 33 addi-
tional small samples were systematically examined by
XPD. Thus, we were able to integrate diverse kinds of
measurements made on essentially identical material.
The seven samples chemically analyzed for this report
received special handling. Possible Fe, Co, Cu, Mn, Ni,
or Zn contamination by the saw blades was reduced by
grinding sawed surfaces with silicon-carbide grit fol-
lowed by cleaning ultrasonically in deionized water and
acetone. These samples were crushed in a hardened steel
mortar (Fe, Cr) and ground with an agate mortar and
pestle. To reduce sample-to-sample cross contamination,
the ﬁrst crushed material for each sample was set aside
and used for XPD rather than trace-element chemical
analyses. The bulk chemical analytical methods are sum-
marized in Baedecker (1987). Total contamination (by the
process of sawing, crushing, and grinding) and cross
contamination (by the previous sample) were investi—
gated by subjecting spectroscopically pure silica glass to
identical procedures (both before and after the aforemen—
tioned seven samples). The second—crushed silica glasses
contained less than 4 ppm Cr, less than 50 ppm Fe, and
less than 1,600 ppm Mn, concentrations which are much
lower than those in the protoliths, indicating that con-
tamination was not a signiﬁcant factor. We are conﬁdent
that observed large differences in trace-element concen—
trations of our Mn-rich samples are real.

The X-ray diffractometer used CuKa radiation, a
proportional counter, and pulse-height analysis that
effectively discriminated against MnKa fluorescence
from Mn-rich samples. Even so, the sheet silicates gave
such poor-quality patterns that useful unit-cell dimen-
sions could not be obtained. (Attempts to obtain good
patterns with a Hague-Guinier camera equipped with a
monochromator were even less successful, perhaps
because, in the Guinier technique, X-rays pass through
the powdered sample rather than reﬂect from its sur-
face.) The goniometer divergence, receiving, and anti-
scatter slits were 1°, 0.1 mm, and 1°, respectively; the
radius was 170 mm; and the takeoff angle was 6°. At a 20
scanning rate of 1° per minute, the patterns routinely
yield a. crystallinity index (Murata and Norman, 1976) of
10.5 to 11.5 for powder obtained by crushing and grind-
ing (to pass ZOO-mesh screen) a large quartz crystal of
quality suitable for electronic oscillators.

Electron microprobe analyses were performed on a
unit with 8 wavelength-dispersive channels operated at
15 kV and 100 nA beam current (15 nA nominal specimen

 

TABLE 2,—Standards used for electron microprobe analysis of sam-

ples from the Buckeye deposit, California Coast Ranges
[See Huebner and Woodruff (1985) for fuller explanation; na, not analyzed]

 

Phase analyzed

 

 

 

 

Element
Oxides Carbonates Silicates

Na ............... FSLC na FSLC

Mg ............... OXTB CDOS PXEN

A1 ................ OXGH na FSLC

Si ................ FSLC na FSLC

K ................ na na FSBO, MSFPl

Ca ............... FSLC CDOS FSLC

Ti ................ OXIL na SPHC

V ................ OXVA OXVA OXVA

Cr ................ OXTB na OXTB, OXBU

Mn ............... OXPA CRAP OLST

Fe ............... OXIL CSIG OLSF, MBLM

Co ............... OLCO na na

Ni ................ OLNI OXNC OLNI

Zn ............... OXGH OXGH OXGH

Sr ................ na CSTR na

Ba ............... GLBA GLBA GLBA1

F ................ na na MSFPl

As ............... OXAS,2 ARDN2 na na
Explanation of acronyms

ARDN ............ Ardennite, Pennsylvania State University
Standard 5—114.

CDOS ............ Dolomite, Oberdorf, Austria.

CRAP ............ Rhodochrosite, Alma Park, N. Mex.

CSIG ............. Siderite, Ivigtut, Greenland.

CSTR ............ Strontianite, Oberdorf, Austria.

FSBO ............ Orthoclase, St. Lawrence Co., NY.

FSLC ............ Labradorite feldspar, Lake County, Oreg.

GLBA ............ Synthetic barium—bearing glass.

MBLM ........... Biotite, Lemhi, Idaho.

MSFP ............ Synthetic ﬂuor-phlogopite, KMgsAlSi3010F2.

OLCO ............ Synthetic 0028104.

OLNI ............ Synthetic NizSiO4.

OLSF ............ Synthetic fayalite, FeZSiO4.

OLST ............ Synthetic tephroite, MnZSiO4.

OXAS ............ Synthetic As203, National Bureau of Standards
Standard Reference Material 833'.

OXBU ............ Chromite, Bushveld Complex, S. Africa.

OXGH ............ Gahnite, Brazil.

OXIL ............. Ilmenite, Ilmen, USSR.

OXNC ............ Synthetic bunsenite, NiO.

OXPA ............ Synthetic bixbyite, Mn203.

OXTB ............ Chromite, Tiebaghi, New Caledonia.

OXVA ............ Synthetic V203.

PXEN ............ Synthetic enstatite, MgSi03.

SPHC ............ Titanite, Hemet quadrangle, California

 

1 Used regularly only for parsettensitelike phase and to check other silicates.
2 Used to analyze phase D only.

current) and using a focused beam (estimated activation
volume less than 10 um3). The standard or unknown was
exposed to the electron beam for 20 to 80 seconds.
Analytical software was described by McGee (1985).
Mineral standards, documented by Huebner and Wood—
ruff (1986), are summarized in table 2. Backgrounds were
obtained by interpolation, based on the mean atomic

REGIONAL GEOLOGY 11

number (Z) of the unknown, between the count rates
obtained on quartz or MgO (low Z) and on tephroite,
NiO, or V203 (high Z). Peak intensity data were cor-
rected by the method of Bence and Albee (1968). Alpha
factors used for constituent oxides were measured by
Albee and Ray (1970) and calculated from ﬁrst principles.
Carbon was not determined, but the carbonate analyses
were corrected in a similar manner, providing that CO2
had been excluded from the analysis of the reference
standard (and the calculation of its set of B—factors).
Analyzed compositions of all phases were compared with
an energy dispersive analysis spectrum to determine
whether any elements present in signiﬁcant quantities
had been missed in the wavelength dispersive analysis.

Fluid inclusion techniques were not used in this study
because ﬂuid inclusions are rare in the Buckeye ore. The
only inclusions found were in a minute amount of barite
that appeared to replace rhodochrosite in sample B29.
These inclusions were minute and had small vapor
bubbles.

REGIONAL GEOLOGY

In lithologic variation and deformation, the thinly
banded ore mimics the host Franciscan Complex but in
miniature. Thus, in addition to providing a geologic
framework for manganese deposition, the regional geol-
ogy may provide clues to the origin of some delicate
textures of the ores. The area containing the Ladd and
Buckeye mines has been mapped by Huey (1948), Cox
and others (in Trask and others, 1950, pl. 4), Maddock
(1955), Raymond (1973a,b), and Raymond and Maddock
(unpub. mapping, 1976—78). Detailed geologic descrip-
tions of the immediate vicinity of the Buckeye mines
were published by Pierce (in Trask, 1950, pl. 17) and by
Volin and Matson (1949). Huebner (1967) recognized that
the metagraywacke sequences at the Buckeye mine were
turbidites that retained clastic texture after low-
temperature and high-pressure metamorphism. Ernst
(1971) correlated the textural grade of the metagray-
wackes with the development of metamorphic assem-
blages containing pumpellyite, albite, lawsonite, and
jadeitic pyroxene but did not ﬁnd jadeitic pyroxene near
the Buckeye. We will demonstrate below that the meta-
graywackes correspond to his textural grades 1 and 2 and
mineral zone 2. Raymond (1973a,b) provided metamor-
phic and tectonostratigraphic maps. He found lawsonite
and albite in the immediate vicinity of the Buckeye mine
and recognized both jadeite and aragonite elsewhere in
the unit that contains the Buckeye mine, but he found no
relationship between metamorphic grade and texture
(Raymond, written commun., Feb. 6, 1988). Huebner
(1967), Hein and others (1987), and Hein and Koski (1987)
provided many other relevant observations, but no
maps.

 

The vicinities of the Buckeye and Ladd mines are each
characterized by sandstones containing an unusual abun-
dance of chert (ﬁg. 1A), most of which is interbedded
with argillaceous layers. Each area contains one large
and several smaller manganese deposits. The eastern
chert-rich sequence contains the large north Buckeye
orebody and the smaller south Buckeye, Sulphur Gulch,
Tiptop, Grummett, Grummett-Knox, Liberty, and
Moran Brothers deposits, each closely associated with
the chert. The western (Ladd) sequence is similar. These
chert sequences pinch out and are truncated by shear
zones to the west (Buckeye) and to the east (Ladd
sequence). The similarities in outcrop pattern, regional
strike (west—northwest), and alignment of both chert and
manganese horizons might suggest that the eastern and
western districts were once a continuous manganiferous
horizon that was broken into distinct segments (depos—
its), but mapping by Raymond (1973a,b; written com—
mun., Feb. 6, 1988) shows that the deposits lie in several
lithologically distinct units.

The Buckeye mines lie in the northern side of Buckeye
Gulch, in an area characterized by steep slopes, a high
ratio of rhythmic chert to graywacke, and prominent
outcrops (reefs) of northerly dipping massive chert (ﬁg.
13). Sulphur Gulch, to the north, is characterized by
gentler topography and by sheared and altered gray-
wacke with much less chert. South of the chert-rich
section, near the bottom of Buckeye Gulch, are blocks or
pods of glaucophane schist and carbonatized ultramaﬁc
rock in sheared graywacke, below a layer of coarse
conglomerate. The south side of the gulch is graywacke
with distinctly less chert than the north side. Raymond
(1973a) recognized that this terrain consists of broken
units and subdivided the area into three broken forma-
tions, which we here use informally: the graywacke-rich
Sulphur Gulch, the chert-rich Grummett, and the south-
erly Oso broken formations (Kstg, KJfg, and KJfo,
respectively, in ﬁg. 18). We have informally adopted
Raymond’s subdivision because it is useful in describing
the graywacke-rich and chert-rich units in the vicinity of
the Buckeye deposits. The Grummett broken formation
contains the Buckeye mines and consists of the ﬁrst
800—900 in of the Buckeye stratigraphic section by Hein
and Koski (1987, p. 723), from the conglomerate at the
base to approximately the greenstone. Glaucophane
schist, greenstone, and carbonate-rich blocks are either
erosional remnants of an overthrust unit or slivers of
basement brought up along faults bounding the Grum-
mett broken formation. In either case, the composition
and metamorphism of these blocks do not relate directly
to the history of the area.

To describe the degree to which the bedding of the
graywackes, chert, and manganiferous sediments was
disrupted, we adopted the classiﬁcation of Raymond

12 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

(1984, his ﬁg. 1, classiﬁcation type VIII). In this scheme
there is a progression from coherent (well-stratiﬁed)
units through broken units (in which the fragments are
moved slightly relative to each other) and dismembered
units (in which there is no suggestion of the initial
relative positions) to a melange (in which disoriented
fragments are engulfed in a matrix). We understand that
there is no consensus on the terminology used to describe
complexes and melanges; for instance, the classiﬁcation
of Cowan (1985) places more emphasis on lithologic
types. We chose the scheme of Raymond because we
found it useful in describing the disruption of unusual
manganese—rich lithologies.

Near the Buckeye mine, the degree of stratigraphic
continuity decreases with decreasing scale of the domain
considered. Raymond (1973a) mapped the regional units
as broken formations. Within the Grummett broken
formation, the chert bodies have irregular shapes in
outcrop pattern (see Trask, 1950, pls. 4, 17) that suggest
intricate folding; the chert appears to ﬂoat in sandstone
yet retains a strike and dip that is predominantly parallel
to the regional trend. One of us (Huebner) had difﬁculty
following sedimentary contacts on a scale exceeding
several hundreds of meters because the chert reefs
consist of slightly displaced blocks (called phacoids by
Raymond, 1973a, 1977). At the scale of the chert reefs,
the structure is transitional between that of the broken
and disrupted units. We will show that much of the ore,
on the scale of a hand specimen and thin section, is a
breccia with cataclastic to mylonitic textures. This
intensely disrupted material may represent the massive
ores described by some of our predecessors. Neverthe-
less, layered and laminated ore, intact on at least the
scale of a hand specimen, does occur. Some breccia
fragments in the disrupted ore are layered and lami-
nated, evidence that they were once part of a well-
layered sediment.

Despite disruption of the stratigraphic section, the
nonmanganiferous host rocks provide important clues to
the sources of sediment, the nature of the depositional
environment, and the subsequent metamorphism. Most
of the sediments are either detritus from a continent or a
magmatic arc or skeletal debris from marine pelagic
organisms. Deposition of each type was predominantly
by turbidity ﬂows. Metamorphism was at very high
pressures and low temperatures, implying rapid subsid—
ence and uplift.

REGIONAL LITHOLOGIES

Coarse conglomerate occurs at the base of the exposed
Grummett broken formation in the western part of
Buckeye Gulch. Rounded cobbles of black and green
chert, white vein quartz, feldspar porphyry with green

 

matrix, and shale form clasts that are suspended in a
matrix of greenish graywacke. This conglomerate is
probably from the same provenance as the matrix of the
massive graywacke elsewhere in the Grummett unit and
may incorporate earlier Franciscan sediments.

The turbidite units are dominantly massive graywacke
and include minor laminated siltstone, shale, and
intraformational conglomerates in which shale chips are
suspended in greenish-gray to gray graywacke matrix.
Graded beds are rare; one unusually complete unit
grades from conglomerate to shale (Huebner, 1967, p. 28)
and indicates that, locally at least, the north-dipping
beds are not overturned. Ten representative graywackes
from the Grummett broken formation and its contact
zones were examined in thin section (table 3) and by XPD
(table 4). All graywackes retain clastic texture. The
framework consists of albite, quartz, chert, micropor-
phyries, shale and clay chips (some black and of moder-
ately high reﬂectivity, presumably due to organic mat-
ter), biotite, muscovite, reddish chromite, and zircon, in
decreasing order of abundance. Most quartz clasts are
monocrystalline and unstrained. Quartz clasts that are
polycrystalline or that do not extinguish uniformly under
crossed nicols are uncommon. The compositions of some
detrital minerals are summarized in table 5. Pyroclastic
components, if present, have been transported and
reworked by sedimentary processes. Postdepositional
processes included compaction of the matrix, corrosion of
framework grains, and incipient recrystallization, corre-
sponding to the textural grades 1 and 2 of Ernst (1971).
The fabric varies from essentially unfoliated to slightly
foliated in hand specimen (distinctly foliated in thin
section). Of the metamorphic minerals, lawsonite is
ubiquitous and develops in framework albite and in the
matrix. Other metamorphic minerals are concentrated in
the matrix. Chlorite (IIb structural type of Bailey, 1980),
phengitic muscovite, and titanite occur in every section.
Calcite and chloritized biotite occur in most sections;
brown stilpnomelane appears in several thin sections and
in one thin section was conﬁrmed by microprobe chemical
analysis (table 5). Glaucophane, clinopyroxene, and pale-
green Ti02 (anatase) are uncommon, and green tourma—
line is rare. J adeite, laumontite, pumpellyite, and clinop-
tilolite were sought but never identiﬁed in our samples.
Aragonite occurs in White veins in sheared or altered
graywacke and greenstone, which occur near the con-
tacts between the broken formations, but, in contrast,
coarsely crystallized (primary) calcite with rhombohe-
dral cleavage is the characteristic polymorph in gray-
wacke of the interior of the broken units. With the
exception of the albite-lawsonite association, petro-
graphic criteria for an equilibrium assemblage are lack-
ing. Representative chemical analyses of metamorphic
minerals are given in table 5. These minerals should be

REGIONAL GEOLOGY

13

TABLE 3.—Constituents of graywackes from the Buckeye deposit, California Coast Ranges
[Modes of clasts are visual estimates only, obtained from thin sections. tr, trace; x, present; -, not observed]

 

 

 

 

 

 

 

Sample B44 B110 65H10 65H27 651-135 65H72 651-173 65H74 65H110 65H119
Modes of clasts, in volume percent

60 50 55 50 40 35 35 50 45 55

5 10 10 10 20 30 15 10 20 5

20 25 20 20 25 20 25 20 25 20

Lithic fragments ............... <5 <5 <5 <5 <5 <5 <5 tr <5 tr
Groundmass ................... 15 15 15 20 15 15 20 20 10 20

Mineralogy

Quartz ........................ x x x x x x x x x x
Albite ........................ x x x x x x x x x x
Lawsonite ..................... x x x x x x x x x -
Glaucophane ................... x - x - - — x - x -
Ca—carbonate .................. x x x x x - x - x -
Biotite (chloritized) ............ x x x x x x x x x x
Chlorite ....................... x x x x x x x x x x
Muscovite ..................... x x x x x x x x x x
Stilpnomelane ................. x - - - - - - - - -
Spinel-chromite ................ x x - x x x x - - -
Titanite ....................... x - x - - x x - x x
Anatase. . x x - x x - - x - x
Zircon .................. . . . x x x - x x x x - -
Tourmaline .................... - x - x - - - x - -
Clinopyroxene ................. x - - - x - - - - -
Fe-oxide or hydroxide .......... x - x x - x x x x x
Green claylike material1 ........ x x x x — x - - x -

 

1Present in very sparse amounts.

regarded as index minerals rather than as metamorphic
assemblages. The critical associations are quartz + albite
+ lawsonite + calcite + chlorite and quartz + albite +
lawsonite + aragonite + chlorite.

The mineralogy of the framework elements and the
bulk chemistry of the graywackes from the Grummett
broken formation and close to the Buckeye deposit are
potential indicators of sediment source and depositional
environment. Some discriminants, reviewed by Taylor
and McLennan (1985), include analyses of modern sands
(Maynard and others, 1982) and Phanerozoic sandstones
(Dickinson and others, 1983). Interpretation of the Buck-
eye graywackes is ambiguous. The high proportion of
quartz clasts (table 3) suggests derivation from a conti-
nental block or uplifted orogenic zone; Without knowl-
edge of the original lithic clast population, which may
have been altered during diagenesis or metamorphism,
we cannot distinguish between these two environments.
The small value of K20/Na20 (table 6) suggests an
island-arc setting. However, the apparent island-arc
signature may merely reﬂect loss of K20 due to metaso-
matism. The range of rare-earth—element patterns for
seven Buckeye graywackes (see ﬁg. 7A; average values
in table 6; Huebner and Flohr, unpub. data) are similar to
patterns for the quartz-intermediate graywackes of Tay-
lor and McLennan (1985). This group includes recycled
orogens, continental blocks, and dissected magmatic arcs
but excludes a forearc basin deriving sediment from a

 

young andesitic arc. We favor a sediment source such as
a deeply dissected continental magmatic arc complex,
augmented by intraformational clasts of shale and chert
derived from near the site of deposition. In studying
graywackes from throughout the Franciscan Complex,
Dickinson and others (1982) observed a smaller propor-
tion of monocrystalline quartz clasts and tentatively
concluded that potassium feldspar was originally present
but has been replaced by albite. They concluded that the
Franciscan graywackes of the Diablo Range were
derived from the southern portion of the Sierran-
Klamath magmatic arc, near the Mojave block. Seiders
and Blome (1988) examined conglomerates in the Fran-
ciscan Complex of the Diablo Range. They concluded
that conglomerates rich in volcanic clasts were trans-
ported directly westward and that chert-rich conglomer-
ates were transported southward. However, they did
not report data for the conglomerates shown in ﬁgure
13.

The contacts between the graywacke units and the
thick sequences of interbedded chert and metashale,
although not well exposed, were originally thought by
Huebner (1967) to be sedimentary. Raymond (1974)
suggested that, at least locally, some contacts between
chert and graywacke are sedimentary, yet Raymond
(1973a; written commun., 1988) stated that most of the
intraformational contacts between chert and graywacke
are now sheared as a consequence of postdepositional

14 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

TABLE 4. —Charactem'zation of metasandstones, metacarbonates, and metavolcanic rocks from the Buckeye deposit, Caltfomia Coast Ranges,
by X—ray powder diﬂmctometry

[Color terms from Goddard and others (1970). Meta-vole, metavolcanic rocks; meta-carb, metamorphosed carbonate rocks. Estimated intensities of X-ray powder diffraction intensities: V, very;
S, strong; M, moderate or moderately; W, weak; -, not observed]

 

 

 

 

 

 

 

 

 

 

Sample B44 B110 65H18 65H20 65H27 65H28 65H28
Color Greenish gray Light olive gray Light olive gray Light olive gray Light olive gray Light olive gray White
Foliation None None Weak None None None None
Rock type Graywacke Graywacke Graywacke Meta-vole Graywacke Graywacke Vein
Quartz ............ VS VS M S VS S VW
Low albite ......... S VS M VS S S —
Chlorite ........... VW VW M M MS M -
Muscovite ......... VW VVW W - - - -
Calcite ............ - W - M VW - W
Aragonite .......... - - - - - - M
Lawsonite ......... - - — - - W -
Sample 65H29 65H29 65H29B 65H30A 65H30B 65HSOC 65H311
Color Grayish olive Very light gray Olive gray White Light gray Olive gray Mottled gray brown
Foliation None None N one None Laminated Laminated Weak
Rock type Greenstone Vein Greenstone Vein Meta-carb Meta-earl) Meta-carb
Quartz ............ MW VVW M M S VS MS
Low albite ......... - - W - - - -
Chlorite ........... S - MS - VVW VW -
Muscovite ......... - - - - - - -
Calcite ............ VW M - M VVW MS MW
Aragonite .......... W W - M W M MW
Lawsonite ......... - - VW - - - -
Dolomite .......... - - - — VS S -
Sample 65H33 65H33 65H34 65H35 65H43 65H43 65H51
Color Medium dark gray White Light olive gray Light olive gray Dark gray White Pale yellowish brown
Foliation Slight None Slight Slight None None None
Rock type Greenstone Vein Graywacke Graywacke Hornfels Vein Hornfels
Quartz ......... VS S S VS S W M
Low albite ...... MW - VS VVS - - -
Chlorite ........ MW - M VW MW - M
Muscovite ...... — - VW? VVW? W - —
Calcite ......... - VVW - - - W W
Aragonite ....... - W - - — MS M
Lawsonite ...... W - VW - W - VW
Glaucophane - - - - MW — -
Sample 65H52 65H55 65H55 65H56 65H57 65H59 65H60
Color Light olive gray Light olive gray White Light olive gray Light olive gray Olive gray Dark gray
Foliation None Slight None None Slight None None
Rock type Greenstone Graywacke Vein Meta-carb Graywacke Meta-vole Hornfels
Quartz ......... M MS W M S MS S
Low albite ......... S M - - S M -
Chlorite ........... M W - M W MS W
Muscovite ......... - - - - - - -
Calcite ............ - - S - - - VVS
Aragonite .......... W - VVW MW - - VW
Lawsonite ......... - - - W - - -
Sample 65H65 65H72 65H73 65H74
Color Light olive gray Medium gray Medium gray Greenish gray
Foliation Slight None None Slight
Rock type Graywacke Graywacke Graywacke Graywacke
Quartz ............ VS VVS VVS VS
Low albite ......... VS W VS S
Chlorite ........... M VVW M M
Muscovite ......... - - - W
Calcite ............ - - VVW -
Aragonite .......... — - - -
Lawsonite ......... - - W -

 

1 XPD patterns from 65H31 also contain a weak reﬂection at 2.894 A.

REGIONAL GEOLOGY

15

TABLE 5. —Compositions of minerals from nonsiliceous host rock, in weight percent, from the Buckeye

deposit, California Coast Ranges, determined by electron mzcroprobe analysis

[Total Fe reported as FeO for all minerals. Ferrous and ferric iron calculated assuming ideal stoichiometry of 3 cations per 4 oxygens for
chromites. Analyses used are average of 3 points for all minerals except chlorite in sample 65H73; nd, not detected; na, not analyzed.
H2O was not analyzed, resulting in low oxide sums for tourmaline, muscovite, glaucophane, stilpnomelane, lawsonite, chlorite, and
possibly titanite; C02 was not analyzed, resulting in low oxide sums for carbonate minerals]

 

Detrital minerals

 

 

 

 

 

 

 

 

 

Titanite1 Augitez Chromite3 Tourmaline Muscovite
Sample 65H73 B44 65H73 B110 65H30 65H31 65H27 65H74
Si02 ........ 29.6 50.9 0.03 0.05 nd nd 36.3 50.5
A1203 ....... 1.09 4.28 35.4 9.86 2.94 8.30 31.7 28.8
Ti02 ........ 36.1 .43 .05 .22 nd nd .29 .28
V203 ........ .50 .05 .20 .24 .21 .21 .15 nd
Cr203 ....... na .24 32.6 51.9 61.1 59.2 .03 na
FeO ........ 1.24 5.81 17.1 33.4 28.6 23.5 6.73 3.15
MnO ........ 17 .12 .24 .57 .52 .41 .09 nd
MgO ........ 10 15.4 14.6 2.33 5.38 7.43 8.05 2.50
ZnO ........ nd na na na .18 .16 na nd
NiO ........ nd nd .10 nd nd .05 nd nd
CaO ........ 28.2 21.9 nd nd .12 .13 1.34 .05
BaO ........ nd nd na na na na na .23
N320 ....... 03 .29 nd nd nd nd 1.81 .13
K20 ........ na nd na na na na nd 9.25
Sum .. 97.03 99.42 100.32 98.57 99.05 99.39 86.49 94.89
FeO ........ - - 15.1 29.1 23.1 20.8 - -
Fe203 ....... - - 2.24 4.81 6.13 2.93 — -
Metamorphic minerals
Glaucophane Stilpnomelane Albite4 Lawsonite4 Chlorite5 Titanite6 Dolomite7 Calcite7 Aragonite7
Sample 651173 B44 65H73 65H73 65H73 65H73 65H3l 65H31 65H30
Si02 ........ 56.2 48.0 67.0 37.7 28.0 30.7 na na na
A1203 ....... 8.50 7.36 20.7 31.4 18.0 5.46 na na na
Ti02 ........ . 10 nd nd .06 .09 30. 8 na na na
V203 ........ .05 .05 nd nd nd .47 nd nd nd
Cr203 ....... na na na na na na na na na
FeO ........ 15.3 24.1 .05 .54 24.6 1.00 6.08 .46 nd
MnO ........ .14 1.87 nd .06 .69 nd .48 1.88 nd
MgO ........ 9.05 5.84 .13 .08 15.5 .16 17.7 .52 .18
ZnO ........ nd nd na na . 10 nd nd nd nd
NiO ........ nd nd nd nd nd nd nd nd nd
CaO ........ 1.23 .89 .65 16.5 .13 28.5 29.3 53.1 54.4
BaO ........ nd .21 nd nd nd nd nd nd nd
NaZO ....... 5.97 .72 11.0 .04 .03 .04 na na na
K20 ........ .04 .59 .17 nd .09 .07 na na na
Sum . . .. 96.58 89.63 99.70 86.38 87.23 97.20 53.56 55.99 856.93

 

1 Fragmented titanite in groundmass.

2 Augite occurs with quartz.

3 First two chromite analyses are of the most aluminous chromite and most chromium-rich chromite, respectively,
analyzed in the graywackes. Third and fourth analyses are from carbonates.

4 Lawsonite occurs as laths within albite clast.

5 Pale green chlorite with Mg/(Mg+Fe+Mn) = 0.531; 4 points averaged.

5 Anhedral titanite grains within chloritized clast.

7 Calculated CO2 values, assuming ideal stoichiometry, are 46.4 weight percent for dolomite, 43.8 weight percent for
calcite, and 43.5 weight percent for aragonite.

8 Sum includes 2.35 weight percent SrO.

soft-sediment and tectonic deformation. Hein and Karl
(1983, p. 35—36) and Hein and others (1987 , p. 210) also
thought that the contacts were sedimentary. Hein and
Koski (1987, p. 725) noted “gradational sedimentary
contacts between the underlying sandstone and the
radiolarian chert-argillite sections.” Subsequently they

reported that “the ﬁeld relations clearly show deposi-
tional contacts between the graywacke sandstones and
the rhythmically bedded chert-argillite sections” (Hein
and Koski, 1988, p. 470). However, the undulatory and
lenslike outcrop pattern (ﬁgs. 1A,B) of slivers of chert in
graywacke may be the result of soft-sediment foundering

16

MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

TABLE 6. —Chemistry and mineralogy of selected protoliths from the Buckeye deposit, California Coast Ranges

[Chert, average of 35 cherts from the Ladd-Buckeye district (Hein and others, 1987); analyses
by XRF, ICP, and INAA methods; metashale, average of 7 metashale samples, Ladd-
Buckeye district (Hein and others, 1987); graywacke, average of 7 graywacke samples (B44,
B110, 65H27, 65H35, 651-172, 651173, 65H74), Buckeye district; modes for these samples are
given in table 3. W237—, sample number preﬁx assigned by Sample Control, Branch of
Geochemistry. USGS. ICP (inductively coupled plasma atomic emission spectrometry) by H.
Kirschenbaum; INAA (instrumental neutron activation analysis) by J .N. Grossman; FAAS
(ﬂame atomic absorption spectrometry) by H. Kirschenbaum; SQES (semiquantitative
DC-arc emission spectroscopy) by Janet Fletcher; XRF-EDA (X-ray ﬂuorescence with

energy dispersive analysis) by R.G. Johnson; GAAS (graphite-furnace atomic absorption
spectrometry) by M.W. Doughten; Sep (extraction preceded analysis) by MW. Doughten,
B. Libby, and P.J. Aruscavage; Coul., coulometric determination; Pen., modiﬁed Penﬁeld
method; Grav., gravimetric determination; pct, percent; ppm, parts per million; ppb, parts
per billion; —, chemistry not reported or not available or mineral not observed in Buckeye
samples Mineral mode not reported for chert and metashale. All Buckeye samples have less
than 2 ppm Rb by XRF and less than 2 ppm Se by INAA. Estimate of mineral modes from
visual examination of X-ray powder diffraction patterns]

 

 

Sample B50 B64 B112 BS4 B95 B104 651179 Chert Metashale Graywacke
W237— 507 503 506 505 511 510 509

SiO2 ........... pct ........ ICP ......... 9.4 38.7 17.0 33.2 16.1 8.3 22.6 94.39 58.13 67.9
A1203 .......... pct ......... ICP ......... .40 .21 2.14 3.25 1.25 .070 .047 1.71 16.0 12.8
F62031 ......... pct ......... INAA ....... .187 .164 1.47 2.15 .892 .189 .095 1.208 8.58 5.3
MgO .......... pct ......... ICP ......... 1.2 .94 3.83 9.28 2.55 .66 2.28 .59 3.21 2.5
MnO ........... pct ......... ICP ......... 60.4 31.9 47.6 35.8 64.9 79.5 63.4 .21 .27 .22
CaO ........... pct ......... ICP ......... .32 4.27 1.68 1.04 .18 .12 .04 .12 .89 1.6
NaZO .......... pct ......... INAA ....... .008 .015 .058 .023 .015 .007 .009 .28 1.52 4.4
K20 ........... pct ......... FAAS ....... <1 <1 <1 <1 <1 <1 <1 .34 4.08 1.0
002 ........... pct ......... Coul ......... 23.6 23.0 19.8 3.09 5.06 4.69 .66 - - .25
H20+ ......... pct ......... Pen. ........ 3.2 .3 4.6 9.0 3.7 2.7 7.5 2.72 26.0 33.3
H20— ......... pct ......... Grav. ....... .1 .1 .9 1.8 .1 .1 .4 - - -
P205 ........... pct ......... ICP ......... <.06 <.06 .12 .37 <.06 <.06 <.06 .05 .35 .09
Ti ............. ppm ........ ICP ......... 93 81 370 970 390 72 63 419 5215 3342

V ............. ppm ........ ICP ......... 91 <5 1200 480 470 460 140 20.4 128.3 127
Co ............ ppm ........ INAA ....... 6.61 1.72 22.1 13.1 26.2 8.8 5.26 8.0 27.8 16.6
Ni ............. ppm ........ ICP ......... 170 79 350 150 840 650 740 22.4 105 53
Cu ............ ppm ........ ICP ......... <5 11 110 64 59 8 <5 44 130 21
Zn ............ ppm ........ FAAS ....... 530 110 320 220 1300 2700 630 24.8 105.4 92
W ............. ppm ........ Sep, ICP .... 40 <1 <1 1.7 22 30 89 - - <1
Th ............ ppm ........ INAA ....... .214 .18 .97 2.26 1.48 .180 .073 - — 6.04
U ............. ppm ........ INAA ....... .55 <.3 1.23 .59 .84 1.07 1.05 - - 1.84
Te ............. ppm ........ Sep, GAAS . . .03 <.02 .13 .13 .14 <.02 <.02 - - -
Cr ............. ppm ........ GAAS ....... <20 14 7.0 16 12 2.3 7.5 10.6 89.4 170
As ............ ppm ........ FAAS ....... 94 5.0 63 110 120 190 210 — - 10
Cd ............ ppm ........ Sep, GAAS .. 1.5 .56 1.5 .43 .64 1.5 3.6 - - -
Sb ............. ppm ........ INAA ....... 12.0 .122 1.62 14.0 6.01 22.6 24.1 - - 1.31
Mo ............ ppm ........ Sep, ICP . . . . 36 1.2 4.0 740 3.5 73 29 — — <2
Be ............ ppm ........ Sep, GAAS .. <2 <2 .28 .52 .43 .25 .78 - - .96
B ............. ppm ........ SQES ....... 71 <3.2 230 580 390 100 74 - - —
Li ............. ppm ........ Sep, FAAS .. <50 12 12 42 12 <50 <50 12.9 36.3 42
Sc ............. ppm ........ INAA ....... .995 .419 2.45 5.51 4.20 1.547 1.06 - - 14.9
Sr ............. ppm ........ ICP ......... <2.5 120 33 27 3.0 <2.5 <2.5 - - 156
Ba ............ ppm ........ XRF-EDA. . . 21 5200 70 400 74 <5 <5 214 870 337
Zr ............. ppm ........ XRF-EDA. .. <5 <5 <5 50 33 25 <5 — - 137
Hf ............. ppb ......... INAA ....... 106 450 331 830 490 177 <200 - - 3420
Nb ............ ppm ........ Sep, ICP <1 <1 <1 2 <1 <1 <1 - - 8
Ta ............. ppb ......... INAA ....... 17 16 90 237 88 <20 <100 - - 624

Y ............. ppm ........ ICP ......... <5 <5 <5 15 <5 9 <5 6.1 37.4 16
La ............ ppm ........ INAA ....... .71 1.59 8.2 14.0 3.26 14.7 .82 - - 20.0
Ce ............ ppm ........ INAA ....... 1.66 1.97 12.2 26.9 14.3 4.5 <.2 - - 36.9
Nd ............ ppm ........ INAA ....... <4 <6 8.0 15.2 3.9 12.6 <7 — - 16
Sm ............ ppm ........ INAA ....... .248 .454 1.91 2.99 .84 3.15 .117 - - 3.60
Eu ............ ppb ......... INAA ....... 57 118 420 690 195 782 <90 - - 818
Tb ............ ppb ......... INAA ....... 22 79 287 422 196 526 35 - - 473
Ho ............ ppb ......... IN AA ....... <600 <700 <900 670 <900 940 <900 - - -
Yb ............ ppb ......... INAA ....... 151 550 860 1230 610 1460 810 - - 1700
Lu ............ ppb ......... INAA ....... 28 86 142 226 138 246 169 - - 256

 

1 Total Fe reported as Fe203.
2 Loss on ignition.

3 Average of total water content of 6 samples, excluding B44; for which there was insufﬁcient sample for analysis.

REGIONAL GEOLOGY 17

TABLE 6. —Chemistry and mineralogy of selected protoliths from the Buckeye deposit, Cahform'a Coast Ranges—Continued

 

 

 

Sample B50 B64 B112 B84 B95 B104 65H79 Chert Metashale Graywacke
W237— 507 503 506 505 511 510 509
X-ray modes

Quartz .......................................... - 35 2 - - - -
Mn—carbonate .................................... 80 65 70 25 25 15 2
Caryopilite ...................................... - — 25 65 20 2
Chlorite/smectite ................................ - - 5 5 - - -
Gageite ......................................... 20 - - - 45 25 90
Hausmannite .................................... — - - - - 60 10
Braunite ........................................ - - - - 55 - -
Barite .......................................... - - 1 - - - -

 

4 Estimate based on thin-section optics; gageite not observed in X-ray pattern.

or tectonic dismembering. We conclude that transitions
between graywacke and chert-metashale sequences did
occur in the sedimentary environment but that some of
the complicated smaller scale features may be of later
origin.

Chert layers consist of more than one lamination and
form beds that are commonly 2 to 20 cm thick. Chert
beds are separated by metashale beds (or interbeds),
which are several millimeters to 10 cm thick (ﬁg. 2A). In
a general way, the thicknesses of the metashale interlay—
ers are correlated with the color of the chert. Interlayers
associated with red chert are comparable in thickness to
the red chert layers, interlayers associated with green
chert are distinctly thinner than the chert beds, and the
metashale partings within white and gray chert beds are
insigniﬁcant (ﬁg. 23), resulting in massive character.
Within a lens of chert and metashale layers, individual
beds of chert are of uniform thickness rather than
thickening in the noses of tight folds or thinning as they
approach the lateral outcrop margin of a lens. In unusual
cases, individual beds bifurcate or terminate abruptly
(see Taliaferro and Hudson, 1943, their ﬁgs. 3—4). These
observations, combined with the assumption that color
was associated with purity of chert, led early students,
including Huebner (1967), to think that the rhythmic
layering developed during diagenesis. The recognition of
laminations and graded bedding within individual layers
of chert (Hein and others, 1987a, p. 213) suggests that
these chert layers are siliceous turbidites (see N isbet and
Price, 1974).

Folding in the cherts is spectacular (see Huebner,
1967, ﬁgs. 11—8 and 11—10). Sharp folds are common; units
of chert on outcrop scale exhibit uniform and symmetric
chevron folding (see Johnson and Ellen, 1974). In other
places, several sequential chert layers are contorted into
S—shapes, yet adjacent layers are not affected. Individual
chert layers pass smoothly and continuously about the
noses of folds. There is no cleavage. These latter features
are characteristic of soft-sediment deformation of layers
that remained intact. The chert beds are now out by

 

quartz veins that do not extend into the adjacent meta-
shale layers. Rare syneresis (shrinkage) cracks in indi—
vidual chert layers and a brecciated chert reef in which
brown chert fragments are engulfed by white silica were
noted by Huebner (1967). However, we have the impres-
sion that there is little evidence for through-going frac-
tures through which ﬂuids could move perpendicular to
layering, although the metashale layers or their contacts
with the adjacent chert might provide pathways through
which ﬂuids could move parallel to the bedding.

Red cherts have a framework of spherical or elliptical
skeletons of radiolaria set in a matrix of radiolarian
fragments, iron oxides and hydroxides, and clays.
Although the clays may represent a slow yet continuous
background deposition of ﬁne clastic detritus, coarse
detritus of the kinds found in the graywackes is absent.
In contrast to the red chert layers, the green, gray, and
white cherts contain only a small proportion of oxide-
hydroxide-clay component, and skeletal debris is pre-
served only as ghosts in microcrystalline quartz. Within
chert layers of all colors, laminations are indicated by
variable concentrations of opaque “dust.”

The north Buckeye orebody is enclosed in white chert
(locally stained black by supergene manganese oxides)
that lacks the metashale interbeds and therefore is
described as massive. Actually, the massive chert is
composed of chert laminations or, adjacent to the ore—
body, chert interlaminated with the manganese silicates
santaclaraite, caryopilite, and taneyamalite. These lam-
inations impart a ﬁssility that is particularly evident at
the surface, above the north orebody. The massive chert
reef and the conformably enclosed orebody are folded in
an S-shaped kink without major fractures perpendicular
to the chert-orebody contact (see Huebner, 1967, ﬁg.
II-lO; Taliaferro and Hudson, 1943, ﬁg. 8).

The siliceous phase in all cherts and cherty materials
examined by us is quartz. Hein and Koski (1987) reported
cristobalite-like peaks in X-ray diffraction patterns of
silica samples that were associated with manganese.
They concluded that emplacement of manganese within

18 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

 

FIGURE 2,—Representative chert from near the Buckeye deposit. A,
Rhythmic gray chert. Some chert beds are clearly composite. The
chert layers are thick relative to the layers of metashale. Scale is 16
cm long. B, Surface outcrop of north orebody in “massive” white
chert. The white chert is actually laminated and has some ﬁssility.
Width of open cut about 5 In; view is eastward. C, Sample 65H66.
Laminated chert from contact with orebody, southeast wall of cut.
Laminations are parallel to regional layering in the chert reefs and
are caused by varying concentrations of manganese silicates,
rhodochrosite, and “dust” (less-than-Z-pm particles of clays or oxy-
hydroxides). The white bleached zones are channels through which
ﬂuids passed, leaching manganese minerals and dust and leaving
behind pure chert that is more coarsely crystallized than the impure,
laminated chert.

    

1 I 1m s

i!
agrmc 1

REGIONAL GEOLOGY 19

the sedimentary pile arrested the transformation of opal
to quartz. In contrast, we found no evidence of any relic
opaline phases in the rhythmic chert, in massive chert, in
chert interlayered with manganese minerals at the mar—
gin of the orebody, or intimately intermixed with man-
ganiferous minerals within the orebody. In the ore lens,
much (presumed) skeletal debris has been replaced by
rhodochrosite or caryopilite, compositions that have
been conﬁrmed by energy—dispersive X-ray analysis.The
best preserved textures are “necklaces” of anhedral
carbonate crystals that lack the orderly arrangement of
skeletal plates of organisms; these necklaces may not
originally have been siliceous. In X-ray diffraction pat-
terns, quartz is well crystallized. The quintuplet of peaks
resulting from the (122), (203), and (301) reﬂections is
generally well resolved, and the few cases of poor
resolution can be attributed to admixed phases. In
particular, the (030) reﬂection of rhodochrosite coincides
With the (212) reﬂection of quartz. The crystallinity index
a/b (table 7) based on these quartz reﬂections, which is a
qualitative measure of grain size, structural order, and
goniometer slit-widths and alignment, averages about 10
on a scale of 11.6. Thus, our values are greater than the
values of 6 to 7 (on the scale'of 10 proposed by Murata
and Norman, 1976) reported by Hein and others (1987).
The existence of such well-crystallized quartz is inconsis-
tent with the presence of opaline phases. In a ﬁnal
attempt to conﬁrm the report of opaline silica, we
dissolved the carbonate from 11 samples, using 2 N HCl,
but found no reﬂections attributable to cristobalite or
tridymite. Mineral impurities found in the cherts and
quartz-rich materials are rhodochrosite, taneymalite,
caryopilite, chlorite, smectites with spacings from 15A to
25A, acmitic clinopyroxene, and iron oxyhydroxide
(table 7).

The discovery of chert and volcanism in modern deep
ocean basins has led to the use of major-element compo-
sitions to distinguish between pelagic (organic), detrital
(abiogenic), and hydrothermal sources of silica (Adachi
and others, 1986; Yamamoto, 1987). By this reasoning,
the hydrothermal component is Fe- and Mn-rich and Al-
and Ti-poor, the detrital component is Al- and Ti-rich,
and the pelagic component has low concentrations of all
four elements. Hein and others (1987) presented average
compositions for 35 cherts, 4 argillaceous cherts, 7 sili-
ceous argillites, and 7 argillites (metashales) from the
Ladd-Buckeye district. The compositional contrast of
these rocks with Franciscan cherts and shales deposited
above metabasalt is striking (Yamamoto, 1987). At the
Buckeye, the cherts appear to be pelagic with a minor
detrital, rather than hydrothermal, component. The
interbedded metashales have A1203 and TiO2 contents
that are similar to those of average shales and gray-
wackes; thus the Buckeye metashales appear to have a

 

signiﬁcant detrital component, perhaps related to the
ﬁnest fraction of the graywacke. Later we will show that
the rare—earth-element (REE) patterns of the cherts and
shales mimic the patterns for the graywackes, reﬂecting
an REE-rich detrital component. However, because the
detrital (graywacke) component tends to be REE-
enriched, even a small proportion of detrital component
could obscure the contribution of a hydrothermal compo-
nent to the REE pattern for the chert or metashale.
Complete trace-element data for all three rock types
might permit better estimates of the proportions of
detritus.

PROVENANCE

Sediments of the Grummett broken formation, host of
the Buckeye manganese deposits, were predominantly
elastic, both biogenic and abiogenic. The metashale inter-
beds associated with the chert may in part represent a
chemically precipitated component, perhaps of hydro-
thermal origin, but the proportion of such a component,
if present, is uncertain. The biogenic ﬂux consists of
radiolarian debris and whatever pelagic contribution is
contained in the metashales and thus was ultimately
derived from a marine water column. The abiogenic ﬂux
is represented by the conglomerates, metagraywackes,
and associated shales, and a background component of
clays that formed the metashales. Clasts of chert, shale,
porphyritic volcanic rocks, plagioclase, coarse quartz,
and chromite found in the conglomerate and graywacke
cannot have come from a single lithologic source. Three
or four sources are needed: one of low metamorphic
grade, perhaps Within the Franciscan Complex itself; a
quartz-bearing igneous or metamorphic terrain, perhaps
a continental margin or magmatic arc; an ultramaﬁc (or
ophiolitic) terrain that could provide the detrital chro—
mite; and an ultramaﬁc to maﬁc source that could provide
the iron— and magnesium-rich graywacke matrix.

REGIONAL METAMORPHISM

The presence of lawsonite + quartz + albite + chlorite
and minor glaucophane, combined with the absence of
zeolite, zoisite, and epidote, indicates metamorphic con-
ditions between 175 and 300 °C and greater than 4 kbar
(ﬁg. 3; see also Cotkin, 1987). Because aragonite is
known to transform readily to calcite at low pressure, the
presence of calcite within the Grummett broken forma-
tion is not diagnostic. However, the presence of
aragonite-bearing veins in the graywacke near the
boundaries of the Grummett broken formation is an
indicator of the minimum pressure achieved in the veins

20 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

TABLE 7. —Characterization of siliceous samples from the Buckeye deposit, California Coast Ranges, by X-my powder diﬂmctometry
[Color terms from Goddard and others (1970). Estimated intensities of X-ray powder diffraction intensities: S, strong; M, moderate or moderately; W, weak; VW, very weak; -, not observed.
a/b, crystallinity index explained in the text]

 

 

 

 

 

 

 

 

Sample B10 B64 B25b B72 B74 B77 B93
Color Moderate Pink with Very pale Pale red Pink, brown, Pale brown Pale olive
yellowish gray veins orange and gray
brown
Rock type Chert Carbonate Vein Chert Carbonate Chert Argillite
Quartz .................. S M M MS M S S
Chlorite ................. - - - - - - M
Caryopilite .............. MW - - - - - -
Smectite ................ W - - - VW VW -
Biotite .................. - - - - - — MW
Muscovite ............... - - - - - — MW
Acmitic pyroxene ........ - - - - - - MW
Rhodochrosite ........... VW S W S M -
Ca—Mn carb .............. - - — - 7 W - -
Hematite ............... - - - ? W - ‘3 W -
Other ................... - - - - - ? W -
a/b ..................... 10.7 6.4 7.4 9.8 (1) 10.2 10.0
Sample B96 B97a 65H23 65H24 65H66Aa 65H66Ab 65H66Ac
Color Dark Grayish Light Pinkish Pinkish Olive gray Dark
yellowish orange brownish gray gray yellowish
brown and pink gray brown
Rock type Carbonate Carbonate Chert Chert Chert Chert Chert
Quartz ................... M M S S S S S
Chlorite .................. VW — - - - - —
Taneyamalite ............. - — - - - MW MW
Smectite ................. W W - - - W -
Acmitic pyroxene ......... MW - - - - - -
Rhodochrosite ............ S S - - W - -
Other .................... W - - - - W -
a/b ...................... 10.0 (I) 11.5 9.0 10.0 11.1 12.0
Sample 65H67 65H68a 65H68b 65H69a 65H69b 65H690 65H69d
Color Very light Pale yellowish Dark yellowish Medium gray White Light olive Brownish
gray brown brown gray gray
Rock type Chert Chert Chert Chert Chert Chert Chert
Quartz ................... S S S S2 S S S
Taneyamalite ............. - - - MW - W W
Smectite ................. - W VW W W - -
Acmitic pyroxene ......... - W - - - - -
Rhodochrosite ............ - - W - MW — W
Other .................... - ? W - - - ~ -
a/b ...................... 11.4 9.5 10.5 8.1—8.8 10.1 11.1 9.4
Sample 65H69e 65H69f 65H71 65H76a 65H76b 65H76c 65H171
Color Speckled Brownish Very light Pale yellowish Pale reddish Dark Brownish
yellowish gray gray brown brown yellowish gray
gray brown
Rock type Chert Chert Chert Chert Chert Chert Chert
Quartz ................... S S S S S S S
Chlorite .................. - - - - - VW W
Taneyamalite ............. - W - - - - -
Smectite ................. - - W ? W - —
Acmitic pyroxene ......... — - MW MW MW - VW
Rhodochrosite ............ W W - — - - -
, a/b ...................... 11.4 11.6 10.2 10.3 10.2 10.1 10.65

 

1 Unable to determine crystallinity index because of interference by (030) peak of rhodochrosite.
2 Weak reﬂection present at 218° is too sharp for opal.

PROTOLITHS 21

 

12

10—

Pressure (kbar)

 

 

 

 

250 350
Temperature (°C)

FIGURE 3.—Pressure-temperature grid for metagraywacke. All reac-
tions are written with the high-temperature assemblage on the right
side of each equation. 1, heulandite = lawsonite + quartz + ﬂuid. 2,
jadeite + quartz = low albite. 3, aragonite = calcite. 4, lawsonite +
glaucophane = epidote + chlorite + albite + quartz + ﬂuid. 5,
lawsonite + albite = zoisite + paragonite + quartz + ﬂuid. 6,
glaucophane + clinozoisite + quartz + ﬂuid = tremolite + chlorite +
albite. In the vicinity of the Buckeye deposit, the assemblage is
lawsonite + quartz + albite + chlorite :- glaucophane. Aragonite is
nearby. Pumpellyite, jadeite, and members of the zoisite group are
absent. Pressures of 4 to 7 kbar at 180 °C or 10 kbar at 300 °C are
possible. We favor temperatures of no greater than 175—250 °C
(stippled area) for the Buckeye deposit because relic gel-like mate-
rials are present and a pervasive metamorphic fabric did not develop.
Curves largely after Cotkin (1987, ﬁg. 3); the calcite-aragonite curve
is from Johannes and Puhan (1971).

and tentatively is regarded as an indicator of minimum
pressure within the Grummett unit, 5 kbar at 180 °C
(Johannes and Puhan, 1971). The occurrence of jadeite +
albite + quartz in other samples from within the Grum-
mett broken formation (Raymond, 1973a,b) suggests
that the peak metamorphic pressure might have been 7—8
kbar (at 180 °C). Although temperature and pressure
estimates of such assemblages are uncertain, the preser-
vation of palimpsest features, the development of only an
incipient foliation, and the disordered state of the sheet
silicates suggest metamorphism at the low-temperature
end of the range of conditions, perhaps 150 to 200 °C and
7 to 8 kbar. Maintenance of relatively lOW temperatures
at such high pressures requires a low geothermal gradi—
ent. If hydrothermal convection of the kind proposed by
Lonsdale (1977) or Crerar and others (1982) was involved
in the deposition of manganese at the Buckeye, meta-
morphism could not have occurred in the presence of the
hydrothermal heat engine. This separation of processes
could have been brought about by allowing the heat
source to decay with time or by moving oceanic crust
with its mantle of sediments and manganese deposits
away from the heat source before subduction.

 

PROTOLITHS

The goal of this study has been to look back through
metamorphic and diagenetic overprints to characterize
the original sediments that resulted in the north Buckeye
and, by analogy, other orebodies of the Franciscan
Complex. Without knowledge of the original nature of
these materials, it is impossible to deduce their origins.
By volume, four phases (rhodochrosite, caryopilite,
braunite, and hausmannite) compose virtually the entire
orebody. Ten phases form monomineralic or almost
monomineralic layers with sedimentary characteristics.
Many of these phases also occur as mixtures that form
other layers. We are conﬁdent that six kinds of layers
and laminations have a sedimentary origin and thus
represent compositions, if not mineralogies, of distinct
sedimentary components, albeit now dehydrated and to
some extent recrystallized. Two more minerals probably
represent sedimentary components. Each of these sedi-
mentary protoliths is named on the basis of its charac-
teristic mineralogical component: chert, rhodochrosite,
caryopilite, chlorite, hausmannite, braunite, and perhaps
taneyamalite and gageite. (Nominal formulas for these
minerals are given in table 1.) A layer rich in diopsidic
acmite also has sedimentary characteristics. In addition,
talc occurs as a green lens in brown chert (B10) but is too
rare to be considered a component that inﬂuences the
composition of the orebody. In the following paragraphs,
we describe the occurrence of these materials, making
reference to the series of photographs arranged to show
increasing disruption of original layering (ﬁg. 4). Seven
homogeneous layers were sampled to yield 25- to 50—g
splits for initial major- and trace-element analyses. The
large sample size, in some cases, prevented removal of a
pure sedimentary component. These seven samples
(table 6) represent rhodochrosite, caryopilite, gageite,
hausmannite, and braunite protoliths. Layers of chlorite
and taneyamalite are too thin and of insufﬁcient lateral
extent to sample for bulk chemical analysis (1 gram)
Without contamination by other protoliths and are chem-
ically characterized only by electron microprobe tech-
niques. Average chert and metashale analyses can be
found in Hein and others (1987) and are summarized in
table 6. In describing the protoliths, we give the charac-
teristic X-ray reﬂections so that later workers will be
encouraged to identify these phases in ﬁne-grained mix-
tures, using X-ray powder diffractometry. We hope that
our observations will be found to be generally applicable
to many deposits of the Franciscan Complex and else-
where.

CHERT

Chert surrounds the orebody but is not interlayered
with the other protoliths. From this fact, we infer that

22 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

 

PROTOLITHS 23

FIGURE 4.—Slabbed hand specimens of different protoliths from the
Buckeye deposit, California, arranged to show increasing degree of
soft‘sediment deformation. Following the scale of Raymond (1984),
coherent units are disrupted, then dismembered, and ﬁnally only
relics of beds are seen engulfed in a once-ﬂuid matrix to form a
melange. A, Sample BS7. Alternating layers of caryopilite, haus—
mannite disseminated in rhodochrosite, and rhodochrosite. Layering
is still fairly coherent despite late polymineralic veins and recrystal-
lization of rhodochrosite (grayish-pink pods). B, 65H79. Thick layer
of gageite (black) in massive hausmannite (dark gray). Thin veins of
gageite originated in the thick layer. C, B25. Coherent layers of
caryopilite can be traced across hand specimen despite brecciation
and the presence of late white veins of quartz, rhodochrosite,
smectite, and relic santaclaraite. D, B26. Alternating layers of
caryopilite (dark gray) and rhodochrosite. Layers are still coherent,
even though much of the carbonate has recrystallized to spherules.
E, BS4. Coherent interlaminations of chlorite, rhodochrosite, and
caryopilite (top). The thick caryopilite layer (bottom) is internally
brecciated, making it impossible to follow individual bedding lamina-
tions across the hand specimen. F, B81—1. Disrupted braunite—rich

 

layers (dark gray) in carbonate-rich material (light gray). The
slightly displaced breccia fragments can be mentally and visually
rearranged, permitting some layers and laminations to be traced
across the specimen. G, B53. Layers of braunite (black), hausmannite
(grayish red), and rhodochrosite (grayish pink); texture is intermediate
between disrupted and dismembered. Upper and lower surfaces of
individual layers and laminations are indistinct. Displacement of brec-
cia fragments is so great that reconstruction into continuous layers is
not certain, but compositional contrast clearly indicates orientation of
original bedding. H, 65H86. Dismembered layers of gageite (black) and
rhodochrosite (grayish pink) in hausmannite and rhodochrosite (grayish
red). Layer contacts are indistinct; reconstruction of the layers is not
possible. Compositional contrast still suggests orientation of bedding.
1, B58. Gray rhodochrosite protolith, with slightly grumous texture,
and gageite (black). We interpret this specimen originally to have been
interlayered rhodochrosite and gageite. Disruption has been so intense
that the orientation of original layering is not evident, giving it a
massive nature. J, B78. Massive micromelange of rhodochrosite (gray-
ish pink), hausmannite disseminated in rhodochrosite, caryopilite, and
braunite (black). Neither individual layers nor the orientation of
bedding is clear. Scale in centimeters.

24

MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

 

FIGURE 4. —Continued.

 

chert does not occur within the ore lens. Massive chert
contains manganiferous laminations (1 to 5 mm thick)
adjacent to the orebody, but the laminations are nonman-
ganiferous away from the contact. The distinction
between laminae is based on impurities. The least pure
chert laminations are made cloudy by opaque “dust”
(presumably oxides and hydroxides) and by smectite.
Impure layers are characteristically fossiliferous, have
ﬁne—grained quartz, and, adjacent to the orebody, may
contain bundles of manganese-silicate ﬁbers. The pure
layers are devoid of dust and smectite, do not preserve
fossils, are composed of more coarsely crystallized
quartz, and may contain irregularly shaped or rhombo—
hedral bodies of carbonate. In the transitional interval of
at least several centimeters surrounding the north ore-
body, the chert contains thin laminations 0f manganifer-
ous sheet silicates, which cause partings, and some of the
chert contains santaclaraite (ﬁgs. 5C,D). Chert lamina-
tions are coherent to disturbed; soft—sediment deforma—
tion has not been extensive. Impure laminae are
transected by irregularly shaped bleached zones of well-
crystallized pure chert; the boundaries between impure
laminae and bleached zones are gradational (ﬁg. 20).
These bleached zones mark the pathways by which fluids
escaped from the soft sediment. Rarely, laminations
have been disrupted perpendicular to layering (ﬁg. 5D),
as if ﬂuid had pushed through the laminations. This
toe-like texture was probably formed by escaping ﬂuids
during compaction and diagenesis. Veins of coarsely
crystallized quartz have sharp contacts, rarely contain
bundles of manganese silicate ﬁbers, and were clearly
formed after the siliceous sediment was indurated.
Although there is evidence that solutions removed man-
ganiferous components from the chert (ﬁgs. 20, 5C),
there is no indication that the manganiferous laminations
were introduced by vein-forming processes (ﬁg. 50).
Within the chert, these manganiferous laminations
appear to be primary sedimentary features.

The average chemical compositions of chert and met-
ashale in the Ladd—Buckeye district (Hein and others,
1987) are compared with the compositions of the manga-
niferous protoliths in table 6. The major distinctions are
that chert has a smaller Mn/Fe and, in terms of absolute
concentrations, more Na and K. There is less Mn, Mg,
Ca, 002, Ni, V, and Zn in the chert than in the seven
manganiferous samples analyzed, but, because the large
amount of silica tends to dilute the concentration of all
other constituents, the smaller values are not unex-
pected. Clearly, the chert is not mineralized. The REE
patterns of the chert and Mn sediments are similar (see
ﬁg. 7). In general, the Ce anomaly (relative to seawater)
is distinct in chert.

The analyses of metashale are dominated by a ﬁne
clastic sediment component (represented by 0.875

PROTOLITHS 25

weight percent Ti02). Copper and nickel concentrations
exceed the small amounts expected of graywackes (Tay-
lor and McLennon, 1985) and could be due to Mn—poor
mineralization. The REE patterns of the metashales are
similar to, but more enriched than, those of the gray-
Wackes (ﬁg. 7A). Ce anomalies are prominent in both.
This similarity suggests that the pattern is being domi—
nated by the ﬁnest fraction of the turbidites; any contri-
bution to the REE pattern that resembles the patterns of
Buckeye ores (ﬁg. 7A) is obliterated by the presence of
the elastic material.

Hein and others (1987) and Hein and Koski (1987)
argued that the Buckeye orebody formed by the replace-
ment of chert by rhodochrosite and suggested that the
silica released by this process contributed to the forma-
tion of the massive cherts. It is important to realize that
many of the laminations in the massive chert contain relic
skeletal debris and that some of these laminations pre-
serve frameworks supported by ﬂattened radiolarians.
We conclude that most or all of the laminations were
originally composed of the debris of pelagic organisms
with variable amounts of smectite and oxide-hydroxide
“dust.” Thus, a direct major ﬂux of silica from the site of
rhodochrosite deposition to the site of massive chert
accumulation is precluded, making it unlikely that mas-
sive chert could form as a direct consequence of ore
formation. Further, if an orebody with the size and
manganese concentration of the Buckeye were formed by
replacement, one might ﬁnd a halo of replacement fea-
tures about the orebody. There are no such features. The
manganese components that occur in massive chert
appear to have been deposited With the massive chert.

RHODOCHROSITE PROTOLITH

Manganese carbonate appears to be the most abundant
mineral at the Buckeye deposit, but this abundance may
be in part an artifact of the selective mining of higher
grade oxide-rich (hausmannite and braunite) ore (Hueb-
ner, 1967). Carbonate protolith is best preserved in
samples B42, B50, B58 (ﬁg. 41), B103, and B107; it now
consists of extremely ﬁne-grained, medium—gray rho-
dochrosite that is so turbid in thin section that the
characteristic optical properties of carbonate cannot be
recognized (ﬁg. 5A). The rhodochrosite protolith can be
identiﬁed by the (012), (104), and (018, 116) reﬂections at
242°, 31.4°, and 518° in XPD patterns made with copper
radiation. The (104) of rhodochrosite is so strong that it
is a sensitive indicator of as little as 2 percent carbonate
in mixtures. However, the (012) reﬂection overlaps with
the (002) of caryopilite. Partial recrystallization to
coarser, lighter gray rhodochrosite results in a grumous
texture with distinct optical properties in thin section.
More advanced recrystallization results ﬁrst in light-

 

gray carbonate and loss of the patchy texture (as in
samples B17, B28, B108) and ultimately in pink rhodo—
chrosite (B64, B69, B73, B74, and B98). Small segrega-
tions of clear yellow, sensibly isotropic gel-like material
(B50, ﬁgs. 6G,H) occur within the gray carbonate. The
carbonate forms a breccia of equant to tabular bodies, 0.5
to 3 cm in size and with irregular to serrated margins,
that are separated by selvages of gageite (B45, ﬁg. 5A;
B50, B58, B108). We interpret the ﬁnest grained dark-
gray carbonate to be relic primary carbonate. The yellow
bodies are relic gel-like materials, residues of a siliceous
component that presumably once was more uniformly
distributed throughout the carbonate. Small amounts of
smectite similarly occur in several samples. The gel-like
material may have imparted a sufﬁciently cohesive qual—
ity to the ﬁne—grained carbonate (micrite) that it frac-
tured into equant bodies, about Which gageite ﬂowed. In
other samples, the micrite ﬂowed about broken layers of
caryopilite (B54) and of braunite and gel-like material
(B71b, ﬁgs. 6A,B).

The chemistry of the rhodochrosite protolith, repre-
sented by B50 (table 6), illustrates the enrichment of
manganese relative to iron and titanium (and most other
major rock-forming oxides). The mineralogy consists of
rhodochrosite, gageite, and devitriﬁed or partly devitri-
ﬁed gel-like material of chlorite composition and is con-
sistent with the bulk chemistry. The 002 concentration
is 23.6 percent by weight, compared with 38 percent for
pure rhodochrosite. Al and Mg are associated with
yellow, partially devitriﬁed gel-like material of chlorite
composition, and silica is contained in gageite and the
gel. Zn, Mo, Ni, As, Ti, B, and W are the only trace
elements that are present in concentration greater than
25 ppm. Compared with other protoliths from the Buck-
eye deposit, rhodochrosite has low concentrations of
REE (ﬁg. 7B). The Ce anomaly is not well deﬁned,
perhaps due to the fact that Nd values for carbonate
were below the detection limit. However, Sm values are
higher than would be expected if the Ce anomalies were
prominent. In comparison with many manganiferous
marine sediments, particularly most ferromanganese
nodules and pavements, the carbonate and other proto-
liths from the Buckeye deposit contain small concentra-
tions of Co, Ni, Cu, and Ba.

CARYOPILITE PROTOLITH

Caryopilite is found in claylike masses (B38; B84, ﬁg.
4E), weakly foliated layers (B20, ﬁg. 58; B25, ﬁg. 4C),
and thin, well-foliated layers (B26, B54) and fragments of
layers (B24, B83). The most poorly crystallized caryopil-
ite is claylike (or rarely gel-like) and represents de—
watered protolith. Caryopilite is identiﬁed by its (001),
(002), (20—1, 130), and (20—2, 131) X-ray reﬂections at

 

A

FIGURE 5.—-Characteristic microscopic features of protoliths in the
Buckeye deposit, California. A, Sample B45. Serrated septa of
dark-red to opaque (black in photograph) gageite surround irregular
fragments of extremely ﬁne grained rhodochrosite mud. This texture
is presumably caused by soft-sediment deformation of interlaminated
carbonate and gageite, not by late-stage hydrothermal veining. B,
B20. Interlaminated caryopilite (c) and taneyamalite (t). Taneyama-
lite is squeezed from primary layer and intrudes caryopilite layers.
Note presence of transparent caryopilite, designated cg. C, 65H76.
Brecciated santaclaraite—bearing chert cut by quartz veins. The

120°, 243°, 320°, and 358° CuKa, respectively, in
accordance with the indexing by Guggenheim and others
(1982). We also observed broad reﬂections at 56.5° and
580°. In thin section, the relatively high birefringence of
claylike caryopilite distinguishes it from chlorite. Evi-
dence that some caryopilite may have originated as
gel-like material is the occurrence of transparent layers
of clay-sized particles that may represent devitrifying gel
(B20, ﬁg. 5B) and one occurrence of isotropic gel-like
material with caryopilite composition (B91, ﬁg. 8D). In
claylike mixtures, the presence of caryopilite is best
indicated by its X-ray pattern or composition. The cary-
opilite protolith commonly occurs as 0.5- to 3-mm layers
interlaminated with carbonate or, less commonly, with
taneyamalite or chlorite. Ghostlike outlines of spheroidal
skeletal debris are present, and transparent orange—
brown ovoids (skeletal casts of radiolaria?) are rare
(B38). Dismembered caryopilite layers are engulfed by
carbonate (B24), caryopilite-carbonate mixtures (B39),

 

 

. i. , y. u».- 7 ,1
f‘ 'a. u
' _.C , .
m4; .«z' "9&9 new

B

abundance of santaclaraite is controlled by bedding laminations, and
this mineral clearly was not introduced by the later, crosscutting vein
ﬂuids. D, 185910. Dewatering texture in laminated chert. Dark layers
are manganiferous sheet silicates (top); discrete crystals are san-
taclaraite. Fluids moved perpendicular to layering, cutting across
earlier manganiferous laminations. Compare with similar texture in
braunite, in E. E, BS5. Dewatering texture in braunite, similar to
dewatering texture in chert (compare with D). Reﬂected light. F, B95.
Thin laminations with braunite composition. Note also the coarse
euhedra of braunite.

or taneyamalite (B20, ﬁg. 53; B29) or form spectacular
intraformational microbreccias (B83; B84, ﬁg. 4E).

The composition of the caryopilite protolith is repre-
sented by B84 (ﬁg. 4E). The analyzed split consists of
caryopilite with minor rhodochrosite and chlorite or
smectite. Again, the mineralogy is consistent with the
chemistry (table 6): Al and Mg reﬂect the chlorite/smec-
tite and 002 the carbonate. The FeO content is greater
in B84 than in the carbonate protolith (B50) and reﬂects
the presence of caryopilite. The value for M0, 740 ppm, is
unusually high and is due to intimate intergrowths of
ﬁne-grained (less than 5 mm) ﬂakes of molybdenite and
caryopilite. This intergrth has a patchy distribution
but is not obviously related to the late veins that cut the
sample (ﬁg. 4E). The increased Ba value, relative to B50,
is probably due to postdepositional replacement by bar-
ite. Although barite has not been found in thin sections of
B84, it occurs as veins in B29 (caryopilite), B31 (dolomite
near contact between Grummett and Sulphur Gulch bro—

PROTOLITHS 27

 

 

 

28 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

FIGURE 6.—Gel-like and partly recrystallized materials from the
Buckeye deposit, California. A, B71b. Braunite (top) and fragments
of gel-rich layers (bottom) displaced by fractures and engulfed in
rhodochrosite. B, Enlargement of previous photograph. The frag-
ments consist of clear, colorless, isotropic gel-like material (white;
table 8, analysis 6) interlaminated with braunite (black) and rhodo-
chrosite (gray). Mottling of the carbonate matrix is caused by
admixed yellow gel-like material. Both gel-like materials have chlo-
rite composition. C, B84. Gel-like laminations (ﬁg. 4E, top) having
chlorite composition (white) interlayered with caryopilite (brown)
and carbonate (brownish gray). Plane-polarized light. D, Detail
under crossed nicols. The margins of the gel-like layer have been
replaced by caryopilite and the interior of each layer has recrystal-
lized to chlorite. E, B24. Fragment of pinkish (yellowish in photo-

 

graph), transparent, gel-like layer (table 8, analysis 4), now surrounded
by carbonate and taneyamalite. Plane-polarized light. F, Detail,
crossed nicols. The gel-like material is slightly recrystallized to chlorite
and contains tiny rhombohedra of rhodochrosite. G, B50. Residual
pale-yellow gel-like material of chlorite composition (table 8, analysis 3)
surrounded by gageite and rhodochrosite. Plane-polarized light. H,
Same, crossed nicols. Gel-like material is isotropic. I , B80—4. Disrupted
braunite layer with internal, curved (conchoidal), branching fractures.
This shattered texture can be more easily explained by a process such
as shrinkage or shock than by pervasive (external) shear deformation.
Reﬂected light. J, B91. Rosettes of pale-yellow gageite growing
against isotropic gel-like material of caryopilite composition (table 8,
analysis 1) that contains crystals of braunite and rhodochrosite.

PROTOLITHS 29

 

 

30 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

 

 
     

10,000 EXPLANATION

O Metashale

|:l Graywacke
<1 Chert

||ll||||

 

Sample / (Seawater x 1,000)

 

La Ce Pr Nd

Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

A

FIGURE 7.—Rare-earth-element (REE) patterns of country rock and
ore from the Buckeye deposit, California, normalized to concentra-
tions for seawater (Hogdahl and others, 1968). A, 4 metashales, the
range for 7 graywackes, and 12 cherts. From Hein and others (1987,
p. 218—219, Ladd-Buckeye district) and unpublished data for the
Buckeye deposit by Huebner and Flohr. B, Mn-rich lithologies from
the Buckeye deposit. Sample B104 is hausmannite rich, B84 is

ken formations), and B64 (quartz vein cutting pink
recrystallized rhodochrosite; table 6). The REE abun-
dances are enriched relative to seawater and other
protoliths (ﬁg. 7B). The Ce/La is 1.9, and the Ge anomaly
(relative to seawater) is very distinct.

TANEYAMALITE

Taneyamalite occurs as rare monomineralic layers
conformable with laminations of caryopilite (B20, ﬁg.
53), as two laminations interlayered with chlorite and
carbonate (B38, ﬁg. 8A), and as veins crossing bedding
planes (B24; B29; B20, ﬁg. 53). We interpret the vein
material as a soft sediment of taneyamalite composition
that ﬂowed during compaction from bedding laminations
into veins, although taneyamalite is so uncommon that
we are not absolutely certain that it represents a sedi-
ment composition. A resinous golden color in hand spec-
imen, yellow color in thin section, and relatively high
birefringence are indicators of the possible presence of
taneyamalite. The XPD pattern is characteristic: the
(010), (—110), (-130), (—231), (032), (023), and (014,
4—23) reﬂections occur at 9.5°, 11.0°, 272°, 320°, 334°,
340°, and 406°, respectively. Insufﬁcient pure material
is available for a bulk analysis. Microprobe analyses
indicate that taneyamalite is the most sodic manganifer-
ous phase at the Buckeye.

 

 

10,000

l lllllll
|l|||lll

       
 

|
I

1,000 H

llllllll

100

Sample/(Seawater x 1,000)

 

 

10lll||ll|lll||
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

 

B

caryopilite rich, B112 is caryopilite and carbonate, B95 is braunite rich,
B64 is a vein of pink recrystallized rhodochrosite and quartz, 65H79 is
gageite, and B50 is unrecrystallized gray carbonate with gageite. The
chemical analyses and mineral modes for these seven samples are given
in table 6 and plotted in ﬁgure 9D. The range in REE contents of 17
additional Mn-rich protoliths (Huebner and Flohr, unpub. data) is
comparable to that of the 7 samples plotted in this ﬁgure.

CHLORITE PROTOLITH

The chlorite protolith is represented by two textures:
gel-like material, which occurs as thin laminae (B84, ﬁgs.
4E, 6C,D) or fragments of gel layers (B24, fig. 6E,F;
B57; B71b, ﬁgs. 6A,B), and massive layers of semi-
oriented, clay-sized crystallites having, collectively, 10W
birefringence (B11, B24, B25, B26, B31, B81a, 65H70).
Texturally, gel-like material of chlorite composition dis-

FIGURE 8. —Minerals and textures of rocks from the Buckeye deposit,
California, discussed in text. A, B38. Interlaminated taneya-
malite—phase A mixture (dark), chlorite (micaceous), and rhodochro-
site (light gray) surrounded by caryopilite. B, B80. Mass of recrys-
tallized gel-like material (dark, near center) in the caryopilite-
chlorite mixture (mottled, medium gray). Note the rim of
transparent caryopilite (white) adjacent to braunite layer (dark). C,
B24. Coarsely crystallized caryopilite, relic gel-like material of
chlorite composition, and a mixture of caryopilite, wispy chlorite, and
relic gel-like material. The chlorite-rich composition of the gel—like
material may be due to removal of Mn by crystallization of caryopilite
from a gel-like material having initial chlorite-caryopilite composi-
tion. D, B91. Colorless to pale-yellow gageite rosettes (white to gray
in photograph; table 9, analysis 2) in isotropic gel-like material of
caryopilite composition (table 8, analysis 1) with disseminated rho-
dochrosite rhombs. Crossed nicols; ﬁeld of view is part of ﬁgure 6J.
E, B24. Bundles of radiating yellowish-orange ﬁbers of phase A
(table 9, analysis 10) in recrystallized rhodochrosite (lightest gray,
table 1], analysis 5) and chlorite (ch) (pale brown, table 8, analysis 5).
F, B6. Layers of braunite (massive) and of hausmannite idioblasts in
carbonate.

PROTOLITHS 31

 

32 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

plays stages of “devitriﬁcation” that range from sensibly
isotropic, through partially crystallized (wispy ﬂakes of
chlorite in residual gel-like material), to fully crystallized
material with the characteristic anomalous low birefrin-
gence of chlorite. Commonly, gel-like material and
recrystallized gel-like material of chlorite composition
are clear in transmitted light, whereas claylike masses
of chlorite composition are cloudy. The occurrence of the
(001) X-ray reﬂection at 14.3° is diagnostic of the pres-
ence of chlorite. Mechanically, the gel-like material
appears to have fractured, forming curved surfaces (B24,
B71b), rather than to have ﬂowed to surround fragments
of other material. The chlorite laminations are too thin to
be sampled for bulk chemistry.

GAGEITE

Gageite forms brownish-black laminations and layers,
0.5 mm to 1.5 cm thick, interlayered with hausmannite
(B106; 65H79, ﬁg. 4B), rhodochrosite, or mixtures of
rhodochrosite and hausmannite (B45; B48; 65H86, ﬁg.
4H). In thin section, the layers are opaque to dark red or
dark brown and massive (granularity is not visible);
orange to yellow anisotropic patches that appear ﬁbrous
are rare (65H79). X-ray diffraction reﬂections at 129°
(200, 020), 14.3° (120), 26° (400, 040), and 27.5° (330) and
a broad reﬂection at 328° are diagnostic. Gageite layers
are commonly monomineralic, but in B106 some gageite
laminations contain hausmannite idioblasts. Gel-like
material of gageite composition does not occur, although
some gageite is closely associated with residual gel-like
material of chlorite composition. Thus, the nature of the
gageite precursor is unknown. Gageite occurs both as
broken layers in dismembered ore (65H86, ﬁg. 4H) and
as veinlets, apparently originating within a primary
layer, intruding hausmannite protolith (65H79, ﬁg. 4B).
The mechanical behavior of the gageite protolith is
intermediate between that of a ﬂuid carbonate mud and
that of the tougher hausmannite layers.

A 1-cm thick layer of gageite (65H79) appears mono-
mineralic, but XPD results and the presence of 0.7
percent CO2 suggest contamination by surrounding haus-
mannite and carbonate, respectively (table 6). The chem-
istry of this sample is almost completely described by the
system MnO-SiOZ-HZO. Compared with other protoliths,
gageite has the highest concentration of tungsten. Tung-
sten was detected only in samples that contained gageite,
and the concentration of tungsten is roughly proportional
to the amount of gageite in the mode (table 6). Gageite is
also associated with relatively high values of As and Ni.
The gageite protolith has the most depleted light-REE
abundances of all sediments, but the enrichment of heavy
REE relative to light REE exceeds that of the other
protoliths (ﬁg. 7B). The Ce concentration is below the

 

detection limit, preventing detection of a Ce anomaly, if
present. Nevertheless, this REE pattern is most
unusual. The only similar pattern of which we are aware
was obtained from “Siliceous Mn-oxide ore” from the
Ladd-Buckeye district (Hein and others, 1987, p. 218).
Because Hein and others did not recognize gageite in any
of their samples, we cannot verify that this unusual REE
pattern is characteristic of gageite.

HAUSMANNITE PROTOLITH

Hausmannite forms well—developed crystals in thin
coherent layers (B55; B56; B87, ﬁg. 4A), in fractured
layers (B53, 65H172), and dismembered layers (65H86,
ﬁg. 4H) and is disseminated in rhodochrosite layers
(B106; B6, ﬁg. 8F). In addition to its occurrence in
sedimentary layers, hausmannite occurs as massive
oxide (B104; 65H79, ﬁg. 4B; 65H171) that is not obvi-
ously sedimentary and may represent the “hydrother—
mal” hausmannite of earlier workers. In all occurrences,
the deep-red color in transmitted light and the deep—red
internal reﬂections in reﬂected light, combined with the
characteristic equant shape, pitted surfaces, and distinct
anisotropy in reﬂected light, distinguish hausmannite
from opaque braunite, deep-red gageite, and oxidized
caryopilite. However, hausmannite from the Buckeye is
not as extensively twinned as is hausmannite from rocks
of higher metamorphic grade. The relative lack of twin-
ning may reﬂect subsidence and deformation while the
orebody was still a soft sediment; conversely, it may be
due to the common occurrence of Buckeye hausmannite
in easily deformed carbonate. Except when disseminated
in rhodochrosite, the hausmannite fractures rather than
ﬂows. Hausmannite is responsible for X-ray reﬂections
at 180°, 290°, 324°, 361°, 585°, and 59.9° (101, 112, 103,
211, 321, and 224 reﬂections, respectively), but its optical
properties are so distinctive that it was not necessary to
rely upon the characteristic X—ray reﬂections.

Massive hausmannite protolith (B104) containing car-
bonate and gageite impurities was analyzed (table 6).
Considering that the mineralogy is basically that of an
oxide, the FeO value of 0.19 percent is extraordinarily
low. Hausmannite has the highest Zn concentration of all
the protoliths analyzed, but no other trace-element con-
centration appears to be characteristic. The 30 ppm W
reﬂects the presence of gageite in the sample. The REE
abundances of B104 are enriched (ﬁg. 7B), but the
pattern is unusual. The Ce/La ratio is only 0.3, resulting
in a slightly negative Ce anomaly. The enrichment of the
heavy REE, relative to seawater, is greater than for any
of the other Mn-rich protoliths. The lack of a Ce anomaly
and the heavy—REE enrichment result in a pattern that
is almost ﬂat.

PROTOLITHS 33

BRAUNITE PROTOLITH

Braunite occurs as thin laminations (10 um) commonly
grouped into layers (B71b, ﬁg. 6A; B81, ﬁg. 4F; B95, ﬁg.
5F) that may be so closely spaced as to appear massive
(65H19, 65H170). Opaque braunite layers range from
poorly crystalline layers (lacking metallic reﬂectance but
giving the characteristic braunite X—ray reﬂections) to
coarsely crystalline mosaics of anhedral crystals that
exhibit the characteristic braunite anisotropy in shades
of brownish gray. X-ray reﬂections at 33.0°, 38.1°, and
553° are characteristic. We found no braunite 11 (see De
Villiers, 1980). Braunite is interlaminated with rhodo—
chrosite, caryopilite, gageite, and, rarely, hausmannite.
The coarsest crystals, 0.2 to 0.5 mm in size, invariably
grow against carbonate at the margins of laminations. In
their extent of deformation, the layers range from coher-
ent through broken (B6; B81, ﬁg. 4F) and dismembered
(B53, B55) to a micromelange in which laminated braun-
ite layers are engulfed in carbonate and give no indica-
tion of their original spatial relationships (B59). Braunite
layers fractured into fragments, but protolith braunite
was never observed to have ﬂowed to surround frag-
ments of another protolith. On a smaller scale, braunite
layers contain many fractures that originate within,
rather than pass completely through, the layers. Some
braunite layers shattered such that the individual frag-
ments are slightly separated but did not change their
relative positions (B80, ﬁg. 61). This behavior suggests
that braunite had gel-like mechanical properties. An
unusual texture in which “toes” of coarsely crystallized
braunite protrude into a mixture of rhodochrosite, chlo-
rite, caryopilite, and ﬁne-grained braunite (B95, ﬁg. 5E)
probably reﬂects the loss of ﬂuids by passage through a
braunite layer. This texture is similar to the dewatering
texture in chert (fig. 50) mentioned previously.

A braunite-rich protolith with carbonate, caryopilite,
gageite, and chlorite as impurities was analyzed chemi—
cally (B95, table 6). The only major constituents (greater
than 2 percent) are Mn, SiOz, COZ, H20, and Mg. This
protolith has relatively high Co and Ni (26 and 840 ppm,
respectively), but no other trace-element concentrations
are characteristic. The 22 ppm W reﬂects admixed
gageite. The REE abundances are higher than those in
the carbonate (B50) and gageite (65H79) protoliths but
lower than those in the caryopilite, graywacke, and
metashales (ﬁg. 7B). The Ce/La is 4.4 and the Ce
anomaly is very well developed. The contrast between
the patterns for the hausmannite-rich and braunite-rich
protoliths is striking. Because Ce has both trivalent and
quadrivalent oxidation states, whereas La is exclusively
trivalent, the variation in Ce/La may indicate relative
redox states during deposition of the protoliths, with the
hausmannite protolith being deposited under more

 

reducing conditions than the braunite protolith. This
interpretation is consistent with the Mn+2/Mn+3 in these
oxides.

DIOPSIDIC ACMITE PROTOLITH

In addition to the chert and metashale, a nonmangan-
iferous bed composed of extremely ﬁne grained diopsidic
acmite (B94, B96) occurs adjacent to braunite in the
exploration adit that lies at approximately the same
horizon as the south Buckeye orebody.

SUMMARY OF PROTOLITHS

Most Mn-rich samples from the Buckeye deposit con-
tain relic layering, suggesting that the orebody was
originally layered rather than massive, composed of
interlaminated and interlayered sediments that are now
represented by eight minerals. We believe that the
surfaces bounding the laminations represent the original
bedding planes, parallel to the sediment—seawater inter-
face. While in a soft-sediment state, the more rigid,
gel-like layers broke or fractured and the less competent
muds ﬂowed, resulting in internal structures that range
from layered and laminated to massive. This range of
structures is probably responsible for the conﬂicting
descriptions of the ore by our predecessors. By this
reasoning, their massive ore corresponds to our most
intensely dismembered material. Because there is no
clear evidence of bioturbation, the disruption and dis-
membering of layers probably is due to compaction or
tectonism.

Organisms had a passive role in providing a detrital
component (radiolaria). Direct evidence for either an
active or passive role in precipitation of manganese is
lacking, even though marine bacteria are known to
oxidize and reduce manganese (Nealson, 1983). At a scale
of 0.1 to 1 pm, we observed no oval structures that might
be indicative of bacterial cell walls (Ghiorse and Hirsch,
1979) and no casts of bacterial or algal ﬁlaments
encrusted with oxides (Alt, 1988) or sulﬁdes and oxides
(J onasson and Walker, 1987). In contrast, Ostwald (1981)
found fossil algal oncolites and probable coccoid bacteria
in Cretaceous shallow-marine manganese ores from
Groote Eylandt, Australia. On the basis of this ﬁnding
and the existence of microorganisms that precipitate
manganese oxides, he concluded that 10- to SO-um—thick
laminations in manganiferous metashales have “organo-
sedimentary structure” (p. 561—562). A recent report of
delicate orange bacterial mats near low-temperature
vents (Malahoff and others, 1987) is tantalizing, in part
because the existence of such mats would be consistent
with the lack of bioturbation, but we have no evidence to

34 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

prove that the laminations in the protoliths at the
Buckeye deposit are of organic origin.

Replacement textures such as islands of relic Mn—poor
sediment in Mn—rich sediment, which would indicate
introduction of manganese after deposition of sediment,
were neither observed by us nor reported by earlier
workers. Indeed, it is hard to imagine that the extremely
ﬁne-grained gel-like materials could survive manganese
metasomatism without recrystallizing or that a lami-
nated nonmanganiferous sediment could be replaced by
the observed Mn-rich compositions while at the same
time preserving the intricate laminations and developing
the striking compositional contrast between layers.
Thus, the manganese-rich nature of the orebody and the
compositional contrast between laminations appear to be
primary features, a conclusion similar to that reached
years ago by Taliaferro and Hudson (1943, p. 272). All
ﬁbrous and sparry rhodochrosite, braunite, hausman-
nite, and ﬁbrous brown Mn-silicates observed by us are
recrystallized primary sediments. These minerals are
not diagnostic of postsedimentary hydrothermal or
supergene activity, a view promulgated by Hewett
(1972).

The Buckeye orebody is extraordinarily manganese
rich; most protoliths can be described by the system
MnO—SiOz—OZ—COZ—HZO. Trace metals occur at much
lower bulk concentrations than in many other marine
manganese deposits, but the range of trace-element
concentrations among protoliths is so large that the
maximum values of some trace elements (Sc, Cr, V, Mo,
Ba, W) equal or exceed average values (Cronan, 1977) for
some trace-element-rich ferromanganese crusts and nod-
ules. The bulk chemistry of the orebody reﬂects pro-
cesses that separated manganese from iron and sulfur
with extreme efﬁciency. The complementary iron does
not occur within the orebody or the surrounding enve-
lope of massive chert. The metashales have high Fe/Mn
and may contain this iron complement, but we have not
determined whether the metashales reﬂect an original
detrital composition, contain a hydrothermal component,
or are a residue (manganese was leached from the
protolith of the metashale). The bed of diopsidic acmite is
evidence that at least some iron-rich sediment was
deposited adjacent to a manganese—rich sediment. No
concentrations of sulfur (pyrite or other sulﬁde) were
found, and the protoliths contain only trace quantities of
iron and copper sulﬁdes.

MINERAL CHARACTERIZATION AND
CHEMISTRY

Representative microprobe analyses are given in
tables 8 to 12. These analyses and additional analyses are

 

plotted in ﬁgure 9. Each plotted point represents either
an average of three or more analyses of homogeneous
phases (hausmannite, braunite, santaclaraite, parsetten-
site(?), talc, pennantite, clinopyroxene, smectite, phases
A, B, C, D, and most rhodochrosite) or an individual
spot—analysis of a heterogeneous phase (gageite, chlorite,
caryopilite, some phase A) or mixture of phases (chlorite-
caryopilite, taneyamalite—phase A). (The nominal com-
positions of most of these phases are listed in table 1;
phases A—D have not yet been identiﬁed as known
mineral species and are discussed later.) Minerals
present in thick laminations were initially identiﬁed
optically and were conﬁrmed by XPD. These identiﬁca-
tions formed the basis for the initial selection of materials
for microprobe analysis. Identiﬁcation of layers that
were too thin for routine sampling and XPD was prob-
lematic. Microprobe analyses of these laminations dem-
onstrated that we could not rely upon optical techniques
alone to identify the ﬁne-grained phases. Thus, micro-
probe chemical analysis became an essential tool in
routine mineral identiﬁcation.

Manganiferous materials display textures ranging
from gel-like through partially recrystallized gel-like
material to crystalline material; from subcrystalline to
granoblastic; and from massive and claylike to well
foliated. Although this variety of textures is shown in the
photographs, it is important to recognize that the diver-
sity is largely due to overprints. The original sediments
were gels, or gel-like materials, clay, and micrite.

GEL-LIKE MATERIALS

Our least recrystallized gel-like materials are color-
less, pale yellow or pink, optically isotropic, and have
compositions of sheet silicates (chlorite, caryopilite). The
gel-like materials appear to be cohesive and to have
fractured, rather than ﬂowed, to form curved surfaces.
Gel-like materials of varied compositions are known to
form in other environments. Nahon and others (1982)
described a colorless to straw-colored trioctahedral Mn-
rich smectite “plasma”; this material may be similar to
our silicate gel-like materials. Hoffert and others (1978)
found gel-like inclusions of Mn-rich oxyhydroxides in Fe-
and Si-rich marine sediments. These inclusions appeared
to crystallize to rancieite and todorokite. Haggerty
(1987; oral commun., 1987) reported dredging a nearly
amorphous white mass with the consistency of petroleum
jelly from a seamount in the Mariana forearc. This
material contains Mg and Si but no Al and gave a very
weak X—ray pattern similar to that of allophane. Brett
and others (1987) reported the occurrence of gel with
serpentine composition from the Juan de Fuca Ridge.
Thus, there is precedent for the formation of silicate and
oxide gel-like materials in recent geologic environments.

MINERAL CHARACTERIZATION AND CHEMISTRY

35

TABLE 8. —Compositions of gel-like materials, in weight percent, from the Buckeye deposit, California Coast

Ranges, determined by electron microprobe analysis

[Total Fe and Mn reported as FeO and MnO, respectively. nd, not detected; na, not analyzed. H20 not determined, resulting in low oxide sums for
caryopilite, caryopilite-chlorite, and chlorite]

 

 

 

 

 

Caryopilite-
Caryopilite Chlorite Chlorite Braunite
Sample B91 B24 B50 B24 B24 B71b B80 B48 B80
Column 1 2 3 4 5 6 7 8 9
Si02 ............................. 37.2 33.8 29.5 31.1 36.2 36.4 34.7 33.9 10.4
A1203 ............................ .95 11.7 17.7 17.2 12.9 14.0 13.9 14.0 nd
Ti02 ............................. nd nd nd nd nd nd nd nd .36
V203 ............................ nd nd nd nd nd nd nd nd nd
MnO ............................. 47.8 23.5 15.6 11.8 11.4 4.20 5.55 4.38 78.4
MgO ............................. 2.90 16.5 23.8 24.0 23.6 28.8 31.0 33.3 nd
FeO ............................. .40 2.39 .83 1.53 3.01 .39 .20 .20 3.20
NiO ............................. .10 .69 .31 1.05 1.64 .33 .41 1.49 nd
ZnO ............................. .17 .31 nd .33 .77 .67 1.12 ha .11
C30 ............................. nd .05 nd .14 .19 .47 .11 nd .48
Na20 ............................ .04 .04 .04 .05 .08 .11 .13 nd nd
Sum ....................... 89.56 88.98 87.78 87.20 89.79 85.37 87.12 87.27 92.95
1. Partly devitriﬁed caryopilite with gageite laths (table 9, analysis 2; ﬁg. 8D); 4 points averaged.
2. Devitriﬁed yellow mass within rhodochrosite (ﬁg. 8C); 9 points averaged.
3. Isotropic gel-like material within rhodochrosite layer (ﬁg. 6G,H); 3 points averaged.
4. Pink, isotropic gel-like material; contains rhombs of rhodochrosite (ﬁg. 6E,F); 4 points averaged.
5. Partly devitriﬁed gel—like material with phase A (table 9, analysis 10); 3 points averaged.
6. Isotropic gel-like material with braunite (ﬁgs. 6A,B); 5 points averaged.

7. Partly devitriﬁed gel-like material with anomalous deep
caryopilite; 3 points averaged.

blue birefringence and a high Zn content within layer of

8. Partly devitriﬁed chlorite with high Mg and Ni found with gageite and rhodochrosite; 3 points averaged.
9. Layer of now-crystallized gel-like material (ﬁg. 61); 3 points averaged.

The preservation of gel-like materials from Jurassic or
Cretaceous time was initially surprising to us but
appears to be possible if metamorphism is at low tem-
peratures and under dry conditions.

Various textures of gel-like materials are shown in
ﬁgure 6, and representative compositions are given in
table 8. Most well-preserved gel-like materials have
chlorite composition with variable concentrations of mag—
nesium and minor elements (table 8, analyses 3—8); one
isotropic gel-like material has caryopilite composition
(table 8, analysis 1). A nearly isotropic fragment of very
pale yellow gel-like material, removed from a thin section
of B57, gave a strong X-ray powder diffraction pattern of
chlorite. Isotropic gel-like material of chlorite composi-
tion occurs as fragments (ﬁg. 6E, table 8, analysis 4; ﬁg.
63, table 8, analysis 6) and as residues within gageite and
rhodochrosite (ﬁgs. 6G,H, table 8, analysis 3). Note the
similarity between the texture of this residue and the
texture of the dismembered carbonate protolith (ﬁg. 5A).
Commonly, gel-like material having chlorite composition
is partly recrystallized, and sometimes it occurs With
gageite (table 8, analysis 8), caryopilite (ﬁg. 83, table 8,
analysis 7), or braunite (ﬁg. 63, table 8, analysis 6).
Totally isotropic material with a gageite composition has
not yet been found. However, the materials from which
gageite has crystallized have residues depleted in gageite

 

component (ﬁgs. 6G,H, table 8, analysis 3). In addition to
rare isotropic gel—like material of caryopilite composition
(ﬁg. 8D, table 8, analysis 1), disrupted layers and frag-
ments of layers with a partly devitriﬁed to wispy texture
yield compositions between those of caryopilite and
chlorite (ﬁg. 8C, table 8, analysis 2). A further indication
that some caryopilite originally was a gel-like material is
the occurrence of remarkably transparent yet anisotrop-
ic caryopilite layers in B20 (ﬁg. 53). Evidence that
braunite layers were originally deposited as gel-like
materials is the shattered texture due to curved and
branching fractures that are conﬁned to a braunite layer.
This texture is probably caused by partial dehydration
and shrinkage (ﬁg. 61, table 8, analysis 9) or by tectoni—
cally induced shock.

The compositions of individual silicate species are not
uniform, even within a single sample or layer, indicating
that equilibration between the silicates (and between
silicates, oxides, and carbonate, as discussed below) did
not occur. In part, the range of compositions is attributed
to very ﬁne scale mixing of two silicates, which cannot be
resolved with the microscope or the microprobe (ﬁg. 9);
the silicates also occur as monomineralic material within
the same sample. In some cases, the actual end-member
composition cannot be deﬁned, making difﬁcult the dis-
tinction between a single phase and a mixture.

PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

MICROBANDED MANGANESE FORMATIONS

36

 

 

 

 

 

 

 

 

 

 

 

 

.owme: 3:6: m ”368585: mix 3:25 ﬁts Swahdﬁﬁcaw .3 mﬁmq .m:
.ommﬁwiw 3:6: m 5.85 55:5 whim miaﬂuwm .3
.vwmﬁwiw 3:6: v ”whim oﬁiﬁggﬁ 323% .«o .853 .m:
ﬁwmwSZN 3:6: v ”3:233:50: «Ea < @923 @832 NH
.omwﬁgm 3:6: m ”wuzmﬁawgu 5:5 :BOHMEQE .3
damage?” 3:6: m wa .mw um Embxsw .: 2:58 BEoEooooi :Ears Emma wiaﬂvﬁ .323: Emtm .oH
.owwﬁwiw 9:6: m 6::th8 ﬁts ESSEBE 958 855253: .m
.oewﬁw: 350: m ”312:0 wE—Rambo Ewmbmou mo .5qu .w
.cwwwEBa 3:6: m ”am .33 8:339 5:5 63:8 :m "E: MEEE.“ Bzﬁoido 25390.0 .N.
.cwwﬁgn 3:6: .0. Mug—Eobdu €2.32 :32 .o
.omwﬁwiw 3:6: my R3 .wm memo waging BEQPCS MS 5932 wEER m: :omﬁmomﬁco MwﬁmoEoocoi ﬁts :93: BEQoEO .m
.cwmﬁwzw 3:6: m ”:93?qu we: B :35 .v
.3998; 3:6: m 5mm .mu J mmmhaﬁw .2 29:6 Sﬁoéoooof SE 36me .«o 3.323.: Eon.“ Mﬁuﬂoﬁ Ewan xoﬂm .m
.owmgoia 3:5: m ”Sm .9: M: wings .m 3.33 35:9?3 35.53% Eta E mﬁﬂ $23.8 3.582 .N
.cmeEZN 3E3 m $5.2 323% 8 $2.830 A
B mSm mph mﬁm mSm mﬁm E E E E E o o o o . . . . . . . . . . . . . . 9515‘
voooﬁ wmmﬁm oomxmw www.mm oomﬁm www.mm obod oowd wmod ombd wood mooé ooos Hoos Eoﬁ ....... . . . .83 :2qu
omo. omw. bmo. bmw. ovo. vNo. Bo. :o. o o o o o o o ........ . . . . . . . . . NZ
mww. vmo. o 3o. oNo. omo. woo. mNo. ms. o o o o o o 50
o o o o «No. So. o :o. No. ooo. moo. go. o o So. . . . . . . . . . . . . . . . .. :N
o o o mmo. o o o o moo. ms. woo. o mmo. woo. o I... I. ....:Z
ooo. mow. owvé momd bwmﬁ oNo.m own. mow. mmo. o3. wow. Hmo. go. So. bmo. . . . . . . . . . . . . . . . . .. mm
Nwﬁ. owvd momﬁ oooé So. o3. moNA ovow mom. mmmé Em. oNN. va. wmﬁ. Sum. . . . . . . . . . . . . . . . . . . ME
www.mw mﬁmo meN Hmmw Emma obmd wom.m moo. 3%.: HmoaV oinv mp: 38.: vowd $5.: ........... . . . . . . . :2
o ooo. o mﬁo. mow. o wHo. 3o. o o o o o o o >
o ooH. Nvo. who. moo. mmo. o o o o ooo. moo. o o o . . . . ..... . . . . . . . . . . a.
N5. m3; owm. mom. wmm. ooN. oNoA Em. 2:. moo. me. o o o o 2...: 12.22.53.
moo. o o o o o «mm. mvo. o o o o woo. go. o ................. >L<
hood MHNNH onNH NoHNH 9%.: 3%.: Sod PNMMH mﬁmd NoNAV momd ooo.N mboA mwoé mum; . . . . . . . . . . . . . . 3.1%
£5: 355:8 .5: 25?: “E: 25550 .«o hoe—:52
£3 3% a: $5 Ems «98 ﬁ.% 2.8 8% 3% 5% 3.3 8.8 5% 3% ..:::....:.:§m
E. 8.: N: w: mo. 3. S. 8. E E E B E E E z... 2. . ommz
mmﬁ 3. w: wo. mo. 2. mo. VN. oo. o: o: o: o: w: o: . . . . . . . . . . . . . . ONO
o: o: a: a: S. no. w: 3. E. Z. 3. ma. a: o: mN. O:N
w: o: w: «A . w: o: o: 63 mo. ow. wo. o: ov. NH. 6: ................. Omz
vo. No.m om.o Ed omd Q: mmﬁ wmﬁ om. SA oﬁw Nv. oN. oN. hm. . . . . . . . . . . . . . . . . Ooh
ooé Saw who Nogm wmé EA NH.» oNN mod Hm.» Hod E; no; woA owé . . . . . . . . . . . . . . . . .033
«no: ﬁsh oAm o.mm usm NNm ﬁmm Hog. «v.5» m4: «Ev ado 93 now mdo . . . . . . . . . . . . . . . . .052
E NV. 2 8. a: E 8. S. E E E E E E E ................ ”09
mv: om. ON. mm. WN. ON. “a: U: U: U: OH. we. U: W: “V: . . . . . . . . . . . . . . . . .Nop
:. Sim wbé EA om. om. o.mH o4; ob. 5V. wk. 6:: mo. 2. o: . . . . . . . . . . . . . . .momjw
93 5.3. mow oéw wow mom wow o.mw mﬁm mom vow o.mN wﬁm WNN bNN .............. . . .momm
:50qu Emma? 5 £355
3 E m: N: 2 S a m h w m w m N M :EEoO
hm Sm wmm «mm wmm «um Hmm wmm owm Sm vmm omm nmwm Sm omm 295w
maﬁaﬂoﬂﬁwm wuzwﬁmzegﬁ €me 44. 3....an 322:0 35%?50 36me

 

838$: 3: 933 3:95? $35 :5 .m :5 .nm .M :8 vwsbmg €me 3:253:42 «Em < $.23 .M :5 am .5. ENE—Ea om? Awmmo < 33: .me .Emo wEEoSO 63235 go: .a: 63933. .8: 65

$33.35: 33%.582: :383» on 39:58:: .mmgsﬂ Ego Egcbgb .ﬁmoaw: 38:35 33 23% 3:238 Kc wggtmoREoDI .m mama;

MINERAL CHARACTERIZATION AND CHEMISTRY

37

TABLE 10. —Compositions of vein minerals from the Buckeye deposit, California Coast Ranges, determined by electron microprobe analysis
[nd, not detected; nay not analyzed]

 

 

 

 

 

 

 

 

Talc Smectite Parsettensitelike phase Pyroxene Phase B Phase C Phase D
Sample B71 B7 B59 B29 B94 815 B15 core B15 rim Bl5 B17
Column 1 2 3 4 5 6 7 8 9 10
Oxides, in weight percent
SiO2 ................. 62.1 46.9 52.9 44.7 52.5 32.9 (32.1—33.7) 36.9 35.4 32.2 0.54
A1203 ................ .39 7.34 2.75 6.13 1.50 .51 (0.38—0.69) 24.4 16.2 3.92 .17
Ti02 ................. nd nd nd .46 .21 .77 (0.36—1.44) 2.39 .27 .47 nd
V203 ................ nd nd nd nd nd 34.0 (32.4—36.0) 6.83 16.5 31.8 22.8
MnO ................. 2.69 2.31 .71 31.6 2.42 1.57 (0.99—2.43) .31 .52 1.23 58.7
MgO ................. 29.1 27.1 26.6 2.10 2.22 1.62 (0.92—2.41) .08 .25 1.33 .65
FeO ................. .37 2.28 nd .16 24.1 1.06 (0.73—1.73) .17 .41 .65 .23
N10 ................. nd .09 1. 66 na nd nd -- .07 nd .15 .13
ZnO ................. nd nd 1.59 na nd .08 (nd-O. 15) .07 .11 .10 . 10
CaO ................. .13 1.52 .66 1.03 3.82 15.1 (12.7—16.9) 15.8 13.8 11.8 nd
BaO ................. na na nd .42 nd .07 (005—0. 17) . 10 . 10 .20 nd
NaZO ................ .09 .22 .17 .18 11.0 1.37 (0.81—1.74) .04 .03 .05 nd
K20 ................. na nd na 2.22 ha nd -- nd nd nd nd
Sum ........... 94.87 87.76 87.04 89.00 97.77 89.05 87.16 83.59 83.90 8551*
Number of cations and anions per formula unit ,
Si ................... 7.984 6.749 7.541 8.000 1.993 6.026 6.054 6.255 6.092 0.157
A1IV ................. .016 1.251 .459 0 .007 0 0 0 0 0
AlVI ................. .043 .007 .003 1.292 .060 .110 4.717 3.384 .874 .057
Ti ................... 0 0 0 .062 .006 .107 .294 .036 .067 0
V ................... 0 0 O 0 0 4.992 .897 2.333 4.829 5.352
Mn .................. .293 .281 .086 4.794 .078 .243 .043 .078 .197 14.567
Mg .................. 5.572 5.803 5.651 .560 .126 .442 .020 .066 .375 .284
Fe .................. .040 .274 0 .024 .020 .163 .023 .061 .103 .055
Ni ................... 0 .010 .190 0 0 0 .009 0 .023 .030
Zn .................. 0 0 .167 0 0 .011 .008 .014 .014 .033
Ca .................. .018 .238 .101 .197 .155 2.966 2.764 2.608 2.390 0
Ba .................. 0 0 0 .029 0 .005 .006 .007 .015 0
Na .................. .022 .061 .047 .062 .810 .488 .013 .010 .018 0
K ................... 0 0 0 .507 0 0 0 0 0 0
Cation sum ........... 13.988 14.674 14.245 15.527T 3.999H 15.553 14.848 14.852 14.997 20.925TH
Anions ............... 22 22 22 1 1 6 24 24 24 24 24

 

* Sum includes 2.19 weight percent A5203. l Cations normalized to 8 Si.

1. Talc vein; 3 points averaged.

” Sum includes 0.744 Fe”.

m Sum includes 0.390 As.

2,3. Smectite within quartz and rhodochrosite veins, respectively; 3 points averaged.

4. Lath within quartz vein; 3 points averaged.

5. Yellow-green, slightly pleochroic diopsidic acmite forming laths and ﬁbers within quartz vein. Calculation of ferric iron, assuming ideal stoichiometry, yields

Fe203 = 26.1, F‘eO = 0.64, oxide sum = 100.41; 3 points averaged.

6. Green laths and ﬁbers in quartz vein; range in parentheses; 9 points averaged.
7,8. Columnar grain in quartz vein, showing core-to-rim zonation, with pink to colorless pleochroism; 2 points averaged.
9. Vanadium-rich lath that appears to be compositionally intermediate between phase B and other phase C grains yet contains essentially no Na compared to phase

B; 3 points averaged.
10. Vanadate, orange-red; 12 points averaged.

GAGEITE

Gageite, the most manganese-rich silicate analyzed, is
most striking when it occurs as black to dark-red
monomineralic septa within massive rhodochrosite (ﬁgs.
41, 5A) and as layers and veins within massive hausman-
nite. The principal reason for believing that this gageite
is a sedimentary protolith, rather than a product of a
postsedimentary process, is the less obvious occurrence
of gageite in layers and laminations that are mixtures of

phases, suggesting that gageite is a sedimentary compo—
nent. In these mixtures, gageite occurs as pale-yellow to
nearly colorless mats of ﬁne crystals associated with
carbonate and chlorite (B50, ﬁg. 6G), as radiating laths
growing from a layer now represented by carbonate and
residual gel-like material of caryopilite composition (B91,
ﬁg. 8D), and intergrown with braunite in laminae (B95).
Gageite was analyzed in seven samples (represented in
table 9 by analyses 1—4). In comparison with the other
manganese-bearing silicates, particularly caryopilite and

38 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

TABLE 11. —Compositions of rhodochrosite from the Buckeye deposit,

California Coast Ranges, determined by electron microprobe analysis
[All carbonates also analyzed for Sr, Ba, Zn, Ni, and V, but these elements were not detected.
C02 calculated from stoichiometry, assuming C=1.000. nd, not detected]

 

 

 

 

 

 

 

 

Sample B45b BS3 B53 B6 B24 B53
Column 1 2 3 4 5 6
Oxides, in weight percent
MnO ........... 59.6 56.9 57.4 57.5 53.3 52.7
MgO ........... 1.17 2.19 1.21 1.51 .41 .74
CaO ............ .13 1.78 2.34 1.68 5.80 6.89
FeO ............ .11 .09 .08 .09 .46 .08
Subsum 61.01 60.96 61.03 60.78 59.97 60.41
CO2 ............ 38.64 39.10 38.88 38.90 39.00 39.21
Sum ....... 99.65 100.06 99.91 99.68 98.97 99.62
Number of cations per formula unit
Mn ............. 0.962 0.902 0.917 0.922 0.862 0.839
Mg ............. .033 .061 .034 .043 .012 .021
Ca ............. .003 .036 .047 .034 .119 .139
Fe ............. .002 .001 .001 .001 .007 .001
C .............. 1.000 1.000 1.000 1.000 1.000 1.000
Cation sum ..... 2.000 2.000 1.999 2.000 2.000 2.000

 

1. Coarsest grained rhodochrosite from rhodochrosite protolith (ﬁg. 5A); 4
points averaged.

2. Rim of coarse—grained rhodochrosite in contact with braunite (table 12,
analysis 1); 3 points averaged.

3. Core of same coarse-grained rhodochrosite; 3 points averaged.

4. Fine-grained rhodochrosite with hausmannite (table 12, analysis 5; ﬁg. 8F);
4 points averaged.

5. Coarse rhodochrosite with phase A (table 9, analysis 10; fig. SE); 3 points
averaged.

6. Rhodochrosite vein cutting braunite; 3 points averaged.

chlorite, the composition range of gageite is very small.
No consistent correlation of composition with color,
texture, or associated minerals was found. The greatest
range in composition was found in samples 65H79 and
65H86. The observed color differences may be due to
variable amounts of Mn in other than the +2 oxidation
state. There is a good inverse correlation between Mn
and Si + Mg (table 9, analyses 1—4). Small amounts of
nickel (up to 0.5 weight percent NiO) were detected in
most gageite.

Chemical formulas obtained for gageite from the Buck-
eye are inconsistent with those reported by Moore
(1969), Dunn (1979), and Ferraris and others (1987) for
gageite from other localities. The crystal structure of
gageite described by Moore (1969) consists of walls of
edge-sharing octahedra corner-linked to bundles of edge-
sharing octahedra. Silicate tetrahedra reside in pipes
or channels and support the framework structure by a
network of oxygen bonds. Moore proposed M42(Si12036)
[06(OH)48] as the formula, where M represents the
divalent cations Mn, Mg, and Zn, and M/Si = 3.50.
Ferraris and others (1987 ), using electron diffraction and
transmission electron microscopy, conﬁrmed the general

 

octahedral structure but found that the distribution of
the silicate tetrahedra within the channels was not
consistent with the crystal-chemical model of Moore.
They found two kinds of interlinked modules built by
chains of edge-sharing octahedra and proposed M4206
(OH)40(Si4012)4 with M/Si = 2.625 as the formula for
gageite. Dunn (1979) analyzed gageite from Franklin,
N.J., and proposed the empirical formula M4oSi15050
(OH)40 with M/Si = 2.67. In the analyses by Dunn, M/Si
ranges from 2.63 to 2.74. Signiﬁcant ZnO and MgO
(3.8—5.2 weight percent and 10.1—12.9 weight percent,
respectively) were also present, leading to the sugges-
tion that Zn and Mg are essential constituents. Unlike
gageite analyzed by Dunn (1979), gageite from the
Buckeye contains little or no Zn, and there is only minor
substitution of Mg for Mn, indicating that Zn and Mg are
not essential structural constituents. Most Buckeye
gageite has M/Si or M/(Si+Al) of about 2.49 to 2.55,
outside of the range of the analyses of Dunn (1979) and
consistent with the formula M5(Si,Al)2O(9_x)(OH)2x (ideal
M/Si = 2.50). Because of scatter in our data, some
analyses are consistent with the formula of Dunn (1979)
and others are inconsistent with any of the discussed
formulae. The best agreement with our empirically
derived formula (MzSi = 5:2) is obtained when there are
4.5 to 4.9 Mn cations per 9 anions; as Mn decreases from
4.5 to 4.1 cations, Si increases to about 2.2 and M/Si
decreases to about 2.11. In a few analyses with fewer
than 3.9 Mn cations, Si decreases with decreasing man-
ganese content. We do not know whether the trace
amounts of tungsten in gageite—bearing protoliths occur
in gageite or another phase.

Despite the extreme color variation—from almost
transparent and colorless, through deep red or brown, to
opaque in thin section—we found no relationship
between gageite composition and color. A possible expla-
nation relates to the intense absorption of the Mn+3 ion,
which dominates the Mn+2 ion in spectra of oxides. Thus,
even minute amounts of Mn+3 ion would color the
gageite, yet we would not be able to see any difference in
gageite composition by electron microprobe techniques.
The occurrence of gageite supports this explanation: the
centers of gageite septa (deformed protolith layers) tend
to be intensely colored, whereas gageite laths growing
into gel-like material or against carbonate tend to be
almost colorless. We initially thought that dark color in
gageite was due to supergene oxidation by ﬂuids that
followed gageite veins and did not alter the enclosing
rhodochrosite. However, we observed that pieces of
carbonate, left on the surface for 20 years, developed a
black oxide rind even in the dry environment of the
Diablo Range. Thus, in the absence of any other evidence
of supergene alteration, such as alteration of the carbon-
ate immediately adjacent to the gageite, we propose

MINERAL CHARACTERIZATION AND CHEMISTRY 39

TABLE 12. —Compositions of the oxides braunite (columns 1—/,) and hansmannite ( columns 5—8) from the Buckeye deposit, California Coast

Ranges, determined by electron microprobe analysis
[nd, not detected; 113, not analyzed. MnO, Mn203, FeO, Fe203, Mn”, Mn”, Fe”, Fe+2 calculated by assuming ideal stoichiometry of 8 cations per 12 anions for braunite and 3 cations per
4 anions for hausmanniteJ

 

Sample B53 B53 B6 65H170 B6 B106 65H79 65H170
Column 1 2 3 4 5 6 7 8

 

Oxides, in weight percent

 

 

SiO2 ................. 9.79 9.88 9.87 10.0 0.22 nd nd nd
Ti02 ................. nd nd nd nd .05 nd nd nd
Mn203 ............... 77.5 76.7 78.4 78.6 67.8 69.2 69.8 69.1
Fe203 ............... .23 .51 .47 .24 .480 0 .21 .40
A1203 ................ nd nd .27 nd nd nd nd nd
Cr203 ................ nd nd na na na na na na
V203 ................ nd nd nd nd .05 .12 .08 nd
MnO ................. 10.9 10.6 11.1 11.6 0 30.3 29.0 30.6
FeO ................. 0 0 0 0 0 0 0 0
MgO ................. .05 .11 .15 nd .03 .36 1.33 .46
NiO ................. .14 .33 nd nd nd .08 .15 nd
ZnO ................. nd nd nd nd .07 .09 nd nd
COO ................. nd nd nd nd nd nd nd nd
CaO ................. .34 .40 .18 .15 .07 nd nd nd
BaO ................. nd .03 nd nd nd nd nd nd
Na20 ................ nd nd nd nd nd nd nd nd
Sum ........... 98.95 98. 56 100.44 100.59 99.77 100.15 100.57 100.56
“Mn0”* .............. 80.52 79.59 81.63 82.19 91.89 92.56 91.71 92.74
“FeO”* .............. .21 .46 .42 .22 .43 nd .19 .36

 

Number of cations per formula unit

 

 

 

Si ................... 0.995 1.006 0.985 1.000 0.008 0 0 0
Ti ................... 0 0 0 0 .001 0 0 0
Mn+3 ................ 5.992 5.947 5.961 5.984 1.967 1.996 1.991 1.982
Fe‘L3 ................ .018 .039 .035 .018 .014 0 .006 .011
A1 ................... 0 0 .032 0 0 0 0 0
Cr ................... 0 ‘ 0 0 0 0 0 0 0
V ................... 0 0 0 0 .002 .004 .002 0
Mn”'2 ................ .938 .917 .944 .980 1.000 .973 .922 .978
Fe+2 ................ 0 0 0 0 0 0 0 0
Mg .................. .008 .017 .022 0 .002 .020 .074 .026
Ni ................... .011 .027 0 0 O .002 .005 0
Zn .................. 0 0 0 0 .002 .003 0 0
C0 .................. 0 0 0 0 0 0 0 0
Ca .................. .037 .044 .019 .016 .003 0 0 0
Ba .................. 0 .001 0 0 0 0 0 0
Na .................. 0 0 0 0 0 0 0 0
Cation sum ..... 7.999 7.998 7.998 7.998 2.999 2.998 3.000 2.997
* Total Mn as MnO and total Fe as FeO, obtained by electron microprobe analysis.
1. Coarse-grained braunite rims in contact with rhodochrosite (table 11, analysis 2); 3 points averaged.
2. Fine-grained braunite with rhodochrosite in cores of grains (rim composition is given in this table, analysis 1); 3 points averaged.
3. Braunite coexisting with hausmannite (analysis 5, this table; ﬁg. 8F); 3 points averaged.
4. Braunite coexisting with hausmannite (analysis 8, this table); 6 points averaged.
5. Hausmannite from hausmannite-braunite-rhodochrosite sample; 3 points averaged. Braunite is analysis 3, this table.
6. Hausmannite from hausmannite-rhodochrosite sample; 6 points averaged.
7. Hausmannite layer within gageite; 3 points averaged.
8. Hausmannite coexisting with braunite (analysis 4, this table); 3 points averaged.
instead that the intense coloration of the gageite proto- CARYOPILITE AND CHLORITE
lith reﬂects its initial oxidation state and that the light-
colored marginal laths have inherited the (presumably) Caryopilite generally is orange or orange brown and

more reduced state of the adjacent carbonate and cary- contains about 0.1 weight percent each of NiO and ZnO;
opilite protoliths. the concentrations of V203 and TiO2 are generally below

40 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

Si

 

Rdc iln protlolithsl
2 Others

I

 

Si

 

D
It®1|1ll|

2 Others

 

 

2 Others

FIGURE 9. —Summary of chemical analyses of samples from the Buck-
eye deposit, California, plotted in the system Mn-Si-Others, Where
Others consists principally of the elements Mg, A1, Fe, V, Ti, Ni, Zn,
Na, and Ca. Cation proportions. A and B, microprobe analyses of
discrete phases; C, analyses of mixtures of phases, plus single-phase
vein carbonates; D, bulk analyses of representative protoliths. The
clustering of microprobe analyses does not represent the bulk
composition of the deposit. Bulk analyses of protoliths fall near Mn on

the detection limits of the microprobe (table 9, analyses
5—7). The iron concentration is commonly less than 3
weight percent FeO. The aluminum and magnesium
concentrations of material with the appearance of cary-
opilite range from those of near endmember caryopilite

2 Others

the Mn-Si join. Symbols: Ta, taneyamalite; Ca, caryopilite; Ga, gageite;
Br, braunite; Mn-chl, Mn-rich chlorite; Mg-chl, Mg-rich chlorite; Rdc,
Mn-rich carbonates; Pa, parsettensitelike mineral; Sa, santaclaraite;
A, B, C, and D, minerals designated phase A, B, C, and D; Tc, talc; Px,
acmitic diopside; Sm, smectite; Ti, titanite; Ha, hausmannite. In C,
squares represent analyses of phase A—taneyamalite mixtures, circles
represent analyses of caryopilite-chlorite mixtures, and crosses repre—
sent analyses of gageite-caryopilite mixtures.

to those of near endmember clinochlore (ﬁg. 9). It is
probable that small amounts of Mg and Fe can substitute
for Mn in caryopilite, and of Mn for Mg in chlorites,
within the two structure types, but the fact that Mg is
positively correlated with Al over much of the composi-

MINERAL CHARACTERIZATION AND CHEMISTRY 41

tion range indicates that most of the compositional
variation is due to mixing between these two phases.
When Mg exceeds about 1 cation per 14 anions (about 5
weight percent MgO) in caryopilite, the concentration of
Al increases linearly with the Mg, indicating that mix-
tures of caryopilite and chlorite have been analyzed. This
is the composition limit separating caryopilite from mix-
tures of caryopilite-chlorite in ﬁgure 9. Mixtures of
Mg-chlorite and caryopilite trend toward two chlorite
compositions, one with lower A1 and higher Mg than the
other (ﬁg. 10). A few points identiﬁed as mixtures (ﬁg.
10) with Mg less than 1 and about 0.5 A1 are from B38 and
are slightly more Fe-rich than caryopilite analyzed from
other samples. With the exception of sample B24, com-
positions within a given sample fall on only one trend. In
B24, two distinct chlorite compositions are found, as
discussed below. Rarely, caryopilite has high Mg (B17,
table 9, analysis 6; ﬁg. 10) without correspondingly high
A1 and low Si, suggesting that more than the arbitrary 5
weight percent MgO can substitute for Mn in the cary-
opilite structure (rather than being due to admixed
chlorite or pennantite). These high Mg values cause the
overlap between the caryopilite ﬁeld and the mixed
caryopilite-chlorite ﬁeld (compare ﬁgs. 93, C). Cary-
opilite also occurs with gageite. In addition to the
discrete caryopilite and gageite, sample B95 contains
deep-orange to deep-red material, suggestive of the color
of gageite but with intermediate composition (ﬁg. 9C)
attributed to intergrowths of caryopilite and gageite.
B17 is similar, but with only a trace of mixed caryopilite-
gageite. B54 contains discrete caryopilite and the mixed
material, but discrete gageite is absent.

Evaluation of microprobe analyses of caryopilite and
determination of its composition limits are made even
more difﬁcult because caryopilite has a modulated struc-
ture (Guggenheim and Eggleton, 1987), which results in
an excess of tetrahedrally coordinated cations and a
deﬁciency of octahedral cations relative to the ideal
serpentine formula unit, Mg68i4010(OH)8. Caryopilite
from the Buckeye deposits has octahedral/tetrahedral
cation ratios of 1.23 to 1.32, similar to but outside of the
range of 1.36 to 1.44 found by Guggenheim and others
(1982). The deﬁciency of octahedral cations in Buckeye
caryopilite may in part be due to the relatively oxidized
state of the orebody, causing substitution of Fe+3 or
Mn+3 for Fe” or Mn+2 (thereby creating vacancies), or
to admixed trioctahedral smectite.

An interesting textural relationship between caryopi-
lite and layers of mixed caryopilite and chlorite is
observed in sample B80 (ﬁg. BB). The sample is com-
posed of disrupted layers of braunite, caryopilite, and
chlorite. At the boundary between caryopilite-chlorite
and braunite layers, a rind of near-end-member caryopi-
lite (table 9, analysis 7) separates the mixed silicate layer

 

 

 

 

 

 

5.0 I l
4.5 — BBO\ o —
0
©0
4.0 — d) —
o
0 ga
O % —
u & o o
o
I]
o _
g mg]:
<1- IEDD _
Ba
2
ea
69
£9 a) _
o I I I I
0 0.5 1.0 1.5 2.0 2.5
AI/14 anions
EXPLANATION
O Chlorite
O Caryopilite
D Caryopilite-chlorite mixture
63 Mn-rich chlorite

FIGURE 10.—Compositions of caryopilite and chlorites from sample
B80 and 14 other samples from the Buckeye deposit, California.
There appear to be two chlorite-composition clusters and two mixing
lines toward caryopilite. Caryopilite analyses from sample B80
cluster at very Mg—poor compositions, chlorite analyses cluster at
very Mg—rich compositions, and caryopilite-chlorite mixtures span a
very large compositional range. The gaps between the endmember
caryopilite and chlorite compositions and the compositions of the
caryopilite-chlorite mixtures in sample B80 illustrate the lack of
equilibrium between caryopilite and chlorite as discussed in the text.

from the oxide layer. The compositions of the caryopilite
rind, mixed caryopilite-chlorite, and chlorite in B80 are
indicated in ﬁgure 10. The mixed layer in B80 has
compositions that span almost the entire range of the
high-Mg trend of the mixed caryopilite-chlorite analyses
from all samples; in B80, however, caryopilite and
chlorite are distinctly more Mg—poor and Mg-rich, respec-

42 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

tively, than the mixing line, illustrating the lack of
equilibrium between the silicate phases. Further, the
rind of caryopilite is not the result of equilibration
between the mixed silicate and oxide layers, as no
correlations between compositions of caryopilite and
braunite were found. The amount of substitution of other
cations for Mn in caryopilite from B80 is among the least
found in the Buckeye samples, and these compositions
(and those from B24, table 9, analysis 5) probably
represent the purest caryopilite in the Buckeye samples.

Most chlorite is the product of incipient recrystalliza-
tion of gel-like material. Rarely, chlorite forms claylike
layers (B24). Most chlorites are poorly crystallized, and
their compositions are generally magnesian and sili-
ceous, with about 4—16 weight percent MnO and as much
as 2 weight percent FeO. Chlorite that forms claylike
layers in B24 has a composition very similar to the pink
isotropic gel-like material in B24 (table 8, analysis 4),
containing only slightly less Mg and Al and slightly more
Mn and Si. There is no correlation between the degree of
crystallization of the gels and their compositions (table 8,
analyses 3—8). Coarsely crystalline chlorite is uncommon
(found only in B38 and B84), and its composition overlaps
that of the gel-like and claylike layers. Cation sums of
the coarsely crystalline chlorite are low (ﬁg. 8A, table 9,
analysis 8), suggesting that some disorder may exist
within the structure or that very minor amounts of
caryopilite may be interlayered with the chlorite. Com-
positions of the Mg-chlorites, including gels having chlo—
rite composition, are variable within single samples,
particularly with respect to the amount of manganese.
This variation may in part be due to ﬁne—scale mixing
with caryopilite. The compositional boundary between
chlorite and caryopilite-chlorite mixtures (ﬁgs. 9, 10)
is arbitrarily deﬁned at 23 weight percent MgO,
corresponding to about 3.2 Mg cations per 14 anions. The
concentrations of minor elements (Fe, Ni, Zn, V, and Ti)
vary considerably between samples. With the exception
of smectite, chlorite has the highest concentrations of Zn
and Ni. In chlorite and caryopilite-chlorite mixed layers,
Ni and Zn concentrations are greatest in the Mg-rich
mixed layers (ﬁg. 11). In some samples, chlorite and
chlorite—composition gel exhibit a range of textures,
again with no correlation between textures and compo-
sitions. In B24, however, two distinct groups of gels with
chloritelike composition were found. One, discussed
above (table 8, analysis 4) forms a layer—fragment of
essentially isotropic pink gel-like material that contains
rhombohedra of rhodochrosite. The second material with
chloritelike composition (table 8, analysis 5) is partially
crystallized and colorless; it occurs within layers of
rhodochrosite with either caryopilite or phase A (ﬁg. 8E)
and has signiﬁcantly less Al and more Si, Fe, Ni, and Zn
than the pink gel-like material.

 

A trace quantity of material with the composition of
the Mn-rich chlorite, pennantite, was found disseminated
within caryopilite in only one sample (B31, table 9,
analysis 9).

TANEYAMALITE AND PHASE A

Taneyamalite, the Mn analogue of howieite (table 1),
and a compositionally similar phase (referred to in this
report as phase A) occur as discrete crystals. More
commonly, they occur as intimate mixtures that show a
range of compositions. Because of the close association,
the two phases are discussed together.

The best formed taneyamalite occurs interlaminated
with caryopilite or chert. Taneyamalite laminations in
chert appear to be thin protoliths. When interlaminated
with caryopilite, taneyamalite also forms soft-sediment
injection veins (ﬁg. 53). Examination of one such vein in
B24 With the scanning electron microscope showed ﬁne
ﬁbers and laths. Taneyamalite also forms isolated bun-
dles or tufts of radiating ﬁbers within chert laminations
adjacent to the orebody. These ﬁbers are morphologi—
cally similar to ﬁbers of phase A. There is considerable
variation in the composition of taneyamalite, both within
a single sample and between samples (ﬁg. 9). Taneya-
malite in layers or veins contains less MgO, A1203, and
SiOZ, and more FeO and MnO, than taneyamalite in
isolated tufts (table 9, analyses 13—14). These differences
may be caused by contamination with ﬁne-grained smec-
tite, which is ubiquitous in some chert layers. All taneya-
malite analyzed in our samples is deﬁcient in sodium
relative to the ideal formula of Matsubara (1981),
NaManSi12(O,OH)44. Under the microprobe beam, the
Na X-ray count rate did not decrease with time, indicat-
ing that Na was not lost due to exposure to the electron
beam. Thus, the sodium deﬁciency appears to be real and
to reﬂect either deviations from the ideal structural
formula or intergrowth with a Na-poor phase such as
phase A (see below). Taneyamalite from other samples
has compositions within the range of the analyses plotted
in ﬁgure 12 (values given are cations per formula unit):

Sample Na Fe

65H66 ............................. 0. 722—0. 860 1229—1. 458
65H68 ............................. 0824—0859 1266—1318
B10 ................................ 0. 732—0872 0.744—1.099
B7 ................................. 0. 683—0. 862 0852—1. 101

In most samples, the nickel and zinc contents of taneya-
malite are at or below the detection limits of the micro-
probe, although in several samples (B24, B38, B71) as
much as 0.74 weight percent ZnO was found.

Phase A occurs as bundles of radiating yellow ﬁbers in
rhodochrosite (B24, ﬁg. 8E, table 9, analysis 10). Com-
pared with taneyamalite, phase A contains notably more

MINERAL CHARACTERIZATION AND CHEMISTRY 43

 

 

 

 

0.2 l I
BE] U
53 X
X§
X
0.15 —— X _
EH
m R:
c a g1 x
E Ba @1555: <§
<- 0.1 — EB EEC] —
E % EWE D
C I:
N a ﬁ 33 33 A
+ 33
2 5533 gm 5 E % Aff
x EB 63 A
0.05 — E)“; E 69 % _
x 9 69
xx x
X
A
A
o g l l l |
0 1 2 3 4 5
Mg/14 anions
EXPLANATION
824 BBC 395 BS4
Caryopilite I o x A
Caryopilite-chlorite EB 69 X
mixture
Chlorite III 0 x A

A

 

 

 

 

0.1
l I l I I
Z
O x
0.08 —- <§ x —
(b X
a; EB X X
U) _ _
.5 0.06 SD
C
3 EB '3 D [I
E e
C X
N
0.04 —% X A E III —
x 336963 A3 A E333 $33 BED
1A A. 53 EAAEEE ﬁgs (#5; CIEI
A f Dmé [III
x I EB BEI _
0.02 EBA ﬂ El D
I
o ' I I | I I I
0 0.02 0.04 0.06 0.08 0.1 0.12 0.14
Ni/14 anions
EXPLANATION
824 B80 395 BS4
Caryopilite l o x A
Caryopilite-chlorite EB 69 X
mixture
Chlorite III 0 x A
B

FIGURE 11. —Relationship between Ni, Zn, and Mg in caryopilite, chlorite, and mixtures of the two minerals from the Buckeye mine, California.
A shows that Ni and Zn are associated with Mg-rich (Chlorite-rich) compositions, but 8 shows that the proportion of Ni to Zn is variable in

the Buckeye deposit.

Fe and much less Na. The formula of phase A (table 9,
analyses 10—11), calculated on the same anion basis as
taneyamalite, yields similar cation values and sums. This
similarity suggests that phase A may be Na-depleted
taneyamalite, analogous to K—depleted zussmanite,
K0‘1A1M138i12042(OH)14, which rims K—rich zussmanite
in ferruginous metachert (Muir Wood, 1980). Although
the K—depleted zussmanite described by Muir Wood
appears to be the result of leaching, phase A occurs
intergrown with taneyamalite, occurs as well-formed
crystals, and does not appear to be a product of alter-
ation. On the possibility that phase A is the Mn-rich
analogue of deerite, (Fe+2,Mn)6(Fe+3,Al)3Si6020(OH)5,
as taneyamalite is of howieite, we calculated analyses
of phase A to the deerite formula (15 cations:22.5 oxy-
gens). Analyses of phase A approach the formula
M;2M+3Si7020(OH)5, although most analyses have
excess Si and a deﬁciency of divalent and trivalent
cations relative to this formula. Recalculation of phase A
analyses to other formula units proved unsuccessful. We

 

calculated phase A by using the taneyamalite formula
unit and effective ionic radii for tetrahedrally and octa-
hedrally coordinated cations (Shannon, 1976) and found it
to be unlike any of the modulated 2:1 structures consid-
ered by Guggenheim and Eggleton (1987, their ﬁg. 5a).
On the basis of their scheme, phase A recalculated
according to the formula unit MQLZM+?’Si7020(OH)5 has a
mean octahedrally coordinated cation radius (all cations
except Na, Ca, and Si) that ranges from 0.78A to 0.79A
and a mean tetrahedrally coordinated cation radius (Si) of
0.26A. These values plot near the boundary of the known
ﬁeld for zussmanite, suggesting that phase A is structur-
ally similar to zussmanite. Phase A most commonly
occurs in layers intergrown with taneyamalite, and its
presence is deduced from microprobe analyses (table 9,
analysis 12) showing compositions that lie between those
of taneyamalite and those of the ﬁbers in B24. Layers of
mixed phase A and taneyamalite also occur in B38 (ﬁg.
8A). In that sample, the analyses that are richest in
phase A component have signiﬁcant vanadium (table 9,

44 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

 

3-5 l I l l

Fe/37.5 anions

 

 

 

 

X X
o I l l I
0 0.2 0.4 0.6 0.8 1
Na/37.5 anions
EXPLANATION
824 BS8 BZO B71 b
Taneyamalite D o x A
Taneyamalite- I o x A
phase A mixture
Phase A EB 69

FIGURE 12. —Distribution of taneyamalite and phase A compositions in
four samples from the Buckeye deposit, California. Compositions
from B24 and B38 appear to deﬁne two distinct mixing lines,
although there is overlap at the more sodic compositions.

analysis 11), and analyses that are richest in taneyama-
lite have minor amounts of titanium, indicating that V
and Ti are associated with phase A and taneyamalite,
respectively.

Mixed layers of phase A and taneyamalite are far less
common than mixed layers of caryopilite and chlorite.
Compositions with at least 1 weight percent NazO,
equivalent to 0.5 cations per formula unit, are here
deﬁned as taneyamalite; those with 0.1 to 1 weight
percent Na20 as phase A—taneyamalite mixtures; and

 

those with less than 0.1 weight percent as phase A. In
mixtures of taneyamalite and phase A, sodium and iron
deﬁne a linear trend when plotted against each other (ﬁg.
12). Each sample deﬁnes its own trend, although the
distinction is lost as phase A composition is approached.
Analyses from sample B71b form two distinct groups,
with taneyamalite analyses displaced to higher Fe values
than those of mixed taneyamalite and phase A, suggest-
ing that an additional phase may be present in the
mixture—perhaps smectite, which would account for the
high Mg contents of the mixtures.

SANTACLARAITE

Santaclaraite is invariably associated with rhodochro-
site, quartz, and smectite. These minerals form irregular
veins or “intrusions” (B7; B20; B25, ﬁg. 40; B84, ﬁg. 4E;
B91) that cut across laminations of protolith and may
have been pathways for escaping ﬂuids. Fragments of
caryopilite layers are suspended in the “intrusion” in
ﬁgure 4C. The details of the paragenesis are not clear. It
is certain, however, that these veins postdate early
consolidation of the caryopilite and chlorite protoliths
and that much of the santaclaraite has subsequently been
replaced.

Santaclaraite is recognized by its rhodonite-like mor-
phology, brilliant second-order birefringence, and char-
acteristic X-ray reﬂections at 11.5°, 12.6°, 18.5°, and
283°. We conﬁrm the discrepancy between the strong
calculated and weak observed intensities of the (0—24)
reﬂection at 332° (Erd and Ohashi, 1984). The laths are
so characteristic that pseudomorphs of santaclaraite by
rhodochrosite can be easily recognized. These are the
pseudomorphs noted by Huebner (1967). We found no
evidence for the occurrence of either rhodonite or pyrox-
mangite at the Buckeye deposit. In B7, B20, and
65H68B, santaclaraite has a very narrow compositional
range, represented by santaclaraite in B7 (table 9,
analysis 15). Analyses yield cation proportions that are in
excellent agreement with the ideal structural formula,
CaMn4[Si5Ol4(OH)](OH)-H20 (Ohashi and Finger, 1981).
Santaclaraite from the Buckeye deposit is more homoge-
neous but has slightly more substitution of Fe and Mg for
Mn than the type material (Erd and Ohashi, 1984), which
is from another Mn deposit in the Franciscan Complex.

RHODOCHROSITE

The carbonate adopts four distinct textural forms:
micritic, spherulitic, tufted, and sparry. Sparry rhodo-
chrosite forms by the direct recrystallization of the
micrite, without passing through intermediate stages.
The spherulites are uncommon, ﬁbrous, exhibit a distinct

MINERAL CHARACTERIZATION AND CHEMISTRY 45

black cross under crossed nicols, occur in micrite, and
appear to be a secondary, probably diagenetic, feature.
The rare tufts are composed of well—aligned ﬁbers; by
analogy with the sperulites we presume that the tufts are
also diagenetic. There is no systematic compositional
distinction between these forms. Rhodochrosite from six
bulk samples (B22, B28, B35, B45, B74, B108) has cell
dimensions ofa = 4.775—4.782A and c = 15.63—15.69A,
close to the values for pure MnCO3, 4.777A and 15.66A
(Huebner, 1967). Primary micritic and recrystallized
sparry rhodochrosites commonly contain less than 2.4
weight percent each of CaO and MgO (table 11, analyses
1—4). Values of 3 to 8 weight percent CaO are uncommon
(ﬁg. 8E, table 11, analysis 5). Within a given sample,
both micritic and sparry rhodochrosite have about the
same compositional range, consistent with the idea that
the sparry carbonate is simply recrystallized micritic
protolith. Except for very small concentrations of iron,
no other elements were detected in the carbonate by
means of wavelength- or energy-dispersive microprobe
techniques. Speciﬁcally, we did not detect Ni, Ba, or Sr
in the carbonate grains. Some rhombohedra are zoned,
having slightly higher calcium in cores than in rims. A
slightly higher concentration of magnesium was found in
rims of rhodochrosite grains in one sample (B53, table 11,
analyses 2—3); these grains are in contact with braunite,
but no corresponding zonation in MgO was found in the
braunite. Rhodochrosite that forms crosscutting veins
can contain as much as about 10 weight percent CaO and
3 weight percent MgO (analysis 6 in table 11 is represen-
tative). Compositions are variable within a single vein
and between samples (see table 11 and ﬁg. 9).

BRAUNITE

Braunite has close to the ideal composition of braunite
I, (Mn+2)6(SiMn+3)024 (Baudracco-Gritti, 1985). The
composition (ﬁg. 9) varies little within a single sample,
but differences between samples were noted, particu-
larly in the iron content. The range of Fe203 is 0.20 to
1.64 weight percent (all Fe is assumed to be ferric).
Minor amounts of Ca, Mg, or Ni occur in some samples.
The compositions of braunite both in contact with and
more distant from rhodochrosite in B53 are given in table
12 (analyses 1 and 2). The calcium and magnesium values
are not signiﬁcantly different, but the braunite in contact
with carbonate contains slightly more nickel and iron and
less manganese. Braunite in B53, which forms distinct
layers, appears to be texturally variable. In parts of the
sample, braunite appears to be ﬁne grained and inter-
grown With equally ﬁne-grained rhodochrosite, but,
where in contact with coarse rhodochrosite, braunite
forms coarse grains. In other layers, cores of braunite
grains contain blebby to irregularly shaped rhodochros-

 

ite grains. Rims of these braunite grains, whether in
contact with other braunite grains or with the adjacent
rhodochrosite layer, show the same coarsening as
observed in the ﬁne-grained braunite and rhodochrosite
layers, with the same variation in composition (cores
contain less nickel and iron). In sample B80 (ﬁg. SB) a
compositional difference between coarse and ﬁne braun-
ite is also found. Coarse braunite contains signiﬁcantly
less iron and calcium and more manganese than the ﬁner
grains. Despite the range of textures and minor-element
compositions, the Si concentration in Buckeye braunite is
constant, at a value of 1 atom per 12-oxygen formula
unit. Braunite can occur with hausmannite (B6, ﬁg. 8F;
65H170), but no partitioning of iron on the basis of
weight percent concentration was found between these
two oxides (table 12, analyses 3,5 and 4,8).

HAUSMANNITE

Hausmannite has close to endmember composition.
Manganese greatly exceeds all other cations, and, when
recalculated on the basis of an ideal Mn+2Mn3304 for-
mula unit, the analyses total approximately 100 percent
(ﬁg. 93; table 12, analyses 5—8). Signiﬁcant but very
small concentrations of Fe, Mg, V, Zn, Ni, Ca, Ti, and Si
were detected. Most of the slight composition variation
was found between samples, With very small ranges of
compositions within a single thin section.

DISCUSSION

The ﬁne-grained textures and microprobe analyses
expand the idea that the ore originated as chemically
precipitated sediments in which eight components
formed layers (protoliths). The new observations are a
ninth component (phase A) and mixed layers. Thus, the
deposit consists of layers of one component and mixtures
of two components. This distribution is what would be
expected if the conditions of a sedimentary environment
changed such that one ﬁeld of deposition or chemical
precipitation was supplanted by another. The mixtures
would form if conditions coincided with those of two or
three depositional ﬁelds or, more likely, if there was
mechanical mixing as chemical precipitates settled into
place. Although the textural and microchemical evidence
indicates a lack of equilibrium during diagenesis and
metamorphism, the ﬁne laminae of contrasting composi-
tions do indicate that processes in the depositional envi-
ronment also changed, presumably in response to chang-
ing conditions, involving the rates of supply of the major
components MnO, C02, 02, H20, SiOz, and A1203 and the
minor components N a20, MgO, and FeO.

46 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

VEIN MINERALS

Two generations of veins occur. Earlier veins tend to
be narrow (less than 100 um), short (less than 2 mm;
rarely to 3 cm), and devoid of fragments of transected
layers. They are monomineralic and reﬂect the mineral-
ogy of adjacent protolith. There is no chemical distinction
between these vein minerals and the corresponding
minerals in the adjacent metasediments. Later veins are
wide (to 2 cm), traverse many layers, and commonly
contain fragments of the transected protoliths. The late
veins are polymineralic; in addition to wall-rock frag-
ments, they contain either quartz or rhodochrosite and
minerals that are chemically distinct from the minerals in
the enclosing protoliths. These veins introduced addi-
tional chemical components to the ore-forming system,
and there is a correspondence between these components
and the characteristic vein mineralogy. The presence in
the veins of titanite, a stilpnomelanelike mineral such as
parsettensite, barite, smectite, clinopyroxene, and vari-
ous unidentiﬁed minerals is evidence that Ti, K, Ba, 804,
Mg, Na, V, and As were mobile after the host sediments
became indurated. These minerals are not as Mn—rich as
the protoliths, and many plot primarily along the join
between Si and “others” (ﬁg. 9B).

Talc forms veins that cut across spherulitic rhodo—
chrosite (B71) and are composed of pale yellow “books”
about 40 um across; it has near endmember composition
with only minor substitution of Mn and Fe for Mg (table
10, analysis 1). The relative time of formation is not
known but is probably late.

Smectite and related interlayer compounds occur as
ﬁne to coarse (100-ptm) colorless ﬂakes in quartz and
rhodochrosite veins and as minute pale brown ﬂakes in
Mn-silicate layers and in laminated chert. These smec-
tites were identiﬁed by their optical properties, compo-
sition, and ex andable basal X-ray diffraction reﬂections
at 14A to 25 . Chemically analyzed (coarse) smectites
from Buckeye samples have the composition of the
trioctahedral smectite, saponite. Some smectite is simi-
lar to talc in that minor amounts of iron and manganese
substitute for magnesium (table 10, analysis 2). Other
smectites contain as much as 1.66 weight percent NiO
and 1.59 weight percent ZnO (table 10, analysis 3). In
minor-element concentrations, these latter smectites are
similar to the gel-like material of Mg-chlorite composi-
tion.

Several laths with the composition of parsettensite, a
manganese-rich stilpnomelane-like mineral, were found
in a quartz—barite—apatite vein in one sample, B29. This
phase has near endmember parsettensite composition,
and only minor substitution of Mg for Mn (table 10,
analysis 4). It is the only silicate analyzed in this study
that contains signiﬁcant barium.

 

Barite occurs in the aforementioned quartz-apatite-
parsettensite vein and as a replacement (or in-ﬁlling) of a
unique patch of colloform rhodochrosite in B29.

Apatite occurs only as a minor constituent in quartz
veins.

Diopsidic acmite with slight pleochroism in green to
yellow green occurs with titanite and apatite in quartz
veins that cut across the sample of the diopsidic acmite
bed collected from the exploration adit (B94 and B96).
The vein-forming clinopyroxene is similar in composition
to the acmitic host and was undoubtedly derived from the
adjacent layer. The concentration of jadeite component is
small, 3 mole percent (table 10, analysis 5).

The vanadium-rich silicate referred to herein as phase
B (B15, table 10, analysis 6) forms pleochroic light- to
medium-green ﬁbers and laths within quartz veins that
cut rhodochrosite. The crystals are very small, the
maximum dimensions being about 5 by 150 um, and were
too small to obtain an interference ﬁgure. The crystals
have approximately parallel extinction, maximum inter-
ference colors of low second order, and relief greater
than that of quartz. The reﬂectivity of phase B is slightly
greater than that of pyroxene. The composition of phase
B does not correspond to that of any vanadium silicate
reported by Evans and White (1987). X-ray or electron
diffraction data are needed.

Another vanadium—rich silicate, presumably a single
phase (table 10, analyses 7—9), is herein called phase C
and occurs in quartz veins in B15, the same sample as
phase B. Phase C is pleochroic pink to colorless and
forms columnar laths; rarely, hexagonal forms are
observed. This mineral has pronounced optical and com-
positional zoning, with cores being enriched in aluminum
and depleted in vanadium compared to rims. A very good
negative correlation between aluminum and vanadium
exists in all grains analyzed. This mineral is uniaxial (—)
and has parallel extinction, relief greater than that of
quartz, and a reﬂectivity slightly greater than that of
pyroxene. Its crystal habit, optical properties, and pro-
nounced zoning are like those of tourmaline-group min-
erals, but cation proportions are not consistent with the
formulae for tourmalines. It is conceivable that phases B
and C form a solid-solution series, Na0'0_0.5Ca2_3(A1,V)5
Si6024. X-ray diffraction data are needed.

An arsenian vanadium- and manganese-bearing phase,
presumably a hydrous vanadate, was found in sample
B17 (table 10, analysis 10). Phase D is orange red, has a
ﬁne-grained texture, and appears as pseudomorphs of a
phase within rhodochrosite. Its deep, color masks its
interference colors. Phase D has a reﬂectivity greater
than that of titanite but lower than that of an oxide. This
mineral has also been found in hausmannite— and

ORIGIN 47

tephroite-rich samples from the Manga—Chrome mine
located in Nevada County, Calif. (Flohr and Huebner,
unpub. data), and is probably related to ﬂinkite, Mn3
(As04)(OH)4.

Well-crystallized titanite occurs in several samples and
is associated with quartz veins, particularly in the mas—
sive diopsidic acmite, and as trains of crystals in caryo-
pilite. Titanite contains less than 1 weight percent V203
and no signiﬁcant amounts of other minor elements. Its
occurrence, particularly in the veins, is evidence that Ti
was mobile at the low temperature involved in modifying
the manganese deposit.

Trace quantities of minute yellow and gray sulﬁde
grains (probably pyrite, chalcopyrite, chalcocite, covel-
lite or digenite, molybdenite, and possibly sphalerite)
occur in veins. The grain size was not sufﬁcient for
quantitative microprobe analysis, but X-ray spectra
revealed peaks for Cu, Mo, and S. The paragenesis of the
sulﬁde-bearing veins is not certain; they are thin and
short, like the early veins, yet they introduce additional
components, like the later veins. The fact that Buckeye
minerals contain Mn+3 coupled with the absence of
appreciable quantities of iron sulﬁde in or near the
Buckeye deposit suggests that the adjacent sediment
column was never so reducing that marine sulfate was
reduced to make sulﬁde available for sulﬁde precipita-
tion. (Thus the presence of even minute quantities of
copper sulﬁdes signals a local origin or later activity.)

The veins provide clues to the mobile components after
the manganiferous sediments were sufﬁciently well indu—
rated to fracture. During the ﬁrst stage of vein forma-
tion, the manganiferous components of the protoliths
were mobile but did not travel far. During later vein
formation, the protoliths were more thoroughly indu-
rated (resulting in brecciation), and, with few excep-
tions, the manganiferous components were not mobile.
The source of the later vein ﬂuids is not known with
certainty, but in composition the vein minerals tend to
have afﬁnities With seawater and nonmanganiferous
marine sediments. Perhaps these late mobile compo-
nents were incompatible components that were carried
by ﬂuids expelled during compaction of the sediment
pile. The precipitation of well-crystallized titanite in a
low-temperature sedimentary deposit was not expected
and is evidence that Ti can be mobile and concentrated
even in an environment in which the bulk composition is
unusually low in titanium. Similarly, Brett and others
(1987) found anatase in hydrothermal vent chimneys. As
a general rule at the Buckeye, occurrence of an element
(or pair of elements) corresponds to the appearance of a
vein mineral, indicating that the chemical system was
responding in a predictable manner to changes in the
ﬂuid compositions. Thus, Ba results in the parsettensite-

 

like mineral, Ba and 804 in barite, PO4 in apatite, and Ti
in titanite.

ORIGIN

Two theories for the origin of the Buckeye and similar
deposits are currently popular and involve hydrothermal
and diagenetic processes. In the following pages we
discuss some marine sedimentary processes known to be
depositing or redistributing manganese today and exam-
ine some modern and ancient deposits in terms of these
processes. None is an excellent analogue for the Buckeye
deposit, which can best be explained on the basis of two
interacting processes such that diagenesis inﬂuences the
composition of ﬂuids that vent at the seaﬂoor. Finally,
we discuss possible tectonic settings for the Buckeye
deposit.

MODERN PROCESSES OF MARINE MANGANESE
DEPOSITION

Modern marine environments contain no interlayered
manganese carbonate-oxide-silicate deposits that are
analogues of the Buckeye deposit, but examination of the
processes that deposit and redistribute manganese at
ocean-ﬂoor environments will help us understand the
origin of deposits such as the Buckeye. In principle,
there appear to be three kinds of processes: hydroge-
nous, diagenetic, and hydrothermal. In practice, many
manganese deposits record more than one of these
processes. Hydrothermal deposition is usually thought to
involve hot ﬂuids, but recent monitoring of some seaﬂoor
vents that are associated with Mn—rich precipitates in
areas of igneous activity reveals water that is barely
above ambient seaﬂoor temperature (Hoffert and others,
1978; Vonderhaar and others, 1987). These cool vents
cast doubt on the necessity of heat for Mn transport, Mn
precipitation, and even ﬂuid movement, thereby sug-
gesting the possibility that some Mn precipitates could
be related to ﬂuid expulsion driven by compaction and
subduction instead of by thermal convection. The three
processes of Mn precipitation are better distinguished if,
for the purposes of this discussion, we deﬁne a hydro-
thermal processs as one that involves transport with a
moving ﬂuid either beneath or above the seaﬂoor, a
hydrogenous process as one that involves precipitation
and gravity settling of particulates through a stationary
ﬂuid above the seaﬂoor, and a diagenetic process as one
that involves the diffusion of dissolved species through a
stationary ﬂuid, commonly below the seaﬂoor. Each of
the three processes can result in a distinct product.
Characteristic features of these products, when com-
pared with the Buckeye ore, may provide clues to the
origin of the Buckeye deposit. These characteristics,
which are discussed below and in summaries by Margolis

48 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

and Burns (1976), Cronan (1977), Raab and Meylan
(1977), Rona (1978), Toth (1980), Fleet (1983), and Lalou
(1983), are compared with features of the deposit in table
13. Because we are discussing marine environments, we,
like Courtois and Clauer (1980), normalized the REE
concentrations of the products to the REE concentra-
tions of seawater (Hogdahl and others, 1968). In this
way, we hope that our REE patterns will help us
distinguish the signatures brought about by the opera-
tion of each of these processes upon ﬂuid that, originally,
was seawater.

Purely hydrogenous deposits form by the precipitation
and settling of manganese and iron oxyhydroxides
directly from cold seawater onto the top of the sediment
column. This process results in nodules and crusts that
are exposed at the sediment surface, thin coatings on
fresh basalt, particles of Fe—Mn-oxyhydroxides in the
sediment, and ﬁlms of oxyhydroxides that coat detrital
particles. Because the manganese concentration in nor—
mal seawater is only 0.03 parts per billion (ppb) (Bru-
land, 1983), the growth rates are slow and large nodules
and thick crusts occur only on old ocean ﬂoor. A related
factor is the occurrence of hydrogenous deposits in
tectonically and volcanically inactive areas having low
sedimentation rates. Thus, nodules and crusts commonly
are found on siliceous ooze and red clay of deep ocean
basins, below the carbonate compensation depth and far
from sources of detritus, or where currents suppress
sedimentation. The ﬁelds of nodules and crusts are
regional in extent, and surface coverage may exceed 75
percent (Glasby, 1976). Nodules and crusts are layered,
and their upper surfaces, which are directly exposed to
seawater, commonly are smoother than the lower sur-
faces, which are exposed to sediment (see Heye and
others, 1979; Marchig and Halbach, 1982). The nodules
and crusts are highly oxidized (most manganese is
quadrivalent and iron is trivalent). A characteristic Mn
phase is vernadite (the 8—Mn02 of Burns and Burns,
1977). Rhodochrosite and Mn-rich silicates, which
require manganese that is predominantly divalent, never
occur in association with the oxidized nodules and crusts
at the sediment surface. Most hydrogenous nodules and
crusts have ferromanganese compositions with Mn/Fe
ranging from 0.1 to 10. Relative to diagenetic and
hydrothermal components, the hydrogenous component
is enriched in Co, Fe, Th, Hf, and the REE (Dymond and
others, 1984). Halbach (1986) noted that hydrogenous
crusts contain, on average, 0.5 parts per million (ppm)
Pt. The seawater—normalized REE signature of hydrog—
enous products (ﬁg. 13A) is enriched overall, has a
pronounced positive Ce anomaly (large Ce/La and
Ce/Sm), and is relatively depleted in the heavy REE.
These enrichments can be attributed to long exposure to

 

seawater, to the high surface areas of the precipitates,
and to the tunnel structures of the constituent oxides.

Because of their tendency to undergo diagenetic alter-
ation and to form on oceanic sediments that will ulti-
mately be subducted, hydrogenous crusts and nodules
are uncommon in the geologic record (Jenkins, 1977). We
should look for thin (perhaps less than a meter thick) but
laterally extensive horizons of nodules or crusts that
contain, at least at low metamorphic grade, quadrivalent
manganese oxides having nodular or knobby structure,
perhaps preserving concentric layering, and having the
characteristic chemical signature. One might look for this
record in slowly deposited seafloor sediments that lie
immediately below an unconformity or that have been
preserved as fragments of ocean crust that were accreted
to continents. Audley-Charles (1965) and Margolis and
others (1978) described nodules and slabs (crusts) in
Cretaceous red clay that forms an exotic block in an
olistostrome on Timor. On the basis of their intermediate
Mn/Fe, enrichment in Co, U, and Th, and high Th/U,
some of these nodules may record a hydrogenous com-
ponent. The crust and another nodule have high Mn/Fe
and low Co and may represent hydrogenous precipitates
that have been altered by diagenetic processes. In con-
trast, the Buckeye deposit records none of the charac-
teristics of the hydrogenous processes (table 13).

Early diagenetic processes involve pore ﬂuids that act
upon substances incorporated within marine sediments.
These processes may redistribute earlier formed
manganese-rich precipitates, progressively reﬁning
their chemistry through repeated dissolution, transport,
and reprecipitation. Diagenesis that is driven by chemi-
cal disequilibrium between oxygenated seawater and
unconsolidated sediments rich in organic material results
in three chemically distinct products: ferromanganese
oxyhydroxides formed by oxic diagenesis, manganese-
rich oxyhydroxides formed by suboxic diagenesis, and
sulﬁdes and manganese carbonate formed by anoxic
diagenesis.

Processes that occur within the zone of dissolved
oxygen (oxic zone) at the top of the sedimentary column
(see Froelich and others, 1979) largely involve the redis-
tribution of components that had a hydrogenous or
detrital origin. Included are the release of trace metals
from decomposing organic tissues and the extraction of
iron and manganese from sedimentary components
(opaline skeletal debris, hydrogenous coatings on sedi-
mentary clasts) to form ferromanganese nodules and
nontronite (Dymond and others, 1984; Lyle and others,
1984). The Mn oxidation state in the top 5 cm of sediment
at two Paciﬁc sites, 3.3 to 3.9, greatly exceeds that in the
suboxic zone where Mn is divalent (Kalhorn and Emer—
son, 1984). Under the oxidizing conditions near the
surface, manganese and iron should remain in the

49

.5033 333 3000330330530003000 00 0.3030»?

 

 

 

.3003 30 > 300,3 000,3 33053 3 3.3033 3003 0303030033 3030 30> 3030 30 > ............ 0303 £30.30

300303030 300303030
00303 303,3 303,3 303,3 3330533 3330533 3330533 3333 , 33333.3 ............ 033030 333313

03330033

03330033 3333.003 03330033 03530033 03330033
3333003 00 0:02 03 0:073 0:073 quobw 333053 3 333053 3 33053 3 398.0% Emcobw ............ “133030330 00
.003 I «03 003 I 003 003 03 330533 3330533 3330533 303 B3 ..... 3333 303030005333
.303 303 303 :30533 3.303 0303030033 303 0303030032 0303030033 .............. 030300 3
.303 303 303 330533 303 303 00303003 03 303 3333.3 3333 ............. 030300 .35
.303 303 303 330533 303 333 8 0303030033 030303003 3 303 3333 33333.3 ............. 030300 00
.303 303 303 303 23.333 303 333 03 303 3333 33333 ............. 030300 32
.303 303 303 303 .393 303 3M3 03 0303030033 333.3 33.33 ............. 030300 :0
.303 30 > 303 30 > 303 30 > T530338 30333000 3003 00003303 033M333 03 303|v 303 303 ............ 030300 0038
.303 30 > 303 30 > 303 30 > T303330033300 30333000 0003 03003303 333 03 303i 030.303.0333 0303030033 ........... 030300 000333
.MW+3N+ VINIT VI? N+ NIT N+ C.V|mw.m+ VIT VAT .............. 00:®HN>§H>H
.owm I 333 33333 303 3333.3 330533 3333 03 I 3 03 I 3.0 03 I 3.0 .................. 0,3353
000.33. 0 Z 02 00 > 0 Z 0 Z 0 Z 0 Z 0 Z ......... 3.3000333 03.03333
N 000.33. 00 > 00 > 0 Z 0 Z 0 Z 0 Z 00 > 00 > .......... 3.3000333 033.3033

nIu 033303033033 3033330300 03300055 00303030 00338

M 3330:3303 03003330333 03300033333 003.308.3303 003033.300 30 03330303000 3333-03 3333-03
0 03350.33 0003333 r033030303003. .0035 3303.332 03.330333 33053 3 3333003939 .0053-» .NO:S.w .............. 3030,3033

0330-30.30 3300.033

3300033 ”303300 #:0333000 30333000
.0330 #30330 3300033 #:033300m .3033300w 033003030030> 3033300w 3033300m 30333.0m 030303333033 030303333033 ............... 030.5033w

033.3033 003.3% 0333330330 050333030 050333030
h0333.303V~ 6333.330 .05033333 0333-330 050.533 03333030 030.573 000.33% ‘030.553 k050.533 ............ 333333030
.02 073 00% 30 0 Z 00 > 3330300033 330350033 0: 30:0 00 > 02 0 Z . . . . 00.330303 3030003033
00> 00> 00> 02 073 02 02 00> 00> ............. 3300333303
30303.33m 3030u3=w 303033.35m 3003.353 3003.353 3003.533 .30 30303.33w 30303.3.Em 303033.35 303033.35m .................. .3303

33:03 .30 30333000 003300: 0003 .30 0300.30 0300.30 30303033
.0303 .0353 $05.30 .30 300.30 .30 .3033 .3033 .30 00330073 .30 00330073 335000 003300 Z ............. 30303330333
300303.33 333330300033 3303:0033 00303033 30030303 3000033303 3000033303 02 0 Z ........... 3.30003 0303
.3000. 30> 30003 30> 30003 30003 3.30003 30:03w0w3 3020303 3303033 3330.333 .................. 30333
.0233 $033 2033 20% oz 02 02 $630 $80 ................. 0:080
030325 303.333 38.3.0.3 303.6002 03053 03x03=m 0305 0305.30 0033002
30333305033033 0330:0333 00033033303333

 

30.305030 330.300.050.300: .3333 ”00:03:20 £300,035 .333 ”3003.30 30: .02.
033003303. 3033.333 03.33 33003. E0308: 0300 330 33803330 3.300030% 033003033 0333 K0 $003033§00| .3 3.3302?

5O

MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

 

 

   
      
    

 

 

 

 

     
   

   

   

   

 

   

 

I | I | I | I | I | | I
100,000 : EXPLANATION E
E E1 Hydrogenous marine I
_ Mn and Fe deposits —
— <1 Hydrothermal marine _
— Mn and Fe deposits _
_ O Shale -
X Plankton
10,000 E
8 Z
8 _
g E _
>< < -
5
E
3
3
‘2 1,000(
E
‘o‘.
E
m
a)
1
E
< J
100 : E
10 I I I I I I I I I I I I
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
A
T I l I I T I I I | | l |
100,000 — :
E EXPLANATION :
: N Red chert, Apennines I
— O Braunite ore, Apennines —
— x Yellow ore, Harlech _
— I Carbonate, average, Harlech -
)1 Red ore, Harlech
10,000 : El Microspherulitic ore, Harlech :
s = E
o _
q _ _
>< _ _
6 _
E
E
£ 1,000 E
E I
a _
E _
N
w _
100
10 | | l I I l l I l I I I I
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
C

100,000

Sample/(Seawater x 1,000)

10

La Ce Pr Nd

Hydrogenous nodules (Elderfield and others, 1981)

 

  
    
 

|||||||ll

|||l||

 
    
 

lllllll

 

 

EXPLANATION

Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

Corresponding sediments (Elderfield and others, 1981)

+
<
o
D
X
B
10,000
a
0
ca
Z
a
a
s
m
0)
L”,
a
a
E
N
(I)

Nodules
Average 5-Mn02thin crust on basalt,
Galapagos mounds

Todorokite—rich crust, Galapagos mounds
Birnessite-rich crust, Galapagos mounds
Average nontronite, Galapagos mounds
Amorphous coatings

 

   
     

l l | | | | I

EXPLANATION

O Mn-rich lenses,
Blue Jay mine

<1 Cherts, Blue Jay
mine

       

I | I l

 

  

La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

ORIGIN 51

 

   

EXPLANATION

Rhodochrosite-rich
nodules

El Surrounding rocks

10,000

 

8
o
o.
2 1,000'
5
‘5
3
(U
('D
‘2
E
Ta.
E 100
(I)
10 | I l | | | l I | I I | J
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
E

FIGURE 13. —Seawater-normalized rare-earth-element (REE) patterns
of Mn- and Fe-rich deposits, exclusive of the Buckeye deposit and
other materials. A, Recalculated from Fleet (1983) and Gromet and
others (1984, average shale). Hydrogenous marine Mn-Fe deposits
have the greatest REE enrichment and Ce anomaly and are depleted
in heavy relative to light REE. Hydrothermal marine Mn and Fe
deposits are not as enriched overall, lack the Ge anomaly, and are ﬂat
(heavy REE neither enriched nor depleted). Shales are less
enriched, overall, but have both the Ce anomaly and heavy REE
depletion. Plankton show the least enrichment, relative to seawater,
and lack fractionation between individual REE. B, Relationship
between REE pattern and kind of precipitate. Hydrogenous nodules
and the sediments on which they sit (Elderﬁeld and others, 1981)
have the greatest enrichment, the most pronounced Ce anomaly, and
the greatest depletion of heavy relative to light REE. Thin 8-Mn02
crusts on basalt dredged near the Galapagos mounds (Corliss and
others, 1978) combine the overall enrichment and Ce anomaly that
are characteristic of hydrogenous deposits but lack the depletion in
heavy REE. Various crusts and sediments dredged from the Galapa—
gos mounds (Corliss and others, 1978), and interpreted as having a
hydrothermal origin, show a weak Ce anomaly but have acquired
neither the overall REE enrichment nor the depletion in heavy REE
that is characteristic of the hydrogenous deposits. REE patterns for
the Buckeye ores (ﬁg. 7B) are distinct from those of the hydrogenous
materials but overlap with patterns of materials dredged from the
Galapagos mounds. C, REE patterns for braunite—rich ore Z—17 and
red chert Z—lR from the Apennines (Bonatti and others, 1976) and

quadrivalent or trivalent state and be insoluble in the
pore ﬂuids. However, some manganese remobilization
occurs and may contribute substantially to nodule
growth. Dymond and others (1984) concluded that oxic
diagenesis was the principal process forming nodules at
the three Paciﬁc sites studied and resulted in Ni- and
Cu-rich manganese oxide (todorokite) with Mn/Fe of 5 to
10. Under steady-state conditions in which oxygen was
being consumed by the sediment, Sundby and others
(1986) and Westerlund and others (1986) found that
surﬁcial sediment immobilized Fe, Mn, and Co from both

 

 

  
 
    
 
   
      

 
   

 

 

 

I I I | | | I | I | I | l
8
O
0.
‘— Stage3
if 5-Mn02
CD
‘5
3
EU
in — —
Q
%_ Caryoleite
E Braunlte
(U
m
Todorokite-birnessite
Hausmannite
Rhodochrosite
Fe-rich
clays and oxides Stage 1
I I I I I I L 1 I I I I I
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

F

for the red, microspherulitic, yellow, and carbonate ores from Harlech,
Wales (Bennett, 1987). D, REE patterns of chert and the Mn-rich
lenses at the Blue Jay mine (Chyi and others, 1984, their table 2). E,
REE patterns for rhodochrosite nodules and associated mudstone,
chert, and shale from the Inuyama district, Japan (Matsumoto, 1987).
The signature of the Inuyama rhodochrosite appears to be due to
contamination by, or inheritance from, the surrounding rocks, which
include chert, shale, and mudstone. The pattern of the Inuyama
rhodochrosite is more enriched in REE, and has a more pronounced Ce
anomaly, than either hydrothermal marine crusts (ﬁg. 138) or Buckeye
rhodochrosite (ﬁg. 78). F', Sketch, not drawn to scale, showing possible
evolution of the REE patterns of sediments precipitated by vent ﬂuids
that are progressively diluted with seawater. Stage 1 represents
Fe-rich clays and Fe-rich oxides precipitated from cool ﬂuids near a
vent. Stages 2 and 3 represent Mn-rich sediments associated with
mounds or vents, perhaps on stage 1 deposits. Stage 4 represents
Mn-rich and Mn-Fe oxides deposited at some distance from the mounds
and includes hydrogenous nodules.

the overlying water and the underlying (and presumably
suboxic) sediment and released Cd, Cu, Ni, and Zn to
both the overlying water and underlying sediment. Con-
sidering the mobility of these trace elements at the
sediment-seawater interface, it is not surprising to ﬁnd
ferromanganese nodules enriched in C0, Cu, Ni, and Zn.
Were oxic processes to have terminated in carbonates
and silicates, instead of todorokite, we would still expect
the deposit to contain unusually high Co, Ni, Cu, and Zn
concentrations and to have relatively low Mn/Fe. These
are not the characteristics of the Buckeye deposit.

52 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

Organic matter, incorporated in the sediment column,
decomposes by a progressive series of suboxic and anoxic
COZ-producing reactions that consume oxygen, reduce
MnOz, form N2, reduce Fe203 (or FeOOH), reduce
sulfate to form sulﬁde, and ﬁnally form methane
(Froelich and others, 1979). These reactions provide a
mechanism for supplying soluble Mn+2 (and Fe”) to
pore ﬂuids. These cations can then diffuse upward to the
oxic zone, where they precipitate as oxyhydroxides near
the top of the sediment column (Lynn and Bonatti, 1965).
The intensity of reduction is favored by excess of organic
matter (high biological productivity) and time: ferric
oxyhydroxides will not be reduced before quadrivalent
manganese oxyhydroxides and nitrate (Sawlan and Mur-
ray, 1983), and methane will not be formed before
manganese, iron, and sulfate have been reduced. Thus,
methanogenesis requires a supply of organic reductant
that exceeds the supply of the preceding oxidizing
agents. The diffusion of dissolved Mn+2 (and Fe”) in
response to an oxygen potential gradient is dependent
upon the presence of nearby oxidized seawater. Because
diffusion of Mn” is a relatively slow process, it is
favored by slow sedimentation (and can be arrested by
deposition of a thick, perhaps organic—rich, overburden
that displaces the oxidizing seawater). Dymond and
others (1984) identiﬁed the product of suboxic diagenesis
in nodules from only one of the three Paciﬁc sites studied
and concluded that the suboxic process was intermittent
and was related to pulses in the supply of organic debris.
The characteristic product forms nodule bottoms, has a
high Mn/Fe (20—70), and is enriched in Sb (124 ppm) and
U (7.7 ppm). Had suboxic diagenesis been sufﬁciently
intense to produce sulﬁde or methane, ferric iron would
have already been reduced, resulting in the formation of
sulﬁde, in situ, or precipitates with smaller Mn/Fe. The
Buckeye deposit offers the organic matter (debris of
radiolarians) required by the suboxic model, but the
model does not explain the absence of replacement
features, low Sb and U concentrations, and the lack of
iron-rich precipitates that would be expected, particu-
larly if methane was produced.

Mn—rich carbonate within modern sediment columns
has been discussed by Suess (1979) and by Pedersen and
Price (1982). The carbonate occurs in coarse-grained or
porous horizons, such as layers of volcanic ash, and the
primary precipitates have spheroidal or botryoidal tex—
ture. Calvert and Price (1977) considered the carbonate
activity to be the critical factor in the presence or
absence of Mn-rich carbonate. In contrast, Pedersen and
Price (1982) regarded a plentiful supply of (hydrogenous)
manganese oxides to be more important for rhodochros-
ite formation than an unusual supply of dissolved carbon-
ate. The latter opinion is important because it suggests
that, following reduction and dissolution, manganese

 

would be reprecipitated as carbonate, close to the site of
original oxides. In either case, it is difﬁcult to imagine
the in-situ production of sufﬁcient Mn or CO2 to produce
a thick layer of rhodochrosite. Transport over great
distances, exceeding that possible by diffusion alone,
appears to be be necessary.

No modern diagenetic carbonate has a composition of
the rhodochrosite endmember. Many recent manganifer-
ous carbonates contain calcium (but only low concentra—
tions of Mg and Fe), and those of Loch Fyne, Scotland,
coexist with calcium carbonate (Calvert and Price, 1970).
Ancient Mn-rich carbonate and rhodochrosite, inter-
preted to have had an origin involving marine diagenesis,
occur in the Inuyama district, Japan (Matsumoto, 1987).
In most respects, Buckeye rhodochrosite is unlike
Inuyama carbonate. Like the modern diagenetic Mn-rich
carbonates, Inuyama rhodochrosite is spherulitic and
nodular. The Buckeye rhodochrosite is layered (and
interbedded with other Mn-rich phases). There is only a
small proportion of radial (tufted) or spherulitic rhodo-
chrosite at the Buckeye, and that material clearly
replaces earlier, extremely ﬁne grained, unoriented
rhodochrosite. Inuyama rhodochrosite occurs in bedded
chert and has similar large Mn/Fe but smaller concentra-
tions of Cu and Ni, much higher REE concentrations, a
better developed Ce anomaly (ﬁg. 13E), and less nega-
tive values of 813C.

Hydrothermal deposits precipitate from mixtures of
ﬂuids that were originally marine: fresh seawater drawn
beneath the seaﬂoor by convection; connate seawater
from the sediment column; possible juvenile ﬂuids; and
more fresh seawater. As it circulates through sediments
and (or) igneous rocks, the ﬂuid is reduced (Eh potential
lowered) by interaction with organic matter (see Froelich
and others, 1979) and (or) ferrous-iron-bearing rocks,
acidiﬁed by the removal of Mg to form smectitelike and
talclike phases (Thornton and Seyfried, 1985), and pos-
sibly heated. This ﬂuid leaches manganese and iron from
the column of sediments and igneous rocks and trans-
ports both dissolved and suspended constituents toward
the seaﬂoor. Dilution and cooling, brought about by
mixing with fresh, relatively oxidizing and alkaline, cold
seawater, causes precipitation within the fracture sys-
tem or near seaﬂoor vents. These precipitates form
rapidly. Manganese that does not settle to the seaﬂoor
can be greatly diluted with normal seawater (Baker and
Massoth, 1986) and carried for thousands of kilometers
between the seaﬂoor and surface by buoyant ﬂuid plumes
(Klinkhammer and others, 1986). These plumes are asso—
ciated with metalliferous sediments (Edmund and oth-
ers, 1982; Klinkhammer and Hudson, 1986) and may be a
source of hydrogenous manganese that is subsequently
reworked by diagenetic processes.

ORIGIN 53

The mineralogy and chemistry of marine hydrothermal
deposits are varied but distinctive. The common Mn-rich
phases are birnessite and todorokite, rather than verna-
dite as in the hydrogenous nodules and crusts. Some
hydrothermal precipitates are described as gel-like (Hof-
fert and others, 1978; Brett and others, 1987; Haggerty,
1987). Nontronite (see Bischoff, 1972; Corliss and others,
1978; and Alt, 1988) and talc (Lonsdale and others, 1980)
are also found. Manganese and iron can be efﬁciently
fractionated from each other in a pH- or an Eh-potential
gradient (Krauskopf, 1957; Hem, 1972), such as might be
brought about when reducing and acid solutions contain-
ing dissolved iron and manganese are progressively
diluted with oxidizing and relatively alkaline seawater
(see Edmund and others, 1979). Progressive dilution
should result in a continuous distribution of Mn/Fe
ratios. However, most deposits considered, on the basis
of close spatial association with seaﬂoor vents that
debouch ﬂuids, to be hydrothermal have Mn/Fe values
that are either low (less than 0.1) or high (greater than
7). The relative scarcity of demonstrably hydrothermal
deposits that combine low trace-element concentrations
and intermediate Mn/Fe (values of 0.1 to 10) may be due
to inadequate sampling or absence of iron in the rising
solutions. (Large Mn/Fe values would occur if iron was
precipitated as sulﬁdes or oxyhydroxides at depth.)
Alternatively, different reaction kinetics or an excess of
MnO2 (relative to the supply of reducing agent) might
cause manganese but not iron to be dissolved by the
circulating solutions (see Thornton and Seyfried, 1985).
Obviously, if iron was not taken into solution, iron-rich
precipitates would not be present. Hydrothermal depos-
its have low concentrations of Co, Ni, Cu, and REE, no
Ce anomaly relative to fresh seawater, and a ﬂat REE
pattern (ﬁg. 13).

Hydrothermal deposits will have diverse manifesta-
tions in the geologic record. One should look for sulﬁde,
oxyhydroxide, and silica vein systems and vents (Boyce
and others, 1983; Herzig and others, 1988) that cut across
sedimentary layering; Mn-rich or Fe-rich crusts that
formed on the seaﬂoor adjacent to vents (Bonatti and
others, 1976) and, in the case of Mn-rich crusts, under-
lying sediments such as ferruginous clays that lack
detrital-terrestrial, pelagic, and volcaniclastic compo-
nents; or umbers on ophiolites. The hydrothermal pre-
cipitates may consist of layers and laminations that differ
in composition, color, or texture. Because of their close
association with ﬂuid-ﬁlled ﬁssures, hydrothermal vein
deposits might be expected to have small lateral extent.
Vented warm ﬂuids would rise and quickly mix with
fresh seawater, causing nearby precipitation. The 56-
km2 sheet of bedded nontronite on the ﬂoor of the Red
Sea (Bischoff, 1972) is an exception, almost certainly
caused by venting of hot but dense brines, thereby

 

restricting vertical mixing with fresh seawater (and
favoring lateral spreading). Long-lasting hydrothermal
activity could result in a suite of deposits that record
either separate hydrothermal pulses (the opening and
closing of subsurface ﬂuid conduits in response to
mechanical disturbances and to blockage of channels by
rapidly formed precipitates) that are progressively
younger, upward, in the sediment column, or the lateral
movement of the ocean ﬂoor relative to sources of heat
(Crerar, 1982). Many Mn-rich deposits described as
having a hydrothermal origin at or beneath the seaﬂoor
are known from midocean ridges and tectonically active
areas. Their preservation may be due to rapid burial by
sediments or volcanic rocks, a process that would pre-
vent upward mobilization of manganese by diagenesis.

MODERN HYDROTHERMAL ANALOGUES

Features of the Buckeye deposit are more similar to
those of hydrothermal deposits formed on the seaﬂoor
than to those of subsurface hydrothermal, hydrogenous,
or diagenetic deposits (table 13). Major differences
between the Buckeye and modern marine hydrothermal
Mn deposits are the greater compositional and mineral-
ogical diversity, the absence of volcanic bedrock, and the
presence of chert, shale, and graywacke at the Buckeye
deposit. Ironically, the best studied system at which the
hydrothermal ﬂuids interact with sediments (Guaymas
Basin) does not precipitate any Mn-rich crusts, probably
because of high ﬂuid temperatures at the vents. How—
ever, other manganiferous lenses in the Franciscan Com-
plex, long considered to have the same origin as the
Buckeye deposit (Taliaferro and Hudson, 1943), are
associated with basalt. Elsewhere, ancient Mn-rich
lenses of apparent hydrothermal origin are intimately
associated with chert. In the following paragraphs we
will examine these differences and attempt to explain
how they could arise from rising ﬂuids, using as the
starting point for the discussion processes operating at
moundlike or ridgelike deposits that are now forming
at midocean ridges.

Mounds occur 20 km south of the Galapagos spreading
center at 86° W. longitude. Lonsdale (1977) described
their distribution and morphology. Corliss and others
(1978) and Schrader and others (1980) gave mineralogical
and chemical analyses of dredged samples. Physical-
property measurements and core examinations were
summarized by Dymond and others (1980), Hekinian and
others (1978), Hoffert and others (1980), and Honnorez
and others (1981). Galapagos mounds are steep-sided
cones, 5 to 20 m high, aligned in chains that are presum—
ably related to fractures in basaltic bedrock. In plan view
they are similar in size to the Buckeye ore lens (25 to 100
m in diameter). The internal structure of a mound is 15 to
30 m thick, extends below the elevation of the surround-

54 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

ing surface, and consists of nontronitic clay that overlies
calcareous and siliceous nannofossil ooze and is capped
with a todorokite—birnessite crust. Cores through the
sediment column adjacent to the mounds revealed only
nannofossil ooze. Although no ﬂuid venting has been
observed on the Galapagos mounds, the heat ﬂow is
greater through the mounds than through the surround-
ing ﬂat sediments. Measured thermal gradients are not
linear, and fluids 2 to 8 °C above ambient bottom-water
temperature seeped from holes punched into the
mounds. These observations indicate the presence of a
hydrothermal system and suggest that Mn (and Fe) could
be leached from underlying basalt and sediments by
convecting ﬂuids. The chemical signatures of the Galapa—
gos sediments are distinctive: Mn/Fe exceeds 82 in the
Mn crusts, Mn/Fe is less than 0.02 in the nontronite
layers, and both materials have small concentrations of
Co, Ni, Cu, and Th, as at the Buckeye. The REE
patterns are also strikingly similar (compare ﬁgs. 78,
133). Thus, the Galapagos mounds, which show evidence
of active hydrothermal processes, have some of the same
morphological and chemical features of the depositional
process recorded by the Buckeye deposit. However, the
graywacke, the chert and shale, and the compositional
diversity Within the ore lens, all of which occur at the
Buckeye, are missing from the Galapagos mounds.

Manganese-rich crusts have been dredged from the
base of the median—valley scarp at the Trans Atlantic
Geotraverse (TAG) site on the Mid-Atlantic Ridge (Scott
and others, 1974). Crust from site 13 is composed of 5- to
10-mm-thick layers of birnessite with minor todorokite.
Radiometric data indicate rapid growth of the crust, 0.1
to 0.2 m/106 yr. The chemical signature is not distinctly
that of either a suboxic diagenetic or a hydrothermal
process. The concentrations of Co (maximum 25 ppm),
Cu (less than 119 ppm), and Ni (less than 660 ppm) are
relatively small, and the U/Th232 ratio ranges from 2.3 to
6.6. The large Mn/Fe (the minimum value is 364) favors
a hydrothermal process, but the concentration of U (9 to
16 ppm) is larger than in other hydrothermal deposits.
Ferromanganese crust with 1- to 5—mm laminations,
dredged from TAG sites 2B and 10G, had similar concen-
trations of U (11 to 18 ppm) but, in contrast, much
smaller U/Th232, suggestive of a hydrogenous process.
Inﬂuenced by a small increase in the temperature of
bottom water and by the presence of Fe-rich suspended
particles, Scott and others (1974) concluded that the
crust from site 13 was deposited by a hydrothermal
spring. Klinkhammer and others (1986) revisited the
area and found anomalously high temperatures (as great
as 0.6 °C above ambient) and concentrations of manga-
nese and silica in ﬂuids collected near vents, supporting
a hydrothermal origin for crust at site 13.

 

Seaﬂoor deposits in the French American Mid-Ocean
Under Sea (FAMOUS) area of the Mid-Atlantic Ridge
were retrieved by Hoffert and others (1978). Water
samples taken near two vents showed neither a temper-
ature nor a composition anomaly, but a thin wedge of
surﬁcial precipitate spread outward from each vent,
decreasing in thickness from 1 to 0.1 m and covering an
area of about 40 by 15 m. The underlying sediment was
fossiliferous and CaCO3-rich. Surﬁcial materials rich in
Fe and Si and composed of glauconite-celadonite and
nontronite occurred near the vents, Whereas Fe-Mn
oxyhydroxides lay farther from the vents. With distance,
there was the following progression of materials:
“gel-like inclusions of manganite, gel-todorokite, gel-
rancieite, birnessite-rancieite, and gel—cryptomelane”
(Hoffert and others, 1981, p. 81). Some interlayering
between clay-rich material and oxide—hydroxides
occurred. The clays had Mn/Fe values of 0.1 to 0.3, low
concentrations of Co, Ni, and Cu, and Co/Zn values less
than 2, thereby falling into the Fe—rich hydrothermal
category. The Mn-rich materials had Mn/Fe values of 2 to
3, like many hydrogenous and diagenetic deposits, but
had lower concentrations of Co, Ni, Cu, and Zn than oxic
diagenetic or hydrogenous deposits. Because of their
association with vents and ferruginous clays, we favor an
origin involving hydrothermal over suboxic-diagenetic
processes. Perhaps these Mn-rich materials are samples
of the intermediate stage in the fractionation path
between Fe-rich and Mn-rich hydrothermal precipitates.
In contrast, basalts collected some distance from the
vents are covered with Fe-rich crusts. These coatings
have Mn/Fe values of 0.3 to 0.5, high concentrations of
Co, Ni, and Cu, and a Co/Zn of 5.7 to 8.2. This different
chemical signature reﬂects hydrogenous or oxic-
diagenetic processes. Although the composition of the
crusts found adjacent to the vents is too Fe-rich to be a
good analogue for the Buckeye, the FAMOUS site
demonstrates that hydrothermal processes develop tab—
ular (almost lenslike) bodies, form gel-like materials,
and develop compositional diversity by fractionation on a
very local scale.

Recently, a new mounds ﬁeld has been discovered in
the Mariana back-arc basin. Some observations may bear
upon the origin of deposits such as the Buckeye. Leinen
and others (1987) reported that the Mariana mounds are
associated with high heat ﬂow and faults, have morphol-
ogies ranging from cones 2 m high to hummocks half that
height, and are associated with ﬁnely laminated sedi-
ments. Leinen and others (1987) also related the forma-
tion of nontronite to a supply of biogenic silica. Fluid
advects from the centers of the mounds over a distance of
30 m (Leinen and others, 1987; Wheat and McDuff, 1987;
Dadey and Leinen, 1987). Although we do not discount
the importance of precipitation at the seawater-sediment

ORIGIN 55

interface, lateral ﬂow parallel to sediment laminations
provides a mechanism for distributing additional ﬂuxes
while preserving earlier formed layers of contrasting
composition.

Vents in the Guaymas Basin, on the northern exten-
sion of the East Paciﬁc Rise (EPR), debouch hydrother-
mal ﬂuids that interacted with a thick sedimentary pile
after leaving the presumed basaltic basement (Von
Damm and others, 1985). The sediments in the southern
trough consisted of siliceous skeletal debris and detrital
clays but, unlike the Buckeye deposit, had 10 to 15
percent Ca003. The Mn content of the solutions was not
as great as at the sediment-starved ridge vents, presum-
ably because the vented solutions were alkaline and with
pH of 6.5, due to dissolution of CaCO3 and decomposition
of organic matter to produce NH3. Nevertheless,
reported Mn concentrations exceed those of seawater by
a factor of 104, and Lonsdale and others (1980) found
surﬁcial Mn-rich crusts on talc in the northern trough of
the basin. All Mg and sulfate could be attributed to
mixing or contamination with fresh seawater, but the
appreciable H28 content suggests that degradation of
organic matter was more than sufﬁcient to mobilize Fe+2
and that ferrous iron sulﬁde should have precipitated at
depth. Vent temperatures exceeded 250 °C. Bowers and
others (1985) modeled both the interaction of hot EPR
ﬂuids with Guaymas sediment at 315 °C and the mixing of
315 °C Guaymas vent ﬂuid with fresh, cold seawater. The
calculations were not carried to temperatures lower than
5 °C, corresponding to ﬂuid mixtures containing less than
1 percent of the original hydrothermal component.
Quartz, talc, sulﬁdes, and smectite were the dominant
precipitates, and the oxygen fugacity did not increase
sufﬁciently to precipitate quadrivalent Mn oxides. Bow-
ers and others did not, unfortunately, consider the
interaction of fresh seawater with sediments alone, par-
ticularly calcium-carbonate—free sediments, as occur at
the Buckeye. Campbell and others (1988) found that,
upon dilution, dissolved Mn emitted from Guaymas vents
precipitated within 1 week (or within several kilometers)
and settled, contributing to the Mn anomaly in the
surface sediments.

Data for the REE-enriched hydrogenous nodules and
crusts abound, but REE analyses of the relatively
depleted hydrothermal deposits, modern or ancient, are
scarce. Preliminary analysis suggests that three charac-
teristics of the REE patterns can be correlated with the
nature of the mineralizing process (table 13). Relative to
seawater, the REE patterns of hydrothermal oxyhy-
droxides and clays are less evolved—depleted and ﬂat—
than are patterns of hydrogenous materials (ﬁg. 13A).
Examination of REE patterns for a variety of materials
from modern deposits (ﬁgs. 13A,B) suggests that a Ce
anomaly can develop without appreciable overall enrich-

 

ment and that, even in the region in which enriched
patterns overlap, the hydrogenous patterns show a
depletion of heavy REE whereas the hydrothermal
patterns are relatively ﬂat. These observations suggest
the following progression of features with increasing
degree of REE enrichment (ﬁg. 13F). The ﬁrst charac-
teristic to develop is the Ge anomaly (pattern 2), which
becomes distinct before appreciable overall enrichment
in REE (pattern 3). The last characteristic to develop is
the depletion in heavy REE (pattern 4) relative to the
intermediate REE (pattern 3). This last pattern proba-
bly develops long after the third stage and may be the
result of diagenetic as well as hydrogenous processes.
The earlier stages of this sequence correspond to the
progressive dilution of the vented ﬂuids. Comparison of
ﬁgures 7A and 13E suggests that the Buckeye protoliths
have characteristics that are little evolved relative to
seawater and plankton and that are similar to patterns
from hydrothermal materials.

It is clear that, although no modern analogue for the
Buckeye deposit has been found, marine hydrothermal
deposits record some of the ﬂuxes that were previously
proposed for the Buckeye. Further, the late-stage and
highly fractionating hydrothermal processes in the
marine environment record some of the geochemical
signatures that were found in the Buckeye protoliths.
Sources for two of the sedimentary ﬂuxes (dissolved Mn,
Si) are found in the venting solutions at the modern
deposits. At vent sites that have no siliceous sediments,
the dissolved silica is precipitated as iron-bearing clays,
late in the progressive mixing of vented ﬂuids and
seawater. At the Guaymas Basin, where hot ﬂuid inter-
acts with a siliceous sediment column that is thicker than
the chert-rich section at the Buckeye, quartz is a major
product of precipitation. Perhaps, were the solutions
cooler or in contact with siliceous sediment for a shorter
time, late-stage Mn silicates might precipitate before the
Mn is oxidized and precipitated as oxyhydroxides. Fresh
seawater, a possible source of O2 to precipitate hausman-
nite and braunite and of Mg necessary for Mg-rich
chlorite and Mg-substituted caryopilite, is ubiquitous
near the vents. Missing from the modern hydrothermal
environments is direct evidence for the dissolved-002
ﬂux, although measured high values of alkalinity and pH
(Von Damm and others, 1985) are suggestive of its
presence. Carbon-dioxide species must have been
present in the modern systems because skeletal CaCO3 is
present in some sediments and because diagenetic pro-
cesses make HCO; available to the solutions during all
stages of the oxidation of organic matter (Froelich and
others, 1979). As in the case involving a deﬁciency of
dissolved silica to precipitate Mn silicates, a higher
concentration of dissolved CO2 is needed to precipitate

56 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

Mn as MnCO3 before quadrivalent oxides. Perhaps abun-
dant siliceous turbidites, such as those at the Buckeye,
can provide both.

Were siliceous ooze to be abundant in the modern
mounds-forming environments, in which cool ﬂuids are
vented, there could be three important chemical differ-
ences. (1) The siliceous skeletal debris could provide
soluble opaline silica to form the Mn-silicate proto-
liths—the gel—like materials and clays of caryopilite,
taneyamalite, and gageite compositions. (Without the
source of easily soluble silica, hot solutions would be
necessary to dissolve silica from silicates in basalt in the
sediment column.) (2) Abundant organic debris would
decompose and ultimately be oxidized to HCO; or COZ,
providing the dissolved carbonate necessary to precipi-
tate rhodochrosite. (3) A consequence of the decomposi-
tion of organic matter and CO2 generation is the con-
sumption of dissolved oxygen in fresh seawater to
produce ﬂuids with intermediate to low redox potentials.
These reduced ﬂuids would stabilize Mn+2 ions so that
Mn could be dissolved and transported from sediments
and (or) basalt. Progressive dilution with fresh seawater
at the vent would reverse the process by increasing the
alkalinity and Eh of the ﬂuids, causing rhodochrosite,
hausmannite, braunite, and Mn-silicates to precipitate.
Pulses in the ﬂux of silica and organic matter, caused by
nearby turbidites of radiolarian sand that occurred at the
Buckeye, provide one possible mechanism for developing
layers and laminations with strong compositional con—
trast. Each turbidity ﬂow would immediately disturb the
ﬂuids near the seawater—sediment interface, more slowly
inﬂuence the dissolved silica concentrations, and ulti-
mately, through diagenetic processes, inﬂuence the ﬂu-
ids that contribute to the supply of Mn and C02. No
modern deposits of rhythmically interbedded chert and
shale protoliths are known (Hein and Parrish, 1987), with
or without hydrothermal precipitates. A clue to their
occurrence might be the association of a zone with very
high productivity of siliceous pelagic organisms, seaﬂoor
vents that discharge cool ﬂuids, and a seaﬂoor that lies
below the carbonate compensation depth.

SOME POSSIBLE ANCIENT ANALOGUES

Many manganese-rich deposits occur in the geologic
record. A review of some of these deposits, particularly
those associated with interbedded chert and shale, may
provide additional clues for the origin of the Buckeye
deposit. Because none of these deposits is as little
metamorphosed as the Buckeye, their mineral associa-
tions or assemblages differ from those of both the
modern deposits and the Buckeye. In the following
paragraphs we will discuss how the Buckeye protoliths

 

might change with increasing metamorphic grade and
discover that some metamorphosed deposits record orig-
inal features that are similar to sedimentary features
recorded by the Buckeye.

PROJECTED METAMORPHIC PATH

Manganiferous sediments that have been subjected to
metamorphic temperatures greater than those experi-
enced by the Buckeye deposit should develop distinctive
suites of Mn-rich minerals. Many possible minerals are
included in the compilation (table 1). Of this list, the
minerals that will actually be present depend as much
upon the bulk composition of the manganese deposit as
upon metamorphic grade. In the case of the Buckeye,
Mn, Si, C, O, H, and perhaps Al and Mg are the
controlling elements. The minerals most likely to form
include tephroite, pyroxmangite, spessartine, bixbyite,
chlorite, and perhaps the rare Mn-chlorites, pennantite
and gonyerite. Tephroite and pyroxenoids will form at
the margins of the lens by reactions involving rhodo-
chrosite and chert (Huebner, 1967). Under appropriate
conditions and suitably high fH20 and sz (or low fCOz),
the humite-group minerals sonolite, manganhumite, and
alleghanyite might appear. We suspect that these min—
erals are more common than has been previously recog-
nized. The scarcity of MgO and FeO in the Buckeye
deposit is not conducive to the formation of amphiboles
(tirodite—dannemorite). The paucity of Na20, MgO, and
Fe would severely limit the appearance of pyroxenes—
donpeacorite, kanoite, johannsenite, or the unnamed
NaMnJr38i206 found by Ashley (1986) and by Lucchetti
and others (1988)—and the amphibole kozulite. The
scarcity of CaO, A1203, and Na20 make the formation of
akatoreite, piemontite, the manganese pumpellyites,
carpholite, and most hydrous pyroxenoids unlikely.
Thus, a consequence of a restricted bulk composition is
that the chemically simple phases should persist over a
wide range of metamorphic conditions. At the Buckeye,
much of the braunite, hausmannite, and rhodochrosite
should persist to higher metamorphic grades.

Huebner (1967, 1976) predicted that, with increasing
temperature and fC02 sufﬁcient to stabilize rhodochros-
ite, 1‘02 would fall or maintain a relatively constant value
but would not rise. There would be a series of decarbon—
ation reactions by which protoliths of rhodochrosite,
hausmannite, braunite, and quartz would react to assem-
blages that contain tephroite and (or) a pyroxenoid.
Bixbyite or quadrivalent oxides would not appear unless
the protolith was more highly oxidized than at the
Buckeye.

The caryopilite, taneyamalite, and gageite protoliths,
which were not originally considered by Huebner (1967,
1976), plot very near the Mn-Si join, shown in

ORIGIN 57

 

 

2 Others

EXPLANATION

Ideal compositions:

Rh-Rhodonite
Sp-Spessartine

Br-Braunite
Rdc-Rhodochrosite

Ca-Caryopilite H-Hausmannite
T-Tephroite Bi-Bixbyite
Ta-Taneyamalite CI-Clinochlore

FIGURE 14,—Predicted greenschist-facies assemblages for Buckeye
protoliths in which the principal Others components are A1203 and
MgO in chlorite and gel of chlorite composition. Sample numbers
refer to analyzed protoliths (table 6). A, Projection from 02, 002, and
H20 onto the plane MnO—SiOZ—EOthers. This diagram is suitable
for silicate-rich and carbonate-rich protoliths in which most manga-
nese is divalent. Mole percent. Symbols for protolith samples are the
same as in ﬁgure 9D. B, Projection onto the plane MnO—Si02-02
from all other components. This diagram is suitable for protoliths

MnO—Si02—02 and MnO—SiOz-EOthers projections
(ﬁg. 14). A gageite-rich protolith, represented by 65H79,
might transform to a mixture of hausmannite, tephroite,
and vapor at relatively high f02 but, if the ambient fC02
was sufﬁciently high, rhodochrosite would take the place
of hausmannite. A carbonate-gageite protolith such as
B50 would also yield carbonate, tephroite, and perhaps
hausmannite. Caryopilite-rich protolith, represented
by B84 (caryopilite—carbonate-chlorite), should react to
carbonate, tephroite, spessartine, and vapor. Our
caryopilite—carbonate protolith (B112) would yield the
same minerals, but in different proportions. Were the
protolith very caryopilite-rich, pyroxmangite would be a
reaction product. Because the manganese in these

 

  

CI

 

 

\
Sio2 Rhysp’Ta ca T 65H79, MnO
B1 12
EXPLANATION
Ideal compositions:
Rh-Rhodonite Br-Braunite
Sp-Spessartine Fldc-Rhodochrosite
Ca-Caryopilite H-Hausmannite
T-Tephroite Bi-Bixbyite
Ta-Taneyamalite Cl-Clinochlore
B

rich in hausmannite, braunite, or bixbyite. The anticipated metamor—
phic assemblages involve tephroite, rhodonite, and spessartine, in
addition to the low-temperature phases braunite, hausmannite,
rhodochrosite, and quartz, which will persist to high metamorphic
grade. Because there are so few minerals in chemically simple systems
such as those at the Buckeye, these metamorphic assemblages should
persist over a range of conditions. Assemblages involving spessartine
and one of the oxides, hausmannite, braunite, or bixbyite, are unlikely
because chlorite does not occur in the “oxidized” protoliths.

assemblages is predominantly divalent, braunite would
not appear. More oxidized layers, represented by a
braunite-rich protolith such as B95, would form a
tephroite-hausmannite or tephroite-carbonate assem-
blage if 02 were free to leave the system, but if f02
remained constant, braunite could persist to higher
temperatures. Hausmannite is more reduced and stable
at higher temperatures, so a hausmannite-rich protolith,
represented by B104, would retain most of its low-
temperature characteristics, with only minor amounts of
tephroite formng from admixed caryopilite and gageite.
The sodium in taneyamalite might stabilize nambulite,
saneroite, or serandite. We think that because CaO is
low in the Buckeye protoliths, the pyroxenoid might be

58 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

pyroxmangite, but rhodonite would be a more likely
product of metamorphosing santaclaraite-rich veins or
horizons.

With increasing metamorphic grade, the possible
assemblages that could form are severely limited by the
restricted bulk composition of the Buckeye ores and
other Mn-rich deposits. The variety of Mn minerals
(table 1) reported from Mn deposits metamorphosed to
greenschist and amphibolite grades reﬂects, in large
part, the availability of elements that are found in only
low amounts or are essentially absent at the Buckeye.

ORIGINS OF ANCIENT DEPOSITS

Of the deposits that might have an origin similar to
that of the Buckeye, few rival its diversity of Mn-rich
lithologies, and most others include lithologies that are
chemically distinct. To identify deposits formed in similar
environments, we must not ignore deposits composed of
fewer protoliths than at the Buckeye or deposits that
contain signiﬁcant additional chemical components, per-
haps reﬂected in new lithologies. Unfortunately, few of
our predecessors described the combination of geologic
setting and compositions of individual layers that will be
necessary to identify true analogues. In the following
paragraphs we will summarize critical features of some
possible analogues to the Buckeye deposit (and features
of some deposits that are not suitable analogues). The
analogues have a worldwide distribution. Like the Buck-
eye, many appear to have their origins in a chert-forming
marine environment.

The Smith prospect, Sierra Nevada Foothills, Nevada
County, Calif. (I.F. Wilson, in Trask, 1950, p. 167—168;
Hewett and others, 1961; Huebner, 1967; Flohr and
Huebner, 1988), lies in a belt of chert, slate, maﬁc
igneous rocks, and graded crystal tuffs of probable
marine origin. The deposit is a lens of rhodochrosite and
laminated hausmannite enveloped by massive white to
gray chert. Metamorphic minerals are rhodonite, teph-
roite, and small amounts of spessartine and sonolite;
their occurrence is controlled by local bulk composition.
At least some of the rhodonite is a reaction product
between carbonate and chert (Huebner, 1967). Interbed-
ded chert and metashales are nearby. Huebner noted the
occurrence of pyrite crystals and “colorless to brown
glass...incipiently recrystallized to a platey yellow sili-
cate with 7A spacing” (p. 75). The “glass” may corre-
spond to the gel-like material of chlorite composition
found at the Buckeye, and the pyrite may signify reduc-
tion of seawater sulfate to form sulﬁde. Spessartine
occurs with rhodonite and chert, but not hausmannite,
much as chlorite layers at the Buckeye avoid the more
oxidized assemblages. The area contains serpentine and
maﬁc volcanic rocks of zeolite- or lower greenschist-

 

facies grade. It is clear that tephroite and rhodonite can
form at very low metamorphic grades. We expect the
compositions of the protoliths at the Smith prospect to
have higher CaO contents but otherwise to resemble the
compositions reported for the Buckeye (table 6). Hueb-
ner (1967) advocated an origin for the Smith deposit that
involved thermal springs. We see no reason to discount
the connection between manganese deposition and a
hydrothermal ore-forming system driven by an igneous
heat source at the Smith prospect.

The setting of the Manga—Chrome mine is similar to
that of the Smith prospect, which lies three miles to the
south (I.F. Wilson, in Trask, 1950; Hewett and others,
1961; Flohr and Huebner, 1988). We examined a suite of
tephroite-bearing samples, one of which is shown by
Hewett and others (1961, their ﬁg. 1). These specimens
show sedimentary layers and laminations on a scale
similar to that of the Buckeye protoliths. The layers
consisted of rhodochrosite, the material called neotocite
by Hewett and others (1961), but which is probably
gel-like material of chlorite composition, and
hausmannite—rich metasediment. Some tephroite forms
layers parallel to bedding and may represent metamor—
phosed gageite-rich layers. We expect that the composi-
tion of these sediments would resemble those of the
Buckeye protoliths. A metamorphosed volcanic fragment
sits in laminated carbonate-rich mud, suggesting a close
spatial association between volcanism and deposition of
manganese-rich sediments.

Manganiferous metasediments at the Hoskins mine,
New South Wales, Australia (Ashley, 1986) lie between
ferruginous metachert (compositionally similar to a mix-
ture of chert and Al—poor nontronite) and overlying
metasiltstone and metasandstone. The metachert over-
lies metabasalt of MORB (midocean-ridge basalt) compo-
sition and probable greenschist-facies metamorphic
grade. This suite was deposited on the seaﬂoor of the
Cowra trough. Ashley compared the contorted and brec-
ciated laminations with those of jasper at the Blue Jay
and South Thomas mines, Franciscan Complex (Crerar
and others, 1982); we would extend the comparison to
deformed protoliths at the Buckeye deposit. The jasper
grades into massive manganese—rich metasediment that
now contains tephroite and rhodonite with minor haus-
mannite, Mn—rich chlorite, Mn-rich garnet, and relic
manganoan calcite and quartz. Texturally and composi—
tionally laminated Mn—rich sediment in which much Mn
and all Fe are trivalent may overlie the massive ore. The
minerals include braunite, manganoan pectolite, amphi-
bole (near kozulite in composition), an unnamed Mn+3
analogue of acmite, and an Mn-rich micaceous phase. The
compositions of Mn-rich metasediments at the Hoskins
mine are similar to compositions of the Buckeye proto—
liths but have slightly higher concentrations of A1203,

ORIGIN 59

Fe203, and 0210 and signiﬁcantly higher concentrations
of alkalis and the trace elements Ba, Co, and Sr. The two
Mn-rich metasediments at Hoskins can be distinguished
by the higher As and lower Co, Cu, Sr, and Zn in the
massive ore. Ashley (1986) proposed an origin in which
laminated silica and iron hydroxide gels were overlaid by
mounds of Mn—carbonate, Mn-oxides and hydroxides,
and silica that were deposited near vents. The possibility
of Mn—silicate protoliths was not discussed. Possible
precursors for the laminated ore, suggested by Ashley
solely on the basis of the chemistry of the layers, include
alkali feldspar or alkali-silicate gel; clays; oxides and
hydroxides incorporating Mn, Ba, and K; barite; and
Ca—Ba—Mn carbonates. Perhaps caryopilite should be
added to his list.

Deposits in the Ashio mountainland, Japan, best
reveal the geology of the bedded manganese deposits in
the Chichibu geosyncline, Japan (Watanabe and others,
1970). These deposits are stratiform and occur with
thin-bedded and massive cherts and basic volcanic rocks.
The amount of Mn ore correlates with that of associated
massive chert. The primary precipitates include
rhodochrosite, hausmannite, bementite, and “siliceous
material represented by colloidal silica” (Watanabe and
others, 1970, p. 139). This material may correspond to
the relic gel-like material found at the Buckeye. Regional
metamorphism resulted in braunite, tephroite, rhodo-
nite, spessartine, and piemontite. The conformably bed-
ded and mineralogically diverse deposits described by
Watanabe and others (1970) are more similar to the
Buckeye deposit than are the spherulitic rhodochrosite
nodules of purported diagenetic origin from the Inuyama
district, Japan, described by Matsumoto (1987).

Banded and massive manganese deposits are associ-
ated with chert-basalt contacts in the Apennine ophio-
lites (Bonatti and others, 1976). Base-metal deposits
occur in the basalts. The manganese deposits are con-
tained in layered chert, some of which is radiolarite, 20 to
30 m above the basalt. Limestone and calcareous shale,
but not graywacke, succeed the chert. Iron-rich sedi-
ments (Fe-oxides, ferruginous chert, or argillite) are
absent; the authors attributed this absence to the early
removal of iron (as sulﬁde) from a hydrothermal system.
Banded manganese deposits are stratiform and contain
layers of Mn minerals interbedded with red chert at a
thickness scale of 1 cm. Massive deposits the size of the
Buckeye lenses occur at the hinges of folds, are brecci-
ated, and grade laterally into banded deposits. Braunite
is the principal manganese mineral; it is commonly
accompanied by traces of rhodochrosite, parsettensite,
chlorite, rhodonite, and piemontite. As is the case with
the Buckeye ores, the bulk compositions of the Apennine
ores fall close to the join Mn+2Mng3Si012—Si02, with
minor Fe, Al, Mg, Ti, and alkali contained in nonmanga—

 

niferous silicates. The Mn/Fe and U/Th values are large
and the concentrations of Cu and Ni are small. The Co
concentration (11—600 ppm range), although higher than
in most deposits thought to have a hydrothermal origin,
is distinctly lower than in typical hydrogenous deposits.
Bonatti and others (1976) presented two REE patterns
that are puzzling even when the labeling of their table 7
is made to conform With their ﬁgure 6 and text. The
seawater-normalized pattern (ﬁg. 130) for a braunite—
rich sample is depleted like the Buckeye carbonates but
has a well-developed (and unique) negative Ce anomaly;
the mid-range is ﬂat. The chert is depleted overall.
Bonatti and others (1976) proposed that the Mn was
deposited as a sedimentary precipitate and that the
observed mineralogy developed during subsequent met-
amorphism. We think that the Apennine deposits dem-
onstrate several sedimentary precipitates with an origin
related to volcanism and hydrothermal processes. The
similarity of the Apennine and Buckeye deposits sug-
gests that the Buckeye may also have had a hydrother-
mal origin.

The Blue Jay (Trinity County) and South Thomas
(Mendocino County) deposits in the Franciscan Complex,
California, are deposits that may also be analogous to the
Buckeye deposit. Crerar and others (1982) described
massive Mn-rich lenses in thinly bedded chert that is
underlain by basalt. Crerar and others (1982) did not
report massive chert, but I.F. Wilson (in Trask, 1950, p.
138—139) noted the presence of massive white chert at
the South Thomas mine. At the Blue Jay mine, sandstone
occurs in contact with both chert and basalt (Crittenden,
in Trask, 1950, p. 308). Crerar and others (1982) referred
to the ore as massive but reported few results of a
thin-section examination of the ore. We have found that
even the most massive hand specimens of Buckeye
braunite and hausmannite are laminated when examined
as polished thin sections, and we suspect that the braun-
ite from the Blue Jay and South Thomas mines would
also prove to be laminated, were it so examined. Trask
and Trauger (in Trask, 1950, p. 309—310) mentioned only
hausmannite, carbonate, bementite, and native copper in
the ore from the Blue Jay mine. Crerar and others (1982)
identiﬁed braunite and minor rhodochrosite, hausmann-
ite, bementite, neotocite, rhodonite, inesite, pyrolusite,
todorokite, birnessite, nsutite, and vernadite, but they
found that much of the Mn-rich material was so poorly
crystallized that it could not be characterized. Trace-
element data were given by Crerar and others (1982) and
Chyi and others (1984). In many respects, the analyses
are comparable to those of an average Buckeye compo-
sition. The ores have high Mn/Fe and U/Th, and low
concentrations of Al and Ti, but Buckeye values for Zn
are higher and for Cu lower. Most REE patterns (ﬁg.

60 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

13D) show little overall enrichment and relatively ﬂat
slopes, similar to patterns for the Buckeye and the
modern hydrothermal materials, but it is puzzling that
the two most REE-enriched samples have the smallest
Ce anomaly. The REE patterns for most of the cherts
are similar to those of the most enriched cherts at the
Buckeye. Clearly, there are morphological, mineralogi—
cal, and chemical similarities among these three deposits.
As in the case of the Apennine deposits decribed by
Bonatti (1976), the similarity of Blue Jay, South Thomas,
and Buckeye deposits suggests processes associated with
seaﬂoor volcanism and hydrothermal activity.

Compared with Mn deposits in the Franciscan Com-
plex, stratabound Mn-rich lenses of Northland and Auck-
land, New Zealand, have similar lenticular shape but are
smaller and lack associated massive chert (Stanaway and
others, 1978, p. 24). Associated rock types are interbed-
ded cherts and argillites, marbles, submarine lavas and
tuffs, massive sulﬁde lenses, and ferromanganese depos-
its. The Mn—rich laminae and lenses consist of hausman-
nite, braunite, bementite, and manganite but not carbon—
ate. Stanaway and others (1978) invoked diagenetic
unmixing and recrystallization of sedimentary Mn-rich
clay precipitates under alkaline conditions to develop the
observed coarser textures and colloform structures. The
Mn/Fe values and trace—element concentrations for most
primary deposits in Auckland are similar to those for the
Buckeye deposits, but silica is lower and alkali higher.
Stanaway and others (1978) regarded the primary ferro-
manganese precipitates as having been amorphous and
now best represented by the cores of oolites. These New
Zealand deposits demonstrate a close association with
volcanic processes and textural diversity that increases
with diagenetic reprocessing. Three stages of marine
hydrothermal ore formation may be represented: early
sulﬁde formation, Fe precipitation, and late Mn precipi-
tation. The Buckeye deposit records only the stage of Mn
precipitation. Further, postdepositional processes such
as diagenesis do not appear to have been important at the
Buckeye.

Other stratiform and banded manganese formations
with mineralogy similar to that of the Buckeye lack the
close association with chert and have lateral dimensions
that greatly exceed those of the Buckeye lenses. The
Hotazel Formation in the Kalahari Basin, South Africa,
contains Mn-rich sediments (braunite, hausmannite,
bementite, traces of rhodochrosite) and banded iron-
formation (De Villiers, 1970; Kleyenstuber, 1984). Vol—
canic shards (in the Mn sediments), siliciﬁed rock, jasper,
and andesitic volcanic rocks, but not graywacke, are
present. Kleyenstuber advocated a volcanogenic—
sedimentary origin. Unlike the Buckeye, the Hotazel has
no interbedded chert-metashale sequences. The lateral
extent of individual manganiferous horizons (more than

 

35 km) exceeds that found in most modern marine
analogue environments (basins, rifts) or that of the
Ladd-Buckeye district. Thus the ore-forming process
must have been capable of delivering tremendous quan—
tities of Mn and dispersing it throughout an entire basin.

Banded greenschist-facies metasediments of Harlech,
Wales, form a 40-cm-thick horizon that covers an area
of 190 km2 and correlates with quartz-spessartine
metapelite horizons located 30 km to the west (Bennett,
1987). The Harlech Mn-rich horizon extends throughout
the area and consists of Mn-rich carbonate, spessartine,
chlorite, and silica. We think that the Harlech protoliths
could have been composed of Mn-carbonate, Mg—chlorite,
and quartz or of taneyamalite and Mg-chlorite (see ﬁg.
14). The mean trace-element concentrations (Mohr, 1956;
Bennett, 1987) are similar to those of modern hydrother-
mal precipitates (table 13). The REE patterns (Bennett,
1987), converted to our seawater-normalized scheme
(ﬁg. 130), resemble Buckeye patterns. Red laminated
ore shows the greatest REE enrichment, consistent with
the aluminosilicate fraction. The pattern of yellow con-
cretionary ore, now composed largely of carbonate,
spessartine, quartz, and chlorite, closely resembles the
pattern for the caryopilite-carbonate protolith at the
Buckeye (ﬁg. 78, B112). Caryopilite and carbonate are
reasonable protoliths for the observed mineral associa-
tion of the yellow layers at Harlech. Microspherulitic ore
is reported to be composed of carbonate, but the REE
patterns are more enriched than those of Buckeye
rhodochrosite. Perhaps this distinction can be attributed
to the opaque phase, presumably a Mn-oxide, that is
evident in Bennett’s photomicrograph (1987, his ﬁg. 7C).
The “average carbonate” (Bennett, 1987, p. 13) is not
described. Like Buckeye carbonate, it has low concen-
trations of REE. Like Buckeye gageite, the Harlech
carbonate is enriched in the heavy rare-earth elements
Yb and Lu. Perhaps Harlech carbonate had a protolith
consisting of rhodochrosite and gageite. As in the case of
the Hotazel deposits, the Mn-rich sediments of Harlech
have similarities to the Buckeye deposit and its modern
analogues, but the rhythmic cherts are lacking and the
areal extent is larger. Needed is a process to introduce
Mn and to disperse it over an area of 190 kmz. Bennett
(1987) proposed that this process was the precipitation of
oxides or, less probably, carbonates directly from a
hydrothermal brine. Were this brine to be denser than
cold seawater and to spread laterally to ﬁll a basin, as at
the Red Sea (see Bischoff, 1972, his ﬁg. 2), Mn could be
spread over an area the size of the Harlech or Kalahari
Mn-rich horizons.

We conclude this section by noting that there may be
more than one environment for the formation of proto—
liths that in some way resemble those found at the
Buckeye deposit. However, the themes of hydrothermal

ORIGIN 61

activity and volcanism are common to descriptions of
both modern seaﬂoor deposits and ancient deposits that
might, before diagenesis and (or) metamorphism, have
had similar protoliths. Lacking is a consistent set of
chemical and isotopic observations that can be uniformly
applied to all deposits. In many cases, we are not
conﬁdent that the reported mineralogy is complete or
correct. Most serious is the almost total lack of petro-
graphic data on internal structures that might help
distinguish primary sedimentary processes from diage-
netic and metamorphic processes. We urge future work-
ers to sample these deposits on the scale of the original
sedimentary units and to present petrographic data in
addition to chemical data.

ORIGIN OF BUCKEYE DEPOSIT

The foregoing descriptions provide clues to the nature
of the environment in which the Buckeye deposit formed
and t0 the processes of manganese concentration and
precipitation that might have operated in that environ—
ment. By extension, other manganiferous lenses in the
Franciscan Complex may have similar origins. We will
summarize the local environment, discuss the combina-
tion of processes that we believe was responsible for Mn
deposition, then move on to discuss the regional geologic
frameworks that might permit the operation of these
processes.

ENVIRONMENT NEAR BUCKEYE DEPOSIT

The association of conglomerate, graywacke and shale,
chert, and manganiferous units to form a thick unit (the
Grummett broken formation) appears to be caused by
primary sedimentary processes rather than secondary
tectonic juxtaposition. Radiolarian cherts are marine,
occur in orogenic belt sequences (Hein and Karl, 1983),
and are deposited in environments that are isolated from
coarse terrigenous elastic material such as that which
would form graywackes. Although it would be conve-
nient to call upon tectonic processes to juxtapose the
chert and graywacke, no ﬁeld evidence supports this
idea. The Buckeye deposit apparently formed in an
environment that permitted the alternation of chert-
forming and graywacke-forming processes.

There is evidence for a high-energy distant environ-
ment and a low-energy nearby environment. The occur-
rence of thick graywacke units coupled With the coarse
basal(?) conglomerate and intraformational shale-chip
conglomerate further indicate that one environment had
high potential energy, such as might be provided by a
continental slope, walls of a rift valley, or slopes leading

 

outward from a midocean ridge or into a trench. The
lateral continuity of the thin chert beds and interbedded
metashales, which are uniform over at least the scale of
outcrop, indicates the presence of a second environment
with 10W potential energy that had a gently sloping or ﬂat
seaﬂoor, perhaps in the form of a small, shallow basin.

The high-energy environment provided differentiated
plutonic components (alkali feldspar, quartz, volcanic
clasts), ultramaﬁc and maﬁc components (chromite, ch10-
rite), chert, and lithiﬁed but not recrystallized sediments
(shale chips). The source must have been distant and
exposed a cross section including recent sediments,
differentiated igneous rocks, and ultramaﬁc to maﬁc
basement. The ribbon cherts, representing the nearby
low-energy environment, contain two components. One
is clearly local in origin and consists of the debris of
siliceous pelagic organisms, Which accumulated on the
ocean ﬂoor and periodically ﬂowed into the site of
manganese deposition. The second component is repre-
sented by the ﬁne, claylike material that formed the
metashale interbeds and mixed with the accumulating
debris of organisms to form the included “dust” and thin
partings in the cherts. The similarity between the aver—
age TiOZ, A1203, and alkali contents of seven metashales
from the Ladd—Buckeye district (the argillites of Hein
and others, 1987) and averages for metashales (Petti-
john, 1957) and graywacke (Pettijohn, 1963) suggests
that the source material for the metashales was detrital,
perhaps ﬁne detritus related to the graywacke. This is
conﬁrmed by the REE patterns, which mimic those of
the graywackes (ﬁg. 7A). Deposition of the metashales
was probably continuous, forming a background that was
periodically overwhelmed (diluted) by a turbidity ﬂow or
deposition of a manganese lens.

The combined lack of current structures, algal struc—
tures, bioturbation, and calcareous fossils together with
the preservation of delicate laminae suggests deposition
in deep water, consistent with the topography necessary
to accelerate the turbidites. The depth of the seaﬂoor is
not known; the absence of calcareous skeletal debris and
the relatively low concentrations of CaO in the ores and
surrounding sediments suggests deposition below the
carbonate compensation depth (or, conversely, suggests
that there were no pelagic calcium-secreting organisms).
The abundance of radiolarian chert suggests oceanic
upwelling that enriches surface waters in nutrients (Hein
and Parrish, 1987). Because even hot solutions vented at
the seaﬂoor have been observed to move laterally rather
than reach the surface of the ocean (Klinkhammer and
others, 1986; Baker and Massoth, 1986), it seems unlikely
that there would be a direct connection between ﬂuids
vented at the ocean ﬂoor and a pelagic bloom near the
surface.

62 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

MANGANESE DEPOSITION AT BUCKEYE DEPOSIT

Traditional interpretation of the origin of deposits such
as the Buckeye involved inorganic hydrothermal pro-
cesses that culminated in thermal springs and precipi-
tates on the ocean ﬂoor. Recently Hein and others (1987)
and Hein and Koski (1987) invoked organic diagenesis at
depth that culminated in the replacement of chert by
manganese carbonate within the sediment pile. This
model provides a plausible explanation for the source
(quadrivalent Mn oxyhydroxides in the sediment column)
and solubilization (reduction to Mn+2 state) of the man-
ganese. It also provides a mechanism (methanogenesis
and subsequent oxidation to COZ) for generating carbon—
ate that is strongly depleted in 13C. An origin involving
only diagenesis and replacement, however, is not com—
patible with all features at the Buckeye deposit. Our
principal objections to a model based on diagenesis and
replacement are the following.

(1) Despite the large number and diversity of samples
examined, we observed no islands of relic (unreplaced)
chert. (2) There is evidence that Mn was removed from,
not introduced into, the chert at the margin of the
orebody. (3) We observed no cases in which the orebody
transected the layering of the surrounding chert. (4) The
scale at which most of the ore is laminated is ﬁner than
the scale of layering within the chert and metashale. (5)
There is a large halo of Mn enrichment, rather than Mn
depletion, surrounding the orebody. The MnO content
decreases from an average of about 50 weight percent in
the orebody to about 5 weight percent in the adjacent
wallrocks drilled by Volin and Matson (1949). Average
values for different lithologies in the series chert-
metashale from the Ladd-Buckeye district range from
0.21 to 1.29 weight percent MnO and 1.20 to 8.57 weight
percent total iron as Fe203 (Hein and others, 1987).
Compared with broader scale averages (Pettijohn, 1957;
Cressman, 1962), these materials appear enriched, not
depleted, in MnO. Additional analyses (Huebner and
Flohr, unpub. data) of chert and shale from the vicinity of
the Buckeye deposit conﬁrm the MnO enrichment. The
MnO concentrations of seven graywackes range from
0.02 to 0.74 weight percent and thus, overall, appear
enriched rather than depleted in MnO (compare with
Pettijohn, 1957, 1963; Taylor and McLennan, 1985). (6)
According to the model, silica replaced by rhodochrosite
contributed to the formation of massive chert (Hein and
others, 1987, p. 220). However, the so-called massive
chert is thinly layered (bedded). The displaced silica has
not been found. (7) The model does not provide a
mechanism to form layers of six to nine distinct compo-
sitions in the orebody. Further, the massive chert and
interbedded chert and metashale lack the lithologic (com-
positional or textural) diversity of the orebody, making it
unlikely that the observed layering in the orebody

 

reﬂects layering of some precursor. (8) Gel-like materials
occur in the orebody, yet we would not expect such
ﬁne-grained materials to form during the replacement of
chert and shale or to survive burial diagenesis in a wet,
warm environment. (9) The spherulitic and botryoidal
textures ascribed by others to precipitation of rhodo-
chrosite during diagenesis are rare at the Buckeye and,
where they do occur, clearly replace earlier, very ﬁne
grained rhodochrosite. (10) The Mn protoliths, including
the Mn-silicates, have depleted REE patterns, relative
to those of the chert and metashale (ﬁg. 7). Claylike
silicates formed by replacement inherit the trace-
element signatures of their precursors (Mosser, 1979;
Mosser and others, 1979). If the Mn protoliths, including
the silicates, formed by replacement, we would expect
less-depleted REE patterns. (11) The orebody and sur-
rounding graywackes are sulﬁde poor, yet the suggested
diagenetic reactions would have brought together dis-
solved Fe+2 and HZS' to precipitate FeSZ. (12) Diagen-
esis of organic matter produces methane only after all
ferric iron is reduced to soluble Fe”. Thus, in the
absence of a sink for iron (precipitation as sulﬁdes), the
model fails to explain the high Mn/Fe observed in the
orebody. (13) The model calls for zones of Mn reduction
and oxidation, separated by a zone of diffusion. Such a
process should result in a single Mn-rich layer or single
set of Mn-rich layers, one of each protolith. It is difﬁcult
to imagine a diffusion process at depth that will produce
cyclically alternating compositions, such as the carbonate
and caryopilite observed at the Buckeye. Were the oxic
zone raised, the zone of Mn precipitates should be
dissolved and reprecipitated at a higher level. There is no
petrographic evidence for such a shift in position. (14) It
is unlikely that diffusion alone would be the mechanism
for scavenging Mn from the surrounding volume of chert,
metashale, and graywacke. Considering the distances
involved (the zone of Mn enrichment extends hundreds of
meters), it is probable that moving ﬂuids played an
important part in transporting the manganese. (15) The
diagenetic model accounts for the absence of modern
deposits like the Buckeye by having them form beneath
the seaﬂoor but does not explain the absence of modern
equivalents of the associated interbedded cherts and
metashales, which must have formed on the ocean ﬂoor.
Until the modern environment for deposition of interbed-
ded chert and shale is found, the modern environment for
deposition of a Buckeye deposit cannot have been found.

Our model, based on a cool hydrothermal system,
combines diagenesis to dissolve the manganese incorpo-
rated in the sediment column, ﬂuid ﬂow to transport the
Mn, and precipitation at the seaﬂoor to develop the ﬁne
laminations and diversity of compositions. These pro-
cesses will be discussed in the sequence in which they
interact with ﬂuid in the sediment pile.

ORIGIN 63

Oxic and mildly suboxic diagenesis occurs near the top
of the sediment column. This process consumes buried
organic matter using, as an oxidant, oxygen that is
dissolved in the pore ﬂuid, that is bound to quadrivalent
Mn, or that diffuses (as molecular 02) downward from
the water column. The Mn+2 produced is dissolved in the
pore ﬂuid. Some CO2 dissolves to form H2003 and
HCOQ. At the same time, the ﬂuid equilibrates with
opaline skeletal debris, taking some silica into solution.
Deeper in the sediment column, intensely suboxic and
anoxic conditions occur, but any Fe”, HZS_, and CH4
produced under these conditions are precipitated in place
or, if they migrate upward, oxidized and perhaps precip-
itated before they reach the upper levels. The oxidants
are carried by ﬂuids that pass through the sediment
column. Iron is precipitated as oxyhydroxide (or sulﬁde),
HS‘ is converted back to SOQZ, and CH4 is oxidized to
CO2 with negative 8130 before reaching the mildly sub-
oxic and oxic zones. Some 002 and CH4 may move as
vapor bubbles. An important element of the diagenetic
portion of the model is that any anoxic ﬂuids are neutral-
ized (made mildly suboxic) by interaction with oxic to
suboxic ﬂuid—either slightly modiﬁed seawater that
circulates beneath the seaﬂoor or pore ﬂuids expelled by
compaction from deeper and less organic-rich (or less
reactive) sediments. The results are a supply of aqueous
ﬂuid carrying dissolved species of Mn, 002, and SiOz, the
absence of sulﬁdes, and Fe-poor precipitates, as
observed at the Buckeye deposit.

The source of the manganese is enigmatic. Maﬁc
volcanic rocks, which appear in models for seaﬂoor
mounds, are not part of the stratigraphic column at the
Buckeye, although such volcanic rocks are spatially
associated with many other manganese deposits in the
Franciscan Complex. The most likely source of manga-
nese is the sediment column, including the interbedded
chert—metashale, as in the diagenesis model of Hein and
Koski (1987), or the graywacke, from which manganese
could be leached and transported in the divalent state by
relatively reducing solutions. The source sediment could
lie below the deposit or be laterally displaced from it. The
source might even be a lithology that is not represented
in the Grummett broken formation.

Dissolved constituents, and perhaps included vapor
bubbles, were transported by ﬂuids that ﬂowed within
the sediment column and vented at the seaﬂoor. The
forces driving the ﬂuids were heat and (or) pressure.
Heat would cause convection that draws fresh, oxygen-
ated seawater into the sediment column; favored pas-
sageways would be channels that open to the seaﬂoor. As
a result of sediment compaction, pressure would expel
ﬂuids into channels and upward toward the seaﬂoor. In a
tectonically active environment, such as that in which the
Buckeye formed, semiconsolidated and consolidated sed-

 

iments would be expected to have been fractured and to
inﬂuence ﬂuid ﬂow at higher levels, much as fractures in
sediment-covered basalts are thought to inﬂuence the
distribution of mounds and vents near midocean ridges.
Fractures, ranging from the quartz- and carbonate-
bearing shear zones that separate broken formations
(ﬁg. 13) to veins that cross Mn-rich sediments but do not
deposit Mn minerals (ﬁgs. 4A,E,J) are evidence that
moving ﬂuids were channeled. In one case (ﬁg. 4A, B87),
hausmannite-rich layers became more carbonate rich
near the vein, suggesting that the vein ﬂuid had lower
redox potential than the original sediment and was
COz—rich. Fluid ﬂow permitted scavenging of Mn from
distances greater than diffusion ﬁelds, separation of the
anoxic region of methane generation from that of ore
precipitation, and rapid delivery of ore-forming constit—
uents. The ephemeral nature of vein systems, which
open and close in response to tectonic disturbances,
subsidence, and precipitation, provides a means for
quickly turning the supply of ore—forming constituents on
and off, leading to groups of orebodies, each with sharp
boundaries against the enclosing cherts.

Mixing of ﬂuid from within the sediment column with
fresh seawater and subsequent precipitation on the sea-
ﬂoor satisﬁes many of the observed constraints. There
will be no residues—as phases or chemical signa-
tures—of a replaced lithology. Because the composition
of the mixedﬂuid will vary with the composition of the
subsurface ﬂuid, the ﬂow rate of the vented ﬂuid, and the
degree of mixing with fresh seawater, the nature of the
precipitated layers may change with time or distance
from the vent and may show cyclical variation, and the
layers may be thick or thin. The model permits rapid
delivery of ore-forming materials, necessary to avoid
contamination (terrestrial, hydrogenous, or diagenetic),
and rapid precipitation from the ﬂuid mixture, necessary
to localize the deposit.

Changes in the ﬂuxes supplying chemical components
to the depositional site varied and caused compositional
layering within the orebody. These variations must have
been rapid, perhaps on the scale of seasonal climatic
variations that, through winds, affected the upwelling of
nutrients and that have been shown to affect the rates at
which diatom tests are supplied to the seaﬂoor (Honjo,
1982; Takahashi, 1986). Fluids from the sediment column
probably provided ﬂuxes of dissolved Mn, SiOz, 002, and
minor to trace amounts of the “hydrothermal” elements
such as Zn. Seawater probably contributed ﬂuxes of 02,
MgO, and minor and trace amounts of Ba, Mo, V, and
804. The composition of the seawater itself may have
varied, particularly with respect to dissolved 02, which
could vary seasonally. Thus, the valence of Mn in indi-
vidual protoliths, and the Ge anomaly, could depend not

64 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

only on the ratio of ﬂuids in the mixture but upon the
composition of the seawater component.

DIAGENESIS AND METAMORPHISM

The relatively oxidized sediments (Mn+3 is present)
were preserved rather than subsequently modiﬁed by
diagenesis. Had the Mn—rich precipitates experienced
suboxic diagenesis after burial, braunite, and perhaps
hausmannite, would not have survived. We presume that
little organic matter was buried with the Mn—rich and
surrounding sediments or, if present, that it was isolated
from the deposit.

Our observations provide seemingly contradictory evi-
dence for the existence of a fluid phase during diagenesis
and metamorphism. By analogy with modern marine
sediments, we presume the initial graywacke sands,
cherts, shales, and Mn—rich layers were very water rich.
Veins of quartz, aragonite, and calcite clearly mark ﬂuid
pathways. However, the non-uniform distribution of
jadeite within the Grummett broken formation, the
presence of relic gel-like materials within the orebody,
and the preservation of oxidized Mn (in braunite) suggest
the absence of a ﬂuid phase that would promote recrys-
tallization or regional redox equilibrium. To reconcile
these observations, we propose that compaction during
burial and subsidence expelled most of the trapped pore
ﬂuids throughout the section, leaving relatively dry (and
thus unreactive) materials, such as chert. Subsequent
ﬂuid loss at metamorphic pressures was concentrated
along the contacts between the broken formations, at
some distance from the Buckeye orebody. These contacts
were kept open to ﬂuid loss by shearing, which appears
to have been sufﬁciently intense, in some cases, to
transport dense blocks of glaucophane-bearing schists.
Fluid in these shear zones may ultimately have reached
seaﬂoor vents. Higher metamorphic temperatures at
greater depths may have caused decomposition of the
clay- and chlorite-rich matrix of the metagraywackes,
providing an additional source of ﬂuids.

The Mn—rich mineralogy at the Buckeye provides a
clear record of high oxygen-fugacity values during meta—
morphism. The occurrences of braunite and hausmannite
clearly record an intrinsic oxygen fugacity that is orders
of magnitude greater than that of the surrounding gray-
wackes (Huebner, 1967, 1976), which contain ferrous
silicates (table 5) and, probably, organic matter. Because
this f02 is so high, it cannot have been brought about by
metamorphic equilibration with surrounding rocks and
must reﬂect the premetamorphic history of the sedimen-
tary environment (Huebner, 1967). Huebner (1969, p.
476) suggested that there might be oxidation halos
surrounding such highly oxidized orebodies. In the case
of the Buckeye, the surrounding interbedded chert and

 

metashale are highly oxidized, but this state is probably
a primary rather than a secondary feature. It is more
difﬁcult to read the metamorphic temperatures and
pressures. One difﬁculty is the absence, in the Buckeye
deposit, of mineral associations that can be proven to be
equilibrium assemblages. By analogy with howieite,
which is thought to be stable only at high pressure and
low temperature (Lattard and Schreyer, 1981), the
occurrence of the mineral taneyamalite might indicate
high pressures of formation. Calculations by Huebner
(1967) indicated that the partial pressure of CO2 at the
stage recorded by the mineralogy was an order of
magnitude less than the total pressure estimated from
the mineral assemblages in the metagraywacke. Either
no ﬂuid phase was present or the ﬂuid was dilute with
respect to C02. At such high f02 values and equilibrium,
CO, CH4, and H2 would not be abundant species. There-
fore, any ﬂuid present during metamorphism must have
been aqueous. But preservation of relic gel-like materials
suggests that an aqueous phase was not present during
metamorphism and therefore that no ﬂuid was present.

REGIONAL FRAMEWORK

The spatial and temporal relations between the sites of
Mn deposition and deep burial (blueschist-facies meta—
morphism) are enigmatic. Convecting ﬂuid, even cool
water, requires a heat source, yet rapid subsidence that
generates 5—8 kbar pressure without exceeding 150—200
°C requires a small regional thermal gradient. If the
Buckeye Mn lenses were precipitated from ﬂuids driven
by a heat source, it must have been highly localized,
decayed before subduction, or moved relative to the
lenses. However, if manganese and 002 were dissolved
in, transported by, and precipitated from ﬂuids expelled
by compaction of the sediment pile, it is conceivable that
no heat source would be necessary.

POSSIBLE ENVIRONMENTS

Mn-rich mounds have been observed to form in oceanic
island arcs and oceanic spreading centers, where they
are associated with high heat ﬂow and nearby hot or
warm ﬂuids. If cooler ﬂuids can leach manganese and
iron from the sediment column, Mn—rich mounds might
occur in two additional environments. One is the abyssal
ﬂoor at great distance from the midocean ridge. Here,
hemipelagic sediments should be underlain by maﬁc and
ultramaﬁc ocean crust. Were serpentinization and other
alteration processes not completed in the crust—forming
ridge environment, they would continue, albeit at much
lower temperatures, as the crust moved away from the
ridge, releasing sufﬁcient heat to drive ﬂuid convection

ORIGIN 65

and causing an increase in rock volume (Fyfe, 1974). This
volume increase might fracture the overlying sediments,
thereby channelling ﬂuids. In the geologic record, these
deposits would be found associated with hemipelagic
sediments but, unless subsequently juxtaposed by tec-
tonics, probably not with graywacke turbidites. Another
environment is the top of a subduction complex. Water-
rich sediments in the trench (or surﬁcial expression of
the subduction complex) would release cool pore water as
they were subducted. The boundaries between slabs of
sediments would serve as channels for escaping ﬂuids.
Vents or chains of vents might occur on the walls of the
trench. If cool ﬂuids are able to leach manganese from
the sediments, mounded deposits of Mn-rich precipitates
might form on the trench walls.

In which paleoenvironment were the Buckeye deposit
and related sediments of the Grummett broken forma-
tion deposited? Raymond (1974, p. 145) originally pro—
posed that the Franciscan rocks of the Mount Oso area
were deposited in an environment characterized by “a
combined trench—abyssal ocean-ﬂoor plus ocean ridge.”
Subsequently (Raymond, 1977) he associated the manga-
niferous sediments with seaﬂoor spreading centers. Sny-
der (1978) reviewed the literature on manganese de-
posits associated with chert-greenstone complexes,
including deposits in the Franciscan Complex, and chose
environments in which oceanic hot springs could form:
oceanic spreading centers, marginal oceans, and the base
of island arcs and oceanic islands. Crerar and others
(1982) favored an oceanic spreading center over island
arcs, back—arc spreading centers, and convergent plate
margins. Hein and others (1987) interpreted the Fran-
ciscan sediments of the Ladd-Buckeye district as having
been deposited near a continental margin, then made
reference to possible fore-arc, back-arc, and rifted con-
tinental margin settings. Subsequently, Hein and Koski
(1987) advocated a deep, rapidly deposited continental
margin facies, such as the Yolla Bolly terrane of northern
California described by Blake and others (1982). Ray-
mond (1988) disagreed with the analogy because it refers
to chert—poor lithologies, whereas the Grummett broken
formation is chert rich.

To choose the most appropriate environment, we need
ﬁrst to review the facts that serve as boundary condi-
tions for the range of possible environments. (1) The
cobble conglomerate exposed near the base of the Grum-
mett unit must have been derived from an environment
with substantial topographic relief and must have been
deposited relatively near its source. (2) There must have
been three or four sources for the elastic sediments and
pathways by Which detritus could move from each source
to the depositional environment. (3) These sources sup—
plied quartz-rich but lithic-fragment—poor sands. (4) The
boundaries between conglomerate, metagraywacke,

 

interbedded chert and metashale, and the manganese
deposit are primary sedimentary contacts. Thus, the
depositional environment must be capable of generating
these four basic lithologies in a continuous stratigraphic
section. (5) The chert layers are turbidites that were
deposited in an environment that was intermittently
exposed to a supply of coarse detritus. (6) The manga-
nese deposit has the size, morphology, and trace-element
composition of a deposit that is genetically associated
with a seaﬂoor vent. (7) The compositional layering of the
orebody signiﬁes an environment in which the deposi-
tional parameters changed frequently. (8) Mn deposits
such as the Buckeye are invariably associated with chert
and shale. Mn deposits are not enclosed in graywacke. (9)
Some Mn deposits, unlike the Buckeye, are associated
with basalt. (10) The metagraywacke and enclosed sedi-
ments at the Buckeye were together subjected to the
high pressures and low temperatures characteristic of
subduction complexes at continental margins. Given
these boundary conditions, we can eliminate two envi-
ronments.

The abyssal ﬂoor model is unsuitable. Unlike the
environment in which the Buckeye formed, the abyssal
environment has low topographic relief and lies far from
possible sources of the cobble conglomerate and quartz-
rich graywacke. Although possible sources of heat are
mantle hot spots or serpentinization of heretofore unal-
tered oceanic crust, serpentinization is unlikely. Were
this process to operate, ﬂuids convecting through the
sediment pile must have supplied the H20 necessary for
serpentinization and, by association, become alkaline
(high pH). Such ﬂuids would be too alkaline to leach and
transport manganese and iron (Barnes and O’Neil, 1969).
Thus, even though the ocean crust With its mantle of
abyssal sediments might be carried into a subduction
complex, the crust is unlikely to arrive with manganese
deposits such as the Buckeye.

Deposition of manganese within the trench is an
appealing hypothesis because it directly involves the
subduction complex and provides a mechanism for bring-
ing together sediments of diverse origins. We postulate
two variations on the model associated with trenches. In
the first, the ﬂuid-expulsion model, compaction Within
the pile forces Mn-bearing ﬂuids upward along shear
zones between subducted slices of terrane. Because the
dominant lithology is graywacke and the observed shear
zones are in graywacke, we would expect at least some
deposition of Mn-rich lenses on graywacke sands. The
fact that the Mn-lenses are invariably associated With
chert, a minor lithology, contraindicates the ﬁrst mech—
anism. Further, the model also fails to explain the
presence of submarine basalts, which underlie the cherts
that are associated with many Mn lenses in the Fran-
ciscan Complex and elsewhere. If the Buckeye formed by

66 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

the same set of processes that formed the Mn deposits
that are closely associated With basalt, the model in
Which ﬂuids were expelled by compaction is also inap-
propriate for the Buckeye deposit.

The second trench—related hypothesis involves the
subduction of a midocean ridge at a continental margin.
A possible modern example is the collision of the Chile
Rise with the Peru—Chile trench (see Forsythe and
others, 1986). In this model, the continental margin is a
region of oceanic upwelling that brings nutrient-rich
waters to the surface and causes high organic productiv-
ity (radiolarians). As the ridge (with its grabenlike
topography) approaches the trench, chert and shale are
deposited in protected basins, either directly on ridge
basalt or on graywacke. Graywacke sands and conglom-
erate, derived from dissected, quartz—rich continental
crust, ﬂow into the trench and periodically spill into
protected basins on the oceanic side of the trench axis,
blanketing all other lithologies. Some coarse conglomer-
ates and graywackes are locally derived from trench
walls that oversteepen and collapse. The distal portions
of both locally derived and distant ﬂows provide a
continuous supply of ﬁne detritus that settles slowly to
form shale. Shale deposition is frequently overwhelmed
by ﬂows of sands derived from the sites where the debris
of radiolarians settles. Thus, the scene is set for Mn
deposition on chert that can be both closely associated
with, and distant from, basalt. Residual heat from the
decaying ridge or from short-lived, near-trench thermal
events (predicted by DeLong and Fox, 1978) drives ﬂuid
circulation that leaches, transports, and deposits Mn by
the processes previously discussed. The residual heat is,
in most cases, insufﬁcient for continued basaltic magma-
tism. Some ﬂuid could also be driven by expulsion from
the sediment pile. Subduction buries the Mn deposit and
may tectonically dissociate the Mn mounds or lenses (and
their host rocks) from oceanic basement and crust formed
at the ridge.

To be acceptable for the Buckeye deposit, the ridge-
trench model must be compatible with the evolution of
what is now western California. Schweikert and Cowan
(1975), Dickinson and others (1982), Ingersoll (1982), and
Seiders and Blome (1988), who drew upon numerous
earlier studies, provide useful summaries. In Middle
Jurassic time, there existed a west-facing continental-
margin arc and, to the west, an east-facing oceanic arc.
Farther to the west, there must have been a spreading
center (source of oceanic crust) such as the back-arc
spreading center of Ingersoll (1982). By Late Jurassic to
Early Cretaceous time (the dates of microfossils within
the Grummett broken formation), the oceanic crust
between the two arcs had been subducted, leaving
the single westward-facing Sierran-Klamath magmatic
arc complex, trench, and remnant of the spreading

 

 

{J .7
.w

TSC FAB

SL OSC

 

 

cc
DMA €72
cc

SL TSC

 

    

MiddIe Cretaceous

FIGURE 15.—Possible mechanism for the creation of the paleo-
depositional environment for the Buckeye protoliths, adapted from
Dickinson and others (1982) and Ingersoll (1982) with embellish-
ments. Drawn not to scale but to emphasize features of the deposi-
tional environment. During the Late Jurassic, an oceanic spreading
center (OSC) with median valley containing restricted basins
approached the continental margin, a source of quartz-bearing sands.
The spreading center was consumed by subduction by Early Creta-
ceous time. If the spreading center was still warm as it approached
the trench and subduction complex (TSC), hydrothermal Mn—rich
deposits (Mn, black), and hemipelagic sediments (HPS, cherts and
shales) could form near the continental margin, perhaps in locally
sheltered basins or along remnants of the median valley (MV).
Turbidity flows of sand, originally derived from the dissected
Sierran-Klamath magmatic arc (DMA), ﬂowed longitudinally north-
ward along the axis of the trench and fanned outward, forming
massive graywackes (stippled). Episodically, major sand-laden cur--
rents overﬂowed the restricted basins, permitting the stratal inter-
layering of hemipelagic sediments and graywackes. Other abbrevia-
tions: SL (sea level); GVS (Great Valley sequence) deposited on CRO
(Coast Range ophiolite); OC and CC (oceanic and continental crust,
respectively); FAB (forearc basin); SC (subduction complex).

center (ﬁg. 15). The Great Valley sequence was depos-
ited in the forearc basin immediately to the west of the
eroded are complex. Farther westward, Franciscan sed-
iments ﬂowed both westward and longitudinally from the
south. These sediments were deposited along (and
fanned westward from) the break in slope formed Where
the eastern trench slope and the oceanic crust met. At
the same time, oceanic crust east of the remnant spread-
ing center was consumed by the trough-subduction

REFERENCES 67

complex, bringing the spreading center toward the
trench. If the spreading center was still warm as it
approached the trench, or if ﬂuids were expelled from
the sediment pile, the scene was set for operation of the
ridge-trench model as outlined above. Manganese depos-
its and cherts, some but not all associated with basalt,
formed. Subsidence and deep burial deformed the sedi-
ments, expelled ﬂuids, and caused the blueschist—facies
overprint.

SUMMARY AND FUTURE WORK

Textural, mineralogical, microchemical, and trace-
element data collected from similar (and spatially
related) metasediments from the Buckeye deposit sug-
gest that manganese carbonate, silicates, and oxides
were deposited at the seaﬂoor from dilute mixtures of
hydrothermal ﬂuid with seawater. Although similar
deposits have been described in the literature, our data
focus on well-characterized individual sedimentary lay-
ers rather than on individual mineral grains within a
layer or on groups of layers. The chemistries of the
individual layers and laminations record ﬂuctuations in
the marine environment, caused in part by periodic ﬂows
of radiolarian sand. At the Buckeye, the contrasting
layer compositions record eight sedimentary compo-
nents, which must contain valuable clues to the geochem-
ical and sedimentological variables in both ancient and
active marine environments. N o doubt other manganese
deposits formed in a similar environment, but to our
knowledge the Buckeye is the only described deposit that
reveals more than one Mn-silicate protolith. The appar—
ent uniqueness of the Buckeye is probably due more to
the inadequacies of available descriptions than to any
singularity in the geologic record. It is also possible that
similar deposits are forming on the seaﬂoor today.

We need to be certain that the reasoning leading to the
characterization of Mn-forming environments and proc-
esses does not become circular. Especially frail is the
path of reasoning (1) chert and layered manganese
deposits that are closely associated with basalt are
volcanogenic, (2) therefore most chert and layered man-
ganese deposits, even those not associated with basalt,
have a volcanogenic origin, and (3) a particular layered
manganese deposit formed by volcanogenic processes.
The discovery and characterization of actively forming,
laterally restricted volcanogenic-hydrothermal Mn-oxide
deposits has done much to strengthen the reasoning, but
interbedded manganese oxides, silicates, and carbonates
have not yet been found in modern environments, and
laterally extensive deposits occur in the geologic record.
Until we know why these ancient layered deposits differ
from the actively forming deposits, we cannot be certain
that the depositional environments are similar.

 

Detailed studies of the Buckeye deposit provide a basis
for two research directions, now being undertaken. (1)
We are attempting to characterize the chemical signa-
ture of each sedimentary component in order to verify
the sources of each of the chemical constituents and to
discover the nature and causes of the variations in the
seaﬂoor environment. In particular, we want to develop
criteria that will enable us to distinguish between maﬁc
igneous rocks and the sediment column as sources of
manganese, iron, and base metals, and we want to make
ﬁrm our hypothesis concerning the role of biogenic silica.
We hope that the results of this effort will help us
understand the origin of laterally extensive stratiform
manganiferous sheets that are not associated with chert.
(2) We are using petrographic and microchemical crite-
ria, developed with reference to the Buckeye deposit, to
determine the protoliths of deposits that have lost their
obvious primary characteristics due to pervasive meta-
morphism. If correctly read, such deposits may be valu-
able paleoenvironmental indicators in high—grade meta-
morphic terranes.

REFERENCES

Adachi, Mamoru, Yamamoto, Koshi, and Sugisaki, Ryuichi, 1986,
Hydrothermal chert and associated siliceous rocks from the north-
ern Paciﬁc: Their geological signiﬁcance as indication of ocean
ridge activity: Sedimentary Geology, v. 47, p. 125—148.

Albee, AL, and Ray, L., 1970, Correction factors for electron probe
microanalysis of silicates, oxides, carbonates, phosphates, and
sulfates: Analytical Chemistry, v. 42, p. 1408—1414.

Alt, J .C., 1988, Hydrothermal oxide and nontronite deposits on sea-
mounts in the eastern Paciﬁc: Marine Geology, v. 81, p. 227—239.

Ashley, P.M., 1986, An unusual manganese silicate occurrence at the
Hoskins mine, Grenfell district, New South Wales: Australian
Journal of Earth Sciences, V. 33, p. 443—456.

Audley-Charles, M.G., 1965, A geochemical study of Cretaceous fer-
romanganiferous sedimentary rocks from Timor: Geochimica et
Cosmochimica Acta, v. 29, p. 1153—1173.

Baedecker, P.A., ed., 1987, Methods for geochemical analysis: US.
Geological Survey Bulletin 1770, 132 p.

Bailey, E.H., Irwin, W.P., and Jones, D.L., 1964, Franciscan and
related rocks and their signiﬁcance in the geology of western
California: California Division of Mines and Geology Bulletin 183,
177 p.

Bailey, S.W., 1980, Structures of layer silicates, in Brindley, G.W.,
and Brown, G., eds., Crystal structures of clay minerals and their
X-ray identiﬁcation: Mineralogical Society [London] Monograph 5,
p. 1—123.

Baker, E.T., and Massoth, G.J., 1986, Hydrothermal plume measure-
ments: A regional perspective: Science, v. 234, p. 980—982.

Barnes, Ivan, and O’Neil, J .R., 1969, The relationship between ﬂuids
in some fresh alpine-type ultramaﬁcs and possible modern serpen-
tinization, western United States: Geological Society of America
Bulletin, v. 80, p. 1947—1960.

Baudracco-Gritti, Celestine, 1985, Substitution du manganese bivalent
par du calcium dans les minéraux du groupe: braunite, neltnerite,

68 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

braunite II [Substitution of bivalent manganese by calcium in
minerals of the braunite-neltnerite-braunite 11 group]: Bulletin de
Minéralogie, v. 108, p. 437—445.

Bence, A.E., and Albee, AL, 1968, Empirical correction factors for
the electron microanalysis of silicates and oxides: Journal of
Geology, v. 76, p. 382—403.

Bennett, M.A., 1987, Genesis and diagenesis of the Cambrian manga-
nese deposits, Harlech, North Wales: Geological Journal, v. 22,
p. 7—18.

Bischoff, J .L., 1969, Red Sea geothermal brine deposits: Their miner-
alogy, chemistry, and genesis, in Degens, E.T., and Ross, D.A.,
eds., Hot brines and recent heavy metal deposits in the Red Sea:
New York, Springer Verlag, p. 368—401.

1972, A ferroan nontronite from the Red Sea geothermal
system: Clays and Clay Minerals, v. 20, p. 217—223.

Blake, M.C., Jr., Jayko, A.S., and Howell, D.G., 1982, Sedimentation,
metamorphism, and tectonic accretion of the Franciscan assem—
blage of northern California, in Leggett, J .K., ed., Trench-forearc
geology: Sedimentation and tectonics on modern and ancient active
plate margins: Oxford, Blackwell Scientiﬁc Publications,
p. 433—438.

Bonatti, Enrico, Zerbi, Marco, Kay, Robert, and Rydell, Harold, 1976,
Metalliferous deposits from the Apennine ophiolites: Mesozoic
equivalents of modern deposits from oceanic spreading centers:
Geological Society of America Bulletin, v. 87, p. 83-94.

Bowers, T.S., Von Damm, K.L., and Edmond, J .M., 1985, Chemical
evolution of mid-ocean ridge hot springs: Geochimica et Cosmo-
chimica Acta, v. 49, p. 2239—2252.

Boyce, A.J., Coleman, M.L., and Russell, M.J., 1983, Formation of
fossil hydrothermal chimneys and mounds from Silvermines, Ire-
land: Nature, v. 306, p. 545-550.

Brett, Robin, Evans, H.T., Jr., Gibson, E.K., Jr., Hedenquist, J .W.,
Wandless, M.-V., and Sommer, M.A., 1987, Mineralogical studies
of sulﬁde samples and volatile concentrations of basalt glasses from
the southern Juan de Fuca Ridge: Journal of Geophysical
Research, v. 92, p. 11373—11379.

Bruland, K.W., 1983, Trace elements in sea-water, in Riley, J .P., and
Chester, R., eds., Chemical oceanography: London, Academic
Press, v. 8, p. 158—220.

Burns, R.G., and Burns, V.M., 1977, Mineralogy, in Glasby, G.P., ed.,
Marine manganese deposits: Elsevier Oceanography Series, v. 15,
p. 185—248.

Calvert, SE, and Price, NR, 1970, Composition of manganese
nodules and manganese carbonates: Contributions to Mineralogy
and Petrology, v. 29, p. 215—233.

1977, Shallow water continental margin and lacustrine nodules:
Distribution and geochemistry, in Glasby, G.P., ed., Marine
manganese deposits: Elsevier Oceanography Series, v. 15, p.
45—86.

Campbell, A.C., Gieskes, J.M., Lupton, J.E., and Lonsdale, RE,
1988, Manganese geochemistry in the Guaymas Basin, Gulph of
California: Geochimica et Cosmochimica Acta, v. 52, p. 345‘357.

Chyi, M.S., Crerar, D.A., Carlson, R.W., and Stallard, R.F., 1984,
Hydrothermal Mn-deposits of the Franciscan assemblage, II.
Isotope and trace element geochemistry, and implications for
hydrothermal convection at spreading centers: Earth and Plane-
tary Science Letters, v. 71, p. 31—45.

Constantinou, G., and Govett, G.J.S., 1972, Genesis of sulphide depos-
its, ochre and umber of Cyprus: Transactions of the Institution of
Mining and Metallurgy, v. 81, p. B34—B46.

Corliss, J .B., Lyle, Mitchell, Dymond, Jack, and Crane, Kathy, 1978,
The chemistry of hydrothermal mounds near the Galapagos Rift:
Earth and Planetary Science Letters, v. 40, p. 12—24.

 

 

 

Cotkin, S.J., 1987, Conditions of metamorphism in an Early Paleozoic
blueschist, Schist of Skookum Gulch, northern California: Contri-
butions to Mineralogy and Petrology, v. 96, p. 192—200.

Courtois, Chantal, and Clauer, Norbert, 1980, Rare earth elements and
strontium isotopes of polymetallic nodules from southeastern
Paciﬁc Ocean: Sedimentology, v. 27, p. 687—695.

Cowan, D.S., 1985, Structural styles in Mesozoic and Cenozoic
melanges in the western Cordillera of North America: Geological
Society of America Bulletin, v. 96, p. 451—462.

Crerar, D.A., Namson, Jay, Chyi, M.S., Williams, Loretta, and
Feigenson, M.D., 1982, Manganiferous cherts of the Franciscan
Assemblage: I. General geology, ancient and modern analogues,
and implications for hydrothermal convection at oceanic spreading
centers: Economic Geology, v. 77, p. 519—540.

Cressman, ER, 1962, Nondetrital siliceous sediments: U.S. Geologi-
cal Survey Professional Paper 440—T, 23 p.

Cronan, D.S., 1977, Deep sea nodules: Distribution and geochemistry,
in Glasby, G.P., ed., Marine manganese deposits: Elsevier Ocean-
ography Series, v. 15, p. 11—44.

Dadey, K.A., and Leinen, Margaret, 1987, Anomalous porosity gradi-
ents in Mariana mounds sediments: Eos, v. 68, p. 1534.

DeLong, SE, and Fox, P.J., 1978, Geological consequences of ridge
subduction, in Talwani, M., and Pitman, W.C., III, eds., Island
arcs, deep sea trenches and back-arc basins: Ann Arbor, Mich.,
Lithocrafters, p. 199—210.

De Villiers, J .P.R., 1980, The crystal structure of braunite II and its
relation to bixbyite and braunite: American Mineralogist, v. 65,
p. 756—765.

De Villiers, F.R., 1970, The geology and mineralogy of the Kalahari
manganese-ﬁeld north of Sishen, Cape Province: Geological Sur-
vey of South Africa, Memoir 59, 65 p.

Dickinson, W.R., Beard, L.S., Brakenridge, G.R., Erjavec, J.L.,
Ferguson, R.G., and others, 1983, Provenance of North American
Phanerozoic sandstones in relation to tectonic setting: Geological
Society of America Bulletin, v. 94, p. 222—235.

Dickinson, W.R., Ingersoll, R.V., Cowan, D.S., Helmold, K.P., and
Suzek, C.A., 1982, Provenance of Franciscan graywackes in
coastal California: Geological Society of America Bulletin, v. 93,
p. 95—107.

Dunn, P.J., 1979, The chemical composition of gageite: an empirical
formula: American Mineralogist, v. 64, p. 1056—1058.

Dymond, Jack, Corliss, J .B., Cobler, Richard, Muratli, C.M., Chou,
Christin, and Conrad, Roberta, 1980, Composition and origin of
sediments recovered by deep drilling of sediment mounds, Galapa-
gos spreading center, in Powell, Robert, ed., Initial reports of the
deep sea drilling project, v. 54: Washington, D.G., U.S. Govern-
ment Printing Ofﬁce, p. 377—385.

Dymond, Jack, Lyle, Mitchell, Finney, Bruce, Piper, D.Z., Murphy,
Kim, Conrad, Roberta, and Pisias, Nicklas, 1984, Ferromanganese
nodules from MANOP Sites H, S, and R—Control of mineralogical
and chemical composition by multiple accretionary processes:
Geochimica et Cosmochimica Acta, v. 48, p. 931—949.

Edmund, J.M., Measures, 0., Mangum, B., Grant, B., Sclater, F.R.,
Collier, R., and Hudson, A., 1979, On the formation of metal-rich
deposits at ridge crests: Earth and Planetary Science Letters,
v. 46, p. 19—30.

Edmund, J.M., Von Damm, K.L., McDuff, R.E., and Measures, C.I.,
1982, Chemistry of hot springs on the East Paciﬁc Rise and their
efﬂuent dispersal: Nature, v. 297, p. 187—191.

Elderﬁeld, H., Hawkesworth, C.J., and Greaves, M.J., 1981, Rare
earth element geochemistry of oceanic ferromanganese nodules
and associated sediments: Geochimica et Cosmochimica Acta, v.
45, p. 513-528.

REFERENCES 69

Erd, R.C., and Ohashi, Y., 1984, Santaclaraite, a new calcium-
manganese silicate hydrate from California: American Mineralo—
gist, v. 69, p. 200—206.

Ernst, W.G., 1971, Petrologic reconnaissance of Franciscan metagray-
wackes from the Diablo Range, central California Coast Ranges:
Journal of Petrology, V. 12, p. 413—437.

Evans, H.T., Jr., and White, J.S., 1987, The colorful vanadium
minerals: a brief review and a new classiﬁcation: Mineralogical
Record, v. 18, p. 333—340.

Ferraris, Giovanni, Mellini, Marcello, and Merlino, Stefano, 1987,
Electron-diffraction and electron-microscopy study of balangeroite
and gageite: Crystal structures, polytypism, and ﬁber texture:
American Mineralogist, v. 72, p. 382—391.

Fleet, A.J., 1983, Hydrothermal and hydrogenous ferro-manganese
deposits: Do they form a continuum? The rare earth evidence, in
Rona, P.A., Bostrom, Kurt, Laubier, Lucien, and Smith, K.L.,
Jr., eds., Hydrothermal processes at seaﬂoor spreading centers:
New York, Plenum Press, p. 535—555.

Fleischer, Michael, 1981, The Ford-Fleischer ﬁle of mineralogical
references through 1977: US. Geological Survey Open—File
Report 81—1169, 221 sheets of microﬁche.

1987, Glossary of mineral species: Tucson, Arizona, Mineralog-
ical Record, 227 p.

Fleischer, Michael, and Richmond, W.E., 1943, The manganese oxide
minerals: a preliminary report: Economic Geology, v. 38,
p. 269—286.

Fleischer, Michael, and Schafer, CM, 1981, The Ford-Fleischer ﬁle of
mineralogical references, 1978—1980 inclusive: US. Geological
Survey Open-File Report 81—1174, 287 p.

1983, The Ford-Fleischer ﬁle of mineralogical references,
1981—1982 inclusive: US. Geological Survey Open-File Report
83—615, 417 p.

Flohr, M.J.K., and Huebner, J.S., 1988, Metamorphic manganese
mineral assemblages from Sierra Nevada, California ore deposits
and proposed origin [abs]: Geological Society of America
Abstracts with Programs, v. 20, p. A300.

Forsythe, R.D., Nelson, E.P., Carr, M.J., Kaeding, M.E., Herve, M.,
Mpodozis, C., Sofﬁa, J.M., and Harambour, S., 1986, Pliocene
near—trench magmatism in southern Chile: A possible manifesta—
tion of ridge collision: Geology, v. 14, p. 23—27.

Froelich, P.N., Klinkhammer, G.P., Bender, M.L., Luedtke, N.A.,
Heath, G.R., Cullen, Doug, and Dauphin, Paul, 1979, Early
oxidation of organic matter in pelagic sediments of the eastern
equatorial Atlantic: suboxic diagenesis: Geochimica et Cosmochim-
ica Acta, V. 43, p. 1075—1090.

Fyfe, W.S., 1974, Heats of chemical reactions and submarine heat
production: Geophysical Journal of the Royal Astronomical Soci-
ety, v. 37, p. 213—215.

Gardner, W.D., Sullivan, L.G., and Thorndike, E.M., 1984, Long-term
photographic, current, and nephelometer observations of manga-
nese nodule environments in the Paciﬁc: Earth and Planetary
Science Letters, v. 70, p. 95—109.

Ghiorse, W.C., and Hirsch, P., 1979, An ultrastructural study of iron
and manganese deposition associated with extracellular polymers
of pedomicrobium-like budding bacteria: Archives of Microbiology,
v. 123, p. 213—226.

Glasby, GR, 1976, Surface densities of manganese nodules in the
southern sector of the South Paciﬁc: New Zealand Journal of
Geology and Geophysics, v. 19, p. 771—790.

Goddard, E.N., Trask, P.D., De Ford, R.K., Rove, O.N., Singewald,
J.T., Jr., and Overbeck, RM, 1970, Rock color chart: Boulder,
0010., Geological Society of America.

 

 

 

Gromet, P.L., Dymek, R.F., Haskin, L.A., and Korotev, R.L., 1984,
The “North American Shale Composite”: Its compilation, major
and trace element characteristics: Geochimica et Cosmochimica
Acta, v. 48, p. 2469—2482.

Guggenheim, Stephen, Bailey, S.W., Eggleton, R.A., and Wilkes,
Peter, 1982, Structural aspects of greenalite and related minerals:
Canadian Mineralogist, v. 20, p. 1—18.

Guggenheim, Stephen, and Eggleton, R.A., 1987, Modulated 2:1 layer
silicates: Review, systematics, and predictions: American Miner-
alogist, v. 72, p. 724—738.

Haggerty, J .A., 1987, Cold-water, deep-sea chimneys from the Mari-
ana forearc serpentinite seamounts: Eos, v. 68, p. 1534.

Halbach, P., 1986, Processes controlling the heavy metal distribution in
Paciﬁc ferromanganese nodules and crusts: Geologische Rund-
schau, v. 75, p. 235—247.

Halbach, P., Scherhag, C., Hebisch, U., and Marchig, V., 1981,
Geochemical and mineralogical control of different genetic types of
deep-sea nodules from the Paciﬁc Ocean: Mineralium Deposita, v.
16, p. 59—84.

Hathaway, J.C., and Degens, ET, 1969, Methane-derived marine
carbonates of Pleistocene age: Science, v. 165, p. 690—692.

Hein, J .R., and Karl, SM, 1983, Comparisons between open—ocean
and continental margin chert sequences, in Iijima, A., Hein, J .R.,
and Siever, Raymond, eds., Siliceous deposits in the Paciﬁc region:
Amsterdam, Elsevier, p. 25—43.

Hein, J .R., and Koski, R.A., 1987, Bacterially mediated diagenetic
origin for chert-hosted manganese deposits in the Franciscan
Complex, California Coast Ranges: Geology, v. 15, p. 722—726.

1988, Reply to Comment on “Bacterially mediated diagenetic
origin for chert-hosted manganese deposits in the Franciscan
Complex, California Coast Ranges”: Geology, V. 16, p. 470.

Hein, J.R., Koski, R.A., and Yeh, Hsueh-Wen, 1987, Chert-hosted
manganese deposits in sedimentary sequences of the Franciscan
Complex, Diablo Range, California, in Hein, J .R., ed., Siliceous
sedimentary rock-hosted ores and petroleum: New York, Van
Nostrand Reinhold, p. 206—230.

Hein, J .R., and Parrish, J .T., 1987 , Distribution of siliceous deposits in
space and time, in Hein, J .R., ed., Siliceous sedimentary rock-
hosted ores and petroleum: New York, Van Nostrand Reinhold,
p. 10—57.

Hekinian, R., Rosendahl, B.R., Cronan, D.S., Dmitriev, Y., Fodor,
R.V., and others, 1978, Hydrothermal deposits and associated
basement rocks from the Galapagos spreading center: Oceanolog-
ica Acta, v. 1, p. 473—482.

Hem, J .D., 1972, Chemical factors that influence the availability of iron
and manganese in aqueous systems: Geological Society of America
Bulletin, v. 83, p. 443—450.

Herzig, P.M., Becker, K.P., Stoffers, Peter, Backer, Harald, and
Blum, Norbert, 1988, Hydrothermal silica chimney ﬁelds in the
Galapagos Spreading Center at 86°W: Earth and Planetary Science
Letters, v. 89, p. 261—272.

Hewett, D.F., 1972, Manganite, hausmannite, braunite: Features,
modes of origin: Economic Geology, V. 67, p. 83-102.

Hewett, D.F., Chesterman, C.W., and Troxel, B.W., 1961, Tephroite
in California manganese deposits: Economic Geology, v. 56,
p. 39—58.

Heye, D., Marchig, V., and Meyer, H., 1979, The growth of buried
manganese nodules: Deep-Sea Research, v. 26A, p. 789—798.
Hodgson, W.A., 1966, Carbon and oxygen isotope ratios in diagenetic
carbonates from marine sediments: Geochimica et Cosmochimica

Acta, v. 30, p. 1223—1233.

Hoffert, M., Perseil, A., Hekinian, R., Choukroune, P., Needham,

H.D., Francheteau, J., and Le Pichon, X., 1978, Hydrothermal

 

70 MICROBANDED MANGANESE FORMATIONS: PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

deposits sampled by diving saucer in transform fault “A” near 37°N
on the Mid-Atlantic Ridge, Famous Area: Oceanologica Acta, v. 1,
p. 73-86.

Hoffert, M., Person, A., Courtois, C., Karpoff, A.M., and Trauth, D.,
1980, Sedimentology, mineralogy, and geochemistry of hydrother-
mal deposits from holes 424, 424A, 424B, and 424C (Galapagos
spreading center), in Powell, Robert, ed., Initial reports of the
deep sea drilling project, v. 54: Washington, D.C., U.S. Govern-
ment Printing Ofﬁce, p. 339—376.

Hogdahl, O.T., Melsom, Sigurd, and Bowen, V.T., 1968, Neutron
activation analysis of lanthanide elements in sea water, in Gould,
R.F., ed., Trace inorganics in water: Washington, D.C., American
Chemical Society, Advances in Chemistry Series 73, p. 308—325.

Honjo, Sosumu, 1982, Seasonality and interaction of biogenic and
lithogenic particulate ﬂux at the Panama Basin: Science, v. 218,
p. 883—884.

Honnorez, Jose, von Herzen, R.P., Barrett, T.J., Becker, Keir,
Bender, M.L., and others, 1981, Hydrothermal mounds and young
ocean crust of the Galapagos: Preliminary deep sea drilling results,
Leg 702 Geological Society of America Bulletin, v. 92, p. 457—472.

Huebner, J .S., 1967, Stability relations of minerals in the system
Mn—Si-C-O: Baltimore, Md., The Johns Hopkins University, Ph.D.
thesis, 279 p.

1969, Stability relations of rhodochrosite in the system

manganese-carbon-oxygen: American Mineralogist, V. 54, p.

457—481.

1976, The manganese oxides—a bibliographic commentary, in
Reviews in Mineralogy, v. 3, Oxide Minerals: Mineralogical Soci-
ety of America, Washington D.C., p. SH—l — SH—17.

Huebner, J.S., Flohr, M.J.K., and Jones, B.F., 1986a, Origin of Mn
protore, Franciscan Complex, CA [abs.]: Geological Society of
America Abstracts with Programs, v. 18, p. 642.

Huebner, J.S., Flohr, M.J.K., and Matzko, J.J., 1986b, Origin and
metamorphism of manganese sediments in melange of the Fran-
ciscan Complex, California [abs], in International Mineralogical
Association Abstracts with Program: Washington, D.C., Mineral-
ogical Society of America, p. 131.

Huebner, J .S., and Woodruff, M.E., 1985, Chemical compositions and
critical evaluation of microprobe standards available in the Reston
microprobe facility: US. Geological Survey Open-File Report
85—718, 246 p.

Huey, A.S., 1948, Geology of the Tesla quadrangle, California: Cali-
fornia Division of Mines Bulletin 140, 75 p.

Ingersoll, R.V., 1982, Initiation and evolution of the Great Valley
forearc basin of northern and central California, U.S.A., in
Leggett, J.K., ed., Trench-forearc geology: Sedimentation and
tectonics on modern and ancient active plate margins: Oxford,
Blackwell Scientiﬁc Publications, p. 459—467.

Jenkyns, H.C., 1977, Fossil nodules, in Glasby, G.P., ed., Marine
manganese deposits: Elsevier Oceanography Series, v. 15,
p. 87—108.

Johannes, W., and Puhan, D., 1971, The calcite-aragonite transition,
reinvestigated: Contributions to Mineralogy and Petrology, v. 31,
p. 28—38.

Johnson, A.M., and Ellen, SD, 1974, Theory of concentric, kink, and
sinusoidal folding and of monoclinal ﬂexuring of compressible,
elastic multilayers: Tectonophysics, V. 21, p. 301—339.

Jonasson, I.R., and Walker, D.A., 1987, Micro-organisms and their
debris as substrates for base metal sulﬁde nucleation and accumu—
lation in some mid-ocean ridge deposits [abs]: Eos, V. 68, p. 1546.

Jones, BF, and Weir, A.H., 1983, Clay minerals of Lake Aber't, an
alkaline, saline lake: Clays and Clay Minerals, v. 31, p. 161—172.

Kalhorn, Susan and Emerson, Steven, 1984, The oxidation state of
manganese in surface sediments of the deep sea: Geochimica et
Cosmochimica Acta, V. 48, p. 897—902.

 

 

 

Kleyenstﬁber, A.S.E., 1984, The mineralogy of the manganese-bearing
Hotazel Formation, of the Proterozoic Transvaal Sequence in
Griqualand West, South Africa: Transactions of the Geological
Society of South Africa, v. 87, p. 257—272.

Klinkhammer, G., Elderﬁeld, H., Greaves, M., Rona, P., and Nelsen,
T., 1986, Manganese geochemistry near high-temperature vents in
the Mid-Atlantic Ridge rift valley: Earth and Planetary Science
Letters, V. 80, p. 230—240.

Klinkhammer, G., and Hudson, A., 1986, Dispersal patterns for
hydrothermal plumes in the South Paciﬁc using manganese as a
tracer: Earth and Planetary Science Letters, V. 79, p. 241—249.

Krauskopf, KB, 1957, Separation of manganese from iron in sedimen-
tary processes: Geochimica et Cosmochimica Acta, v. 12, p. 61—84.

Lalou, Claude, 1983, Genesis of ferromanganese deposits: hydrother-
mal origin, in Rona, P.A., Bostrom, Kurt, Laubier, Lucien, and
Smith, K.L., J r., eds., Hydrothermal processes at seaﬂoor spread-
ing centers: New York, Plenum Press, p. 503—534.

Lattard, Dominique, and Schreyer, Werner, 1981, Experimental
results bearing on the stability of the blueschist-facies minerals
deerite, howieite, and zussmanite, and their petrological signiﬁ-
cance: Bulletin de Minéralogie, V. 104, p. 431—440.

Leinen, Margaret, McDuff, R., Delaney, J., Becker, K., and Schul-
theiss, P., 1987, Off-axis hydrothermal activity in the Mariana
mounds ﬁeld: Eos, V. 68, p. 1531.

Lonsdale, Peter, 1977, Deep-tow observations at the Mounds Abyssal
Hydrothermal Field, Galapagos Rift: Earth and Planetary Science
Letters, V. 36, p. 92—110.

Lonsdale, P.F., Bischoff, J.L., Burns, V.M., Kastner, M., and
Sweeney, RE, 1980, A high-temperature hydrothermal deposit
on the seabed at a Gulf of California spreading center: Earth and
Planetary Science Letters, V. 49, p. 8—20.

Lucchetti, G., Cortesogno, L., and Palenzona, A., 1988, Low-
temperature metamorphic mineral assemblages in Mn—Fe ores
from Cerchiara mine (northern Apennine, Italy): Neues J ahrbuch
fur Mineralogie Monatschefte, no. 8, p. 367—383.

Lyle, Mitchell, Heath, GR, and Robbins, J .M., 1984, Transport and
release of transition elements during early diagenesis: Sequential
leaching of sediments from MANOP Sites M and H. Part I. pH 5
acetic acid leach: Geochimica et Cosmochimica Acta, v. 48,
p. 1705—1715.

Lynn, D.C., and Bonatti, E., 1965, Mobility of manganese in diagenesis
of deep—sea sediments: Marine Geology, v. 8, p. 457-474.

Maddock, M.E., 1955, Geology of the Mount Boardman quadrangle,
California: Berkeley, Calif, University of California, Ph.D. thesis,
167 p.

Malahoff, A., McMurtry, G.M., Smith, J.R., and Hunneke, W.F.,
1987, Neovolcanism, neotectonics and hydrothermal activity on
Loihi submarine volcano, Hawaii [abs]: Eos, V. 68, p. 1553.

Marchig, Vesna, and Halbach, P., 1982, Internal structures of manga-
nese nodules related to conditions of sedimentation: Tschermaks
Mineralogische und Petrographische Mitteilungen, V. 30,
p. 81—110.

Margolis, S.V., and Burns, R.G., 1976, Paciﬁc deep-sea manganese
nodules: their distribution, composition, and origin: Annual
Review of Earth and Planetary Sciences, v. 4, p. 229—263.

Margolis, S.V., Ku, T.L., Glasby, G.P., Fein, C.D., and Audley-
Charles, M.G., 1978, Fossil manganese nodules from Timor: Geo-
chemical and radiochemical evidence for deep-sea origin: Chemical
Geology, V. 21, p. 185—198.

Matsubara, Satoshi, 1981, Taneyamalite, (Na,Ca)(Mn2*,Mg,Fe3+,Al)12
Si12(O,OH)44, a new mineral from the Iwaizawa mine, Saitama
Prefecture, Japan: Mineralogical Magazine, v. 44, p. 51-53.

Matsumoto, Ryo, 1987, Origin of manganese nodules in the Jurassic
siliceous rocks of the Inuyama District, central Japan, in Hein,

REFERENCES 71

J .R., ed., Siliceous sedimentary rock-hosted ores and petroleum:
New York, Van Nostrand Reinhold, p. 181—205.

Maynard, J .B., Valloni, Renzo, and Yu, Ho-Shing, 1982, Composition
of modern deep-sea sands from arc-related basins, in Leggett,
J .K., ed., Trench-forearc geology: Sedimentation and tectonics on
modern and ancient active plate margins: Oxford, Blackwell Sci-
entiﬁc Publications, p. 551—561.

McGee, J .J ., 1985, ARL—SEMQ electron microprobe operating manual:
US. Geological Survey Open-File Report 85—387, 62 p.

Mohr, P.A., 1956, Trace-element distribution in garnet-chlorite rock
from Foel Ddu, near Harlech, Merionethshire: Mineralogical Mag-
azine, v. 31, p. 319—327.

Moore, P.B., 1969, A novel octahedral framework structure: Gageite:
American Mineralogist, v. 54, p. 1005—1017.

Mosser, C., 1979, Elements traces dans quelques argiles des altér—
ations et des sediments, in Ahrens, L.H., ed., Origin and distri-
bution of the elements: New York, Pergamon Press, p. 315—329.

Mosser, C., Leprun, J .C., and Blot, A., 1985, Les elements traces des
fractions <2 mm a kaolinite et smectite formées par alteration des
roches silicatées acides Afrique de L’Ouest (Sénégal et Haute-
Volta): Chemical Geology, v. 48, p. 165—181.

Muir Wood, Robert, 1980, The iron-rich blueschist-facies minerals: 3.
Zussmanite and related minerals: Mineralogical Magazine, v. 43,
p. 605—614.

Murata, K.J., and Norman, M.B., II, 1976, An index of crystallinity for
quartz: American Journal of Science, v. 276, p. 1120—1130.

Nahon, D., Colin, F., and Tardy, Y., 1982, Formation and distribution
of Mg,Fe,Mn-smectites in the ﬁrst stages of the lateritic weather-
ing of forsterite and tephroite: Clay Minerals, v. 17, p. 339—348.

Nealson, K.H., 1983, The microbial manganese cycle, in Krumbein,
W.E., ed., Microbial geochemistry: Oxford, Blackwell Scientiﬁc
Publications, p. 191—221.

Nisbet, E.G., and Price, Ilfryn, 1974, Siliceous turbidites: bedded
cherts as redeposited ocean ridge-derived sediments: International
Association of Sedimentology, Special Publication 1, p. 351—366.

Ohashi, Yoshikazu, and Finger, L.W., 1981, The crystal structure of
santaclaraite, CaMn4[Si5014(OH)](OH)-H20: the role of hydrogen
atoms in the pyroxenoid structure: American Mineralogist, V. 66,
p. 154—168.

Okita, P.M., Maynard, J.B., Spiker, E.G., and Force, ER, 1988,
Isotopic evidence for organic matter oxidation by manganese
reduction in the formation of stratiform manganese carbonate ore:
Geochimica et Cosmochimica Acta, v. 52, p. 2679—2685.

Ostwald, J ., 1981, Evidence for a biogeochemical origin of the Groote
Eylandt manganese ores: Economic Geology, v. 76, p. 556-567.

Pedersen, T.F., and Price, N.B., 1982, The geochemistry of manga-
nese carbonate in Panama Basin sediments: Geochimica et Cosmo-
chimica Acta, v. 46, p. 59—68.

Pettijohn, F.J., 1957, Sedimentary rocks: New York, Harper and
Brothers, 718 p.

1963, Chemical composition of sandstones—excluding carbon-
ate and volcanic sands: US. Geological Survey Professional Paper
440—S, p. S1—S21.

Phillips, HA, 1911, Notes on a recent ﬁnd of zincite crystals:
American Journal of Science, 4th ser., v. 31, p. 464—465.

Raab, W.J. and Meylan, M.A., 1977, Morphology, in Glasby, G.P., ed.,
Marine manganese deposits: Elsevier Oceanography Series, v. 15,
p. 109—146.

Raymond, L.A., 1973a, Franciscan geology of the Mt. Oso area,
California: Davis, Calif., University of California at Davis, Ph.D.
thesis, 223 p.

1973b, Tesla-Ortigalita fault, Coast Range thrust fault, and

Franciscan metamorphism, northeastern Diablo Range, Califor-

nia: Geological Society of America Bulletin, v. 84, p. 3547—3562.

 

 

 

1974, Possible modern analogues for rocks of the Franciscan
Complex, Mount Oso area, California: Geology, v. 2, p. 143—146.
1977, Structural control of manganese deposits in subduction
complexes: California Coast Range and Southern Appalachian
examples [abs]: Geological Society of America Abstracts with
Programs, v. 9, p. 486—487.

1984, Classiﬁcation of melanges: Geological Society of America

Special Paper 198, p. 7—20.

1988, Comment on “Bacterially mediated diagenetic origin for
chert-hosted manganese deposits in the Franciscan Complex,
California Coast Ranges”: Geology, v. 16, p. 469—470.

Rona, P.A., 1978, Criteria for recognition of hydrothermal mineral
deposits in oceanic crust: Economic Geology, v. 73, p. 135—160.

Roy, Supriya, Dasgupta, Somnath, Majumdar, N., Banerjee, H.,
Bhattacharya, PK, and Fukuoka, M., 1986, Petrology of manga—
nese silicate-carbonate-oxide rock of Sausar Group, India: Neues
Jahrbuch fur Mineralogie Monatschefte, no. 12, p. 561—568.

Schrader, E.L., Rosendahl, B.R., Furbish, W.J., and Mattey, DR,
1980, Mineralogy and geochemistry of hydrothermal and pelagic
sediments from the mounds hydrothermal ﬁeld, Galapagos spread-
ing center, DSDP Leg 54: Journal of Sedimentary Petrology, v. 50,
p. 917—928.

Schweikert, R.A., and Cowan, D.S., 1975, Early Mesozoic tectonic
evolution of the western Sierra Nevada, California: Geological
Society of America Bulletin, v. 86, p. 1329—1336.

Scott, M.R., Scott, R.B., Rona, P.A., Butler, L.W., and Nalwalk,
A.J., 1974, Rapidly accumulating manganese deposit from the
median valley of the Mid—Atlantic Ridge: Geophysical Research
Letters, v. 1, p. 355-358.

Seiders, V.M., and Blome, GD, 1988, Implications of upper Mesozoic
conglomerate for suspect terrane in western California and adja-
cent areas: Geological Society of America Bulletin, v. 100, p.
374—391.

Shannon, RD, 1976, Revised effective ionic radii and systematic
studies of interatomic distances in halides and chalcogenides: Acta
Crystallographica, v. A32, p. 751—767.

Snyder, W.S., 1978, Manganese deposited by submarine hot springs in
chert—greenstone complexes, western United States: Geology,
v. 6, p. 741—744.

Stanaway, K.J., Kobe, H.W., and Sekula, J ., 1978, Manganese depos-
its and associated rocks of Northland and Auckland, New Zealand:
New Zealand Journal of Geology and Geophysics, v. 21, p. 21—32.

Stoffers, P., Glasby, GR, and Frenzel, G., 1984, Comparison of the
characteristics of manganese micronodules from the equatorial and
southwest Paciﬁc: Tschermaks Mineralogische und Petrographi-
sche Mitteilungen, V. 33, p. 1—23.

Suess, Erwin, 1979, Mineral phases formed in anoxic sediments by
microbial decomposition of organic matter: Geochimica et Cosmo-
chimica Acta, v. 43, p. 339—352.

Sundby, Bjorn, Anderson, L.G., Hall, P.O.J., Iverfeldt, Ake, Rutgers
van der Loeff, M.M., and Westerlund, S.F.G., 1986, The effect of
oxygen on release and uptake of cobalt, manganese, iron and
phosphate at the sediment-water interface: Geochimica et Cosmo-
chimica Acta, v. 50, p. 1281—1288.

Takahashi, Kozo, 1986, Seasonal ﬂuxes of pelagic diatoms in the
subarctic Paciﬁc, 1982—1983: Deep Sea Research, v. 33, p.
1225—1251.

Taliaferro, N.L., and Hudson, F.S., 1943, Genesis of the manganese
deposits of the Coast Ranges of California: California Division of
Mines Bulletin 125, p. 217—275.

Taylor, S.R., and McLennan, SM, 1985, The continental crust: Its
composition and evolution: Oxford, Blackwell Scientiﬁc Publica-
tions, 312 p.

Thornton, E.G., and Seyfried, W.E., Jr., 1985, Sediment—sea-water
interaction at 200 and 300°C, 500 bars pressure: The role of

 

 

 

 

72 MICROBANDED MANGANESE FORMATIONS; PROTOLITHS IN THE FRANCISCAN COMPLEX, CALIFORNIA

sediment composition in diagenesis and low-grade metamorphism
of marine clay: Geological Society of America Bulletin, v. 96,
p. 1287—1295.

Toth, J .R., 1980, Deposition of submarine crusts rich in manganese and
iron: Geological Society of America Bulletin, v. 91, p. 44—54.
Trask, P.D., 1950, Geologic description of the manganese deposits of

California: California Division of Mines Bulletin 152, 378 p.
Trask, P.D., Wilson, LR, and Simons, RS, 1943, Manganese deposits
of California—a summary report (including tabulated data on
manganese properties of California): California Division of Mines
Bulletin 125, p. 51—215.
Treng‘rove, R.R., 1960, Reconnaissance of California manganese
deposits: US. Bureau of Mines Report of Investigations 5579, 46

p.

Volin, M.E., and Matson, E.J., 1949, Investigation of the Buckeye
manganese deposits, Stanislaus County, Calif: US. Bureau of
Mines Report of Investigations 4560, [16] p.

Vonderhaar, D.L., McMurtry, G.M., Fryer, P., and Malahoff, A.,
1987, Geochemistry of hydrothermal deposits from active subma-
rine volcanoes in the Mariana arc: Eos, v. 68, p. 1533.

Von Damm, K.L., Edmond, J.M., Measures, C.I., and Grant, B., 1985,

 

Chemistry of submarine hydrothermal solutions at Guaymas
Basin, Gulf of California: Geochimica et Cosmochimica Acta, v. 49,
p. 2221—2237.

Watanabe, Takeo, Yui, Shunzo, and Kato, Akira, 1970, Bedded man-
ganese deposits in Japan, a review, in Tatsumi, T., ed., Volcanism
and ore genesis: Tokyo, University of Tokyo Press, p. 119—142.

Westerlund, S.F.G., Anderson, L.G., Hall, P.O.J., Iverfeldt, Ake,
Rutgers van der Loeff, M.M., and Sundby, Bjorn, 1986, Benthic
ﬂuxes of cadmium, copper, nickel, zinc and lead in the coastal
environment: Geochimica et Cosmochimica Acta, v. 50,
p. 1289—1296.

Wheat, C.G., and McDuff, RE, 1987, Advection of pore waters in the
Marianas Mounds hydrothermal region as determined from nutri-
ent proﬁles [abs]: Eos, v. 68, p. 1531.

Yamamoto, Koshi, 1987, Geochemical characteristics and depositional
environments of cherts and associated rocks in the Franciscan and
Shimanto terranes: Sedimentary Geology, v. 52, p. 65-108.

Yeh, H—W., Hein, J .R., and Koski, R.A., 1985, Stable isotope study of
volcanogenic- and sedimentary—manganese deposits: U.S. Geolog—
ical Survey Open-File Report 85—662, 20 p.

 

 

COVER.—Mississippian formations in Burnside quadrangle, Pulaski County, Kentucky; Monteagle Limestone, Hartselle
Formation, Bangor Limestone. and Paragon (formerly Pennington) Formation.

“LUUMENTS DEPARTMENT
JAN 14 1991

  

1 a LIBRARY~l
g. NWERSITYWQLQ ﬂown

Mississippian Rocks in Kentucky

By EDWARD G. SABLE ana’ GARLAND R. DEVER, JR.

 

U.S. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1503

Work done in cooperation with the Kentucky
Geological Survey

Strata representing Kinderhookian, Osagean, Meramecian,
ana’ Chesterian Series and equivalents reﬂect largely
marine deposition in shallow cratonic basins and on

shelves and platforms

 

 

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON: 1990

DEC 1 71990

DEPARTMENT OF THE INTERIOR

MANUEL LUJAN, JR., Secretary

U.S. GEOLOGICAL SURVEY

Dallas L. Peck, Director

 

Library of Congress Cataloging-in-Publication Data

Sable, Edward G. (Edward, George), 1924—
Mississippian rocks in Kentucky / by Edward G. Sable and Garland R. Dever, Jr.
p. cm. — (U.S. Geological Survey professional paper ; 1503)

“Work done in cooperation with the Kentucky Geological Survey.” “Strata representing Kinderhookian, Osagean,
Meramecian, and Chesterian series and equivalents reﬂect largely marine deposition in shallow cratom'c basins and on
shelves and platforms.”

Includes bibliographical references.

Supt. of Docs. no.: I 19.16: 1503

1. Geology, Stratigraphic—Mississippian. 2. Geology—Kentucky.

I. Dever, Garland R. II. Kentucky Geological Survey. III. Title. IV. Series.
QE672.822 1990
551.7'51'09769—dc20 89—600367

CIP

 

For sale by the Books and Open-File Reports Section, U.S. Geological Survey,
Federal Center, Box 25425, Denver, CO 80225

Any use of trade, product, or ﬁrm names in this publication is for descriptive purposes only
and does not imply endorsement by the U.S. Government.

CONTENTS

 

Page Page
Abstract ........................................... 1 Mississippian rocks—Continued
Introduction ....................................... 1 Rocks of mostly Osagean age—Continued
History of investigations ......................... 2 Age ---------------------------------------- 53
Scope and methods .............................. 2 Paleotectonic implications ..................... 53
Acknowledgments ............................... 3 Osagean-Meramecian series boundary ............... 54
Geologic framework .............................. 3 Rocks of late Osagean and Meramecian age ......... 55
Physiography, outcrop, and scenic features .......... 20 aﬁﬁbﬁiegggmm """""""""""" g:
D”Eifiiiﬁ‘séifiéfi’xgi‘ffid' ensue; : : :1: 1:: : :3 Seleee Lieeeeeee end seem eee Weeeew Feeeeeem
Bedford Shale and Berea Sandstone ................ 31 “m :"2 """""""""""""""""""" 56
Devonian-Mississippian systemic boundary 32 St' Louis Limestone """""""""""""" 58
_ , , , """"""" Ste. Genevieve Limestone ..................... 61
MISSISSIPPI” rocks ' ’ ‘ " """"""""""""""" 33 Stratigraphic relationships ..................... 66
R001“ of @ﬂeﬁhmklan and early Osagean age (Early St. Louis and older unit relationships ........ 66
MISSISSIPPI”) """""""""""""""" 33 St. Louis-Ste. Genevieve relationships ....... 67
Sunbury Shale ---------------------------- 33 Sources of sediments and depositional environments 69
Maury Formation equivalent ................... 33 Paleotectonic implications ..................... 71
Rockford Limestone .......................... 35 Meramecian-Chesterian series boundary ............. 71
Kinderhookian-Osagean unit relationships ........ 35 Rocks of Chesterian age .......................... 72
Paleotectonic implications ..................... 36 Rocks west 0f the Cincinnati arch .............. 73
Rocks of mostly Osagean age ..................... 36 Carbonate-dominated units ................. 73
Borden Formation ............................ 36 Terrigenous elastic units ................... 74
Farmers Member ......................... 33 Rocks east of the Cincinnati arch ............... 77
Nancy and New Providence Members ''''''' 39 Lithologic trends and interbasin correlations ..... 82
Cowbell, Halls Gap, and Holtsclaw Members . . 40 Sources of sediments and depositional environments 85
Nada and Wildie Members ................. 40 Paleogeography and paleotectonic implications ' ‘ ' 87
Re fr M be 42 Mississippian-Pennsylvaman-younger rock relationships . . A 87
n O em 1' ........................... . . .
Summary of thicknesses and lithologic trends ........... 92
Muldraugh Member ....................... 45 C li d 't' 103
yc c eposi ion ....................................
Floyds Knob _Bed """"""""""""" 45 Paleontology ....................................... 104
Gramger Formation .......................... 48 Macrofossils __________________________________ 104
Fort Payne Formation ------------------------ 48 Microfossils ..................................... 107
Strata overlying Borden and Fort Payne Formations 51 Conodont identifications ....................... 107
Depositional history of the Borden and Fort Payne Flora .......................................... 107
Formations ................................ 51 References cited .................................... 113
ILLUSTRATIONS
Page
PLATE 1. Map of Kentucky showing 71/2-minute quadrangles ................................................. In pocket
FIGURE 1. Map showing generalized structural elements of Kentucky and adjoining areas ........................... 4
2. Map showing distribution of Mississippian, pre-Mississippian, and post-Mississippian rocks exposed in Kentucky 5
3. Map of Kentucky showing 71/2-minute quadrangle boundaries and geographic divisions used in this report . . . . 5
4. Index map of Kentucky showing quadrangles in which Mississippian rocks are exposed .................... 6
5. Map of Kentucky showing control point localities of Mississippian surface and subsurface information other than
that from geologic quadrangle maps ............................................................ 7
6. Pre-Pennsylvanian—post-Ordovician Megagroups and Sequences in the Eastern Interior basin and adjoining areas 20
7. Correlation chart showing stratigraphic nomenclature of the Mississippian System in Kentucky and Illinois . 21
8. Correlation chart showing stratigraphic nomenclature of the Chesterian Series‘and equivalent strata in Kentucky 22
9. Cross section showing Mississippian rock units and relationships in Kentucky ............................ 23
10. Physiographic diagram of Kentucky ................................................................ 24
11. Map of Kentucky showing locations of selected exposures of Mississippian rocks, highways. and other features 25

III

IV

FIGURE

12.

13.

14.

15.
16.

17.

18.

19.
20.

21.
22—27.

28.
29.
30.
31.
32.
33.
34.
35.

36-40.

41.
42.
43.
44.
45.
46.

47—49.

50.
51-64.

CONTENTS

Map showing thickness of Upper Devonian rocks and their relationship to major structural elements in the east-
central United States ........................................................................
Diagram showing generalized relationships between Berea Sandstone and ‘Bedford Shale in Vanceburg quadrangle,
northeastern Kentucky ........................................................................
Photograph of Sunbury Shale overlying the Bedford Shale and underlying the Farmers Member of the Borden
Formation ..................................................................................
Photograph of Maury Formation equivalent .........................................................
Map showing distribution, thickness, and sedimentary transport directions of selected sandstone units in the Borden
and Fort Payne Formations in Kentucky ........................................................
Restored cross section of the Borden Formation delta “front” in Howardstown quadrangle, north-central Kentucky,
showing relationships of Borden units ...........................................................
Generalized stratigraphic diagram of lower part of the Borden Formation and underlying units in northeastern
Kentucky ...................................................................................
Photograph of Borden Formation; contact of Cowbell Member and underlying poorly resistant Nancy Member
Sections showing relationships of siltstone units in the Borden Formation in the area between Brodhead and Bighill,
east-central Kentucky .........................................................................
Diagrams showing relationships of the clastic units of the Borden Formation, central and south-central Kentucky
Photographs showing:
22. Nancy Member and overlying Muldraugh Member of the Borden Formation ........................
23. Cowbell Member of Borden Formation (siltstone and shale) ......................................
24. Halls Gap and Muldraugh Members of Borden Formation .......................................
25. Nada Member of Borden Formation and overlying Slade Formation consisting of Renfro Member and limestone
and dolomite mostly of the St. Louis and Holly Fork Members ...............................
26. Dolosiltite and calcarenite 1n Muldraugh Member of the Borden Formation ........................
27. Upper part of Muldraugh Member of Borden Formation showing westward- -thickening calcarenite lenses
Generalized diagram showing relationships of Salem- Warsaw, Salem, Harrodsburg, Borden, and Fort Payne Forma-
tions in three Kentucky quadrangles ............................................................
Photograph of Fort Payne Formation in roadcut along launching ramp road, Lake Cumberland State Park
Photograph of Salem and Warsaw Formations undivided and lower part of St Louis Limestone .............
Map showing approximate distribution of Science Hill Sandstone and Garrett Mill Sandstone Members of Warsaw
Formation and J abez and Kniﬂey Sandstone Members of Fort Payne Formation 1n south-central Kentucky
Photograph of cherty dolomite in upper part of St Louis Limestone ....................................
Photograph of dolomite 1n middle part of St Louis Limestone ..........................................
Generalized diagram showing relationships between Ste Genevieve and St. Louis Limestones in part of western
Kentucky ...................................................................................
Stratigraphic section and description showing lithologies of St. Louis and Ste. Genevieve Limestones and Lost River
Chert Bed, west-central Kentucky ..............................................................
Photographs showing:
36. Lost River Chert Bed ......................................................................
37. Ste. Genevieve and St. Louis Limestones ......................................................
38. Upper part of Mooretown Formation, Beaver Bend Limestone, and shale and sandstone of the Sample Sand-
stone .................................................................................
39. Big Clifty Sandstone Member of the Golconda Formation .......................................
40. Hardinsburg Sandstone conformably overlying Haney Limestone Member of the Golconda Formation . .
Generalized map and cross section showing distribution and relationships of Mooretown Formation-Bethe] Sand-
stone channel fill in central to western Kentucky ..................................................
Diagram showing relationships of Glen Dean Limestone. Tar Springs Formation, and lower part of Buffalo Wallow
Formation, west-central Kentucky ..............................................................
Photograph of roadcut mostly in Slade (Newman) Formation, Rockcastle County ..........................
Photograph of section at Strunk Crushed Stone Company quarry. Tateville, Pulaski County ................
Diagram showing relationships of Girkin Formation with coeval units of Chesterian age. west-central Kentucky
Map showing distribution of sandstone and shale in Big Clifty Sandstone Member of Golconda Formation in west-
central, central, and western Kentucky, and Hartselle Formation in south-central Kentucky .............
Photographs showing:
47. Sandstone of Caseyville Formation with basal thin coal bed overlying limestone equivalent of Kinkaid Limestone
48. Caseyville Formation disconformably overlying Buffalo Wallow Formation .........................
49. Sandstone of Caseyville Formation overlying Vienna Limestone ..................................
Generalized cross sections showing Pennsylvanian and and Mississippian unit relationships in northeastern Kentucky
Maps showing:
51. Total thickness of Mississippian rocks in Kentucky .............................................
52. Cumulative thickness of Bedford Shale, Berea Sandstone, and Sunbury Shale in eastern Kentucky, and thickness
of approximate age-equivalent strata in other parts of Kentucky ..............................
53. Thickness of terrigenous elastic units of the Borden Formation, New Providence and Grainger Formations,
and basal shale beds in the Fort Payne Formation ..........................................

Page

31

32

34
35

37

38

39
39

40
41

42
43
44

45
46
47

49
50
56
58
59
60
61
63

64
65

74
75
76
77
78
79
80
83
85
88
89
90
92
93
93

94

FIGURE

TABLE

54.
55.
56.
57.
58.
59.
60.
61.
62.
63.

64.

CONTENTS

Thickness of Muldraugh Member of Borden Formation and of Fort Payne Formation exclusive of basal shale
units, and distribution of unusual lithic components .........................................
Combined thickness of Muldraugh—Fort Payne-Salem-Warsaw-Harrodsburg units in central to western
Kentucky and probable equivalent Renfro Member of Slade Formation in east-central and southeastern
Kentucky .............................................................................
Thickness of Harrodsburg Limestone in central, south-central, west-central, and western Kentucky . . . .
Cumulative thickness of Salem, Salem-Warsaw, St. Louis, and Renfro (excluding Muldraugh equivalents) units
in Kentucky ...........................................................................
Thickness of St. Louis Limestone and St. Louis Member of Newman Limestone and Slade Formation . .
Thickness of Ste. Genevieve Limestone and probable and inferred equivalents ......................
Thickness and generalized lithofacies of Chesterian rocks and equivalents in Kentucky ...............
Thickness of Kidder Limestone Member of Monteagle Limestone and equivalents in eastern to south-central
Kentucky, and cumulative thickness of Paoli Limestone through Elwren Sandstone and Renault Forma-
tion through Cypress Sandstone in west-central and western Kentucky .........................
Thickness of combined Hartselle Formation and Bangor Limestone in eastern to south-central Kentucky and
thickness of Golconda Formation in west-central and western Kentucky ........................
Thickness of Paragon and Pennington Formations in eastern to south-central Kentucky and combined thickness
of Glen Dean—Tar Springs—Vienna—Waltersburg formational units in west-central and western Kentucky
Generalized thickness of interval from base of Menard Limestone and equivalent units in Leitchfield and Buffalo
Wallow Formations to base of Caseyville Formation .........................................

65. Diagram of stratigraphic occurrences of Mississippian fossil fauna helpful in recognition of map units in Kentucky
66. Diagram comparing major conodont zonal classifications for the Mississippian System in North America .....
67. Map showing quadrangles of Mississippian and Devonian conodont-bearing samples in central, east-central. and south-

central Kentucky .............................................................................

TABLE S

Kentucky 71/z-minute US. Geological Survey geologic quadrangles in which Mississippian rocks are exposed ......
Selected exposures of Mississippian rocks in Kentucky ....................................................
Conodonts from top one inch of New Albany and Sunbury Shales in east-central, eastern central, and south-central Kentucky
Conodonts from basal part of Nancy (New Providence) Member of the Borden Formation and from basal beds of the Fort

Payne Formation in east-central and south-central Kentucky ............................................
Conodonts from basal (New Providence) part of Nancy Member of the Borden Formation in east-central and south-central

Kentucky .......................................................................................
Conodonts from basal part of New Providence Shale Member of the Borden Formation in central Kentucky .......

Page
94
95
95
96
96

97
97

98

98

99

99

106

108

109

Page
8
26
1 10

111

112
112

MISSISSIPPIAN ROCKS IN KENTUCKY

By EDWARD G. SABLE and GARLAND R. DEVER, JR.l

ABSTRACT

Mississippian rocks in Kentucky reflect a history of largely marine
deposition in shallow cratonic basins and on shelves and platforms.
Strata represent the Kinderhookian, Osagean, Meramecian, and
Chesterian Series and their equivalents. Maximum thicknesses of more
than 2,800 feet of Mississippian strata were deposited in western Ken-
tucky in the Eastern Interior basin, and about 2,000 feet in eastern
Kentucky on the margins of the Appalachian basin. Positive struc-
tural features from which Mississippian strata have been largely
eroded and which were alternately submergent and emergent during
the Mississippian were the Cincinnati arch and parts of the adjacent
Waverly arch. The area of the Pascola arch in westernmost Kentucky
was a negative feature during most of Mississippian time. Variations
in depositional thicknesses resulted from differential tectonic move-
ments, from deposition on surfaces of uneven topographic relief, and
locally, from subaerial and submarine erosion. Strata include a wide
variety of oomplexly related detrital, chemical, and biologically derived
sediments. Crustal instability within parts of the region is recorded
by disconformities representing missing strata and by features inter-
preted to represent subaerial conditions, but few multiple hiatuses
due to widespread epeirogenic movements or eustatic sea-level changes
have been recognized within the stratigraphic succession.

Limestone and dolomite compose about two-thirds of the Mississip-
pian rocks in Kentucky. Carbonate deposition, probably encroaching
from the west, reached western Kentucky in Early Mississippian
(Kinderhookian) time. During Osagean time deltaic sediments pro-
graded westward as far as west-central Kentucky, and carbonates,
which were initially restricted to western areas. spread as deposition
of terrigenous detritus waned. Carbonate strata reached their max-
imum extent during Meramecian time. Although areally more
restricted during Chesterian time, carbonate rocks periodically
accumulated over large areas in and beyond the Eastern Interior basin
and Appalachian basin. Evaporites and associated dolomitic strata
were deposited during mid-Meramecian time and marked an episode
of restricted circulation related to cratonic tectonism or eustatic
changes.

Land-derived detrital rocks constitute about one-third of the total
preserved Mississippian succession. The greatest volume of sediment
was contributed from source areas in highlands northeast and east
of the present Appalachians and in the eastern Canadian Shield areas;
in Late Mississippian (Chesterian) time, a possible minor source area
was a highland to the south or southwest of Kentucky. Large volumes
of detritals from northeastern sources were deposited in Kentucky
by major river systems and accumulated as deltaic complexes dur-
ing deposition of the Lower Mississippian Borden Formation and in
Late Mississippian (Chesterian) time. The Cincinnati arch, chieﬂy a
barrier to sediment dispersal, may have contributed small amounts
of detritus during very early Osagean and Meramecian time. Silioeous
and cherty rocks are abundant in strata of Osagean age in southern

Manuscript approved for publication April 20, 1989.
1Kentucky Geological Survey, Lexington, KY.

 

and western Kentucky, and southern or eastern sources may have con-
tributed clay-size detrital silica in these areas.

Marine sedimentary environments were mostly shallow to very
shallow-water neritic over large areas during the Mississippian. but
deeper water perideltaic neritic environments were present in western
Kentucky during Kinderhookian, Osagean, and early Meramecian
time. Seas are surmised to have opened and deepened southward and
westward, with western connections to widespread Mississippian seas
across and north of the Ozarks, and eastern connections to Appala-
chian areas through the Cumberland saddle of south-central Kentucky.
Lower delta-plain environments characterized large areas of Kentucky
periodically during Chesterian time.

In general. westward-deepening, low-energy environments are in-
dicated by the fining grain size of Kinderhookian rocks. and shallower
higher energy environments are indicated by coarser and better sorted
rocks in Osagean and early Meramecian time. Deposits of middle
Meramecian age indicate a low-energy environment, and alternations
of high- and low-energy environments characterize Chesterian time.
Mississippian rocks attest to a mild climatic regimen during the period,
with aridity in mid-Meramecian time, and possibly an intermittently
wet climate during Chesterian time.

N ortheast- and northwest~trending structures, mostly inherited from
Late Devonian time, characterize the Mississippian tectonic frame
work. N ortheast-trending negative structural features responded most
actively to tectonic stresses. Positive structural features such as the
Cincinnati arch and Waverly arch inﬂuenced some Mississippian
depositional patterns. Crosscutting linear features. now expressed as
the Rough Creek and Kentucky River fault systems, were active fault
zones or hinge lines during the Late Mississippian.

Tectonic movements affecting the region during Mississippian time
probably included relative uplift and partial emergence of the Trans-
continental arch west of Kentucky. and corresponding subsidence of
a subparallel northeast-trending trough, the Michigan and Eastern
Interior basins. following Kinderhookian time. Regional southwest-
and west-dipping paleoslopes developed during Osagean time and per-
sisted into Pennsylvanian time. Relative uplift of preexisting positive
structural features resulted in restricted seas during Meramecian time.
At about the close of Chesterian time, a major episode of southward
tilting and general emergence took place, preceding deposition of Penn-
sylvanian sediments. These structural movements were accompanied
by marked beveling of strata near basin margins and on some gentle
uplifts within the basins, and by widespread channel-cutting.

INTRODUCTION

This report summarizes the lithostratigraphy of Mis-
sissippian rocks in Kentucky. It includes descriptive
and nomenclatural discussions and some resulting en-
vironmental and tectonic interpretations.

2 MISSISSIPPIAN ROCKS IN KENTUCKY

Many significant features described and discussed
in this report are referenced to the 1961—1978 GQ
maps or other relatively recent publications, although
some of these features had been recognized by earlier
workers. Where appropriate, the previous work is cited
as well.

HISTORY OF INVESTIGATIONS

Summary reports on Mississippian rocks of Kentucky
and adjoining States include chapters in Miller (1919),
McFarlan (1943), and Cumings (1922), and reports by
Weller and Sutton (1940), Swann (1963), Shaver and
others (1970), Pryor and Sable (1974), Willman and
others (1975), Sable (1979a), and Rice and others (1979).
A sobering review of difficulties and inconsistencies per-
taining to mappability of some formational contacts
bearing on series and other time-equivalent boundaries,
encountered during the 1960—1978 mapping program,
was given by Pohl (1970).

The development of stratigraphic nomenclature of
Mississippian strata in Kentucky and adjacent States
falls into four general stages. During 1854—1900, sys-
temic, series, and formational names for gross units
based on lithology and fauna] content were first intro-
duced. Most stratigraphic names such as St. Louis and
Ste. Genevieve were extended from adjoining States,
although some, such as Knobstone and Mammoth Cave,
were of Kentucky origin.

Between about 1900 and 1930, units were subdivided
and in part renamed by workers such as Stuart Weller,
E.O. Ulrich, and Charles Butts. Intensive systematic
studies of macrofossil taxonomy and its use for correla-
tion purposes marked this period. The Mississippian
and Pennsylvanian outcrop belt along the margins of
the Eastern Interior basin was mapped in relative detail
by cooperative arrangement of State surveys (J .M.
Weller and Sutton, 1940); and meaningful regional cor-
relations, particularly of Upper Mississippian rock
units, were accomplished.

During the period 1930—1960, concepts of time-
equivalent but lithologically different sedimentary rock
facies were stressed. Stockdale (1931, 1939) used these
concepts in his studies of Lower Mississippian rocks in
Indiana and Kentucky. Divisions and correlations of
Upper Mississippian rock units in west-central and
western Kentucky were refined by Stouder (1938, 1941)
and McFarlan and others (1955), in eastern Kentucky
by McFarlan and Walker (1956), and in nearby southern
Indiana by Malott (1952). In northeastern Kentucky
and adjoining States, a pioneer surface and subsurface
regional study incorporating lithologic correlations of
lowermost Mississippian and uppermost Devonian

 

clastic units with interpretations of sedimentary
environments and paleogeography was completed by
Pepper and others (1954). In cooperative arrangement
between the US. Geological Survey and the Common-
wealth of Kentucky, all 71/2-minute quadrangles in Ken-
tucky were mapped topographically on a 1:24.000 scale,
setting the stage for detailed geologic mapping.

Between 1960 and 1978, Kentucky was mapped
geologically on 71/2-minute quadrangles, at a scale of
1:24,000 (Cressman and N oger, 1981). Many refine-
ments were made in Mississippian lithostratigraphic
interpretations and nomenclature. During this period
mappers recognized the Early Mississippian Borden
delta complex and delineated its morphology. Relation-
ships of Upper Mississippian strata in eastern Ken-
tucky were also clariﬁed during the mapping program.
Mapping of major sandstonefilled channels in Chester-
ian rocks in western and westccentral Kentucky led to
development of a deltaic model of deposition like that
previously interpreted for similar deposits in Illinois.
Resource investigations by the Kentucky Geological
Survey of oil, gas, limestone, and clays in Mississippian
rocks accompanied the stratigraphic studies. Finally,
other investigators in academic, industry, and govern-
ment circles used the geologic maps as bases for
research on Mississippian stratigraphic, sedimentation,
fauna], and engineering studies—as evidenced by the
many reports cited in this text.

Since 1978, further studies of Lower and Upper Mis-
sissippian rocks, mostly in eastern Kentucky, have been
accomplished by Kentucky Geological Survey person-
nel and faculty and students from many universities.
These studies continue, using the GQ maps as practical
framework references.

SCOPE AND METHODS

Sources for this report are mainly published surface
information. US Geological Survey Geologic Quad-
rangle (GQ) maps, products of the 1960—1978 US. Geo-
logical Survey and Kentucky Geological Survey
cooperative program, provided most of the basic data.
Borehole data were obtained from published and un-
published reports and Kentucky Geological Survey
files, and included sample-study logs, drillers’ logs, and
geophysical logs interpreted by us and other in-
vestigators. Paleontological determinations and use of
identifiable fossils by field geologists are incorporated
here mostly as practical biostratigraphic tools evaluated
during the mapping program. Other data used in this
report are cited from dissertations, theses, and other
studies which developed concurrently with or subse-
quent to the mapping program, until about 1983. The

INTRODUCTION 3

reader is also directed to the most recent geologic map
of Kentucky (McDowell and others, 1981; McDowell,
1986), which depicts the distribution of Mississippian
rocks and those of other systems.

The bases for changes in Mississippian stratigraphic
nomenclature during the 1960—1978 mapping program,
such as those of the Borden Formation and its
members, and new names for units of Meramecian and
Chesterian age in southern and eastern Kentucky, are
the criteria of mappability of lithologic units and of con-
tinuity of strata as proved or inferred by detailed map-
ping. However, suggested nomenclatural changes for
some Mississippian units based on the above criteria
have not been adopted, and therefore some nomen-
clatural inconsistencies remain. One of these is the con-
tinued use of the name Warsaw in Kentucky; another
is the use of Chesterian formational names over wide
areas within which some named units vary considerably
in lithology. For the most part, however, Mississippian
terminology in Kentucky used in the Geologic Quad-
rangle maps follows the Code of Stratigraphic N omen-
clature (American Commission on Stratigraphic
Nomenclature, 1961, 1970, 1983).

Regional thicknesses of Mississippian rock units,
shown herein in figures 51—64, are based on approx—
imate thicknesses calculated for each Geologic Quad-
rangle (GQ), thicknesses derived from data given in
columnar sections and from measurements based on
map patterns. Data from about 200 surface and sub-
surface stratigraphic sections were also incorporated.
Plotting of lithofacies for individual units or intervals,
in most cases, resulted in rather poor definition of
trends; instead a more specific portrayal of lithic types
such as sandstone, shale, and limestone and their ap-
proximate limits are shown in some thickness maps in
this report.

Field photographs in this report were taken by GR.
Dever, Jr., in 1986 and 1987.

ACKNOWLEDGMENTS

W.W. Hagan, Director of the Kentucky Geological
Survey, and his staff, and AD. Zapp, P.W. Richards,
W.W. Olive, and ER. Cressman, successive Chiefs of
the Kentucky Geology Branch of the US. Geological
Survey, deserve special thanks for the successful im-
plementation and completion of the cooperative map-
ping program.

We wish also to acknowledge the work done by the
many US. Geological Survey field geologists who
mapped Mississippian rocks in Kentucky. G.W. Weir,
R.C. Kepferle, R.D. Trace, J .C. Phﬂley, W.L. Peterson,
R.Q. Lewis, Sr., D.H. Amos, R.C. McDowell, A.R.

 

Taylor, and Benjamin Gildersleeve made special
contributions to the description, measurement, analysis,
and synthesis of Mississippian strata in several Ken-
tucky areas.

GEOLOGIC FRAMEWORK

The centrally located Eastern Interior basin (fig. 1),
also referred to as the Illinois Basin by some investi-
gators (McDowell, 1986), is a major negative element
of the eastern midcontinent region. It encompasses
much of western Kentucky, Illinois, and Indiana, and
parts of Missouri and west-central Tennessee. The
western margin of the larger Appalachian basin trends
southwest through eastern Kentucky. Major positive
cratonic elements within Kentucky are the Cincinnati
arch, including the Jessamine dome and Cumberland
saddle, the east side of the Pascola arch, and the Waver-
ly arch.

Maximum structural relief of the Precambrian sur-
face in the Eastern Interior basin is about 8,500 ft; that
of the Appalachian basin is about 15,000 ft in Kentucky,
but is as much as 20,000 ft farther east (King, 1969).
In the Eastern Interior basin, Mississippian strata are
more than 2,700 ft thick in western Kentucky and about
3,300 ft thick in southern Illinois (Sable, 1979a); in
the Appalachian basin they are as much as 2,000 ft
thick in southeastemmost Kentucky and are more than
7,000 ft thick farther east in southwestern Virginia (de
Witt and McGrew, 1979).

Mississippian rocks crop out near the major posi-
tive structures in Kentucky such as the Cincinnati
arch, Pascola arch, and Waverly arch; a narrow belt
of outcrop also occurs along the west side of Pine
Mountain in southeastern Kentucky (fig. 2). Because
of an extensive cover of Pennsylvanian rocks, however,
most Mississippian strata in Kentucky occur in the
subsurface.

Strata of Mississippian age constitute the bedrock in
about two-thirds of the area of Kentucky, about 25,850
of the total 40,395 mi2. The volume of Mississippian
rocks in Kentucky is about 1,965 mi3, of which roughly
60 percent are marine carbonate and siliceous rocks, and
40 percent are dominantly marine and marginal marine
shale, siltstone, and sandstone.

Faults that displace Mississippian rocks are the east-
southeast-trending Rough Creek—Shawneetown and
Pennyrile fault systems in western Kentucky, and the
colinear east-northeast-trending Irvine—Paint Creek
fault system in eastern Kentucky. Numerous faults also
strike north-northeast to east-northeast from the Mis-
sissippi Embayment of westernmost Kentucky and ad-
joining States. The Kentucky River fault system strikes

86°

 

MISSISSIPPIAN ROCKS IN KENTUCKY

84° 32°

 

 

   
  
  
    

13096»

!

OZARK UPL|F4

‘Q
\ .0.

~ 7 f4

M

N |
CL N>E’>’

 

/\
A KKK ’r ka/Y
~‘___~ 190%“ Vka -» , .
____.\ ..__ ..........
K’T‘rr—‘(z Q‘Y’Kk \(VKV V?) y
36a __ I.” 4
L“~‘~‘« >190} ‘< y
7‘7—1’ 2) *1 \L
q\ 4‘ »‘ NASHVILLE
a e}
ARKANSAS @ $0
to .x i
3’ v r
és‘b/ “if"“wm
q, MIssIsswm \

AlABA MA

 

l WARRIOR BASIN l

Mia’JHH} w“ J)’

   

. QR
W355
044
(law \NMOORMAN? 7

 

,1,

N i4 49 ( EXPLANATION
) K MICHIGAN 4/

-r ‘44,) r—r

Structural form lines on
Paleozoic and Precambrian
surfaces—Hachures point
towards structural depression

  
   

JESSAMINE

éié‘iiﬁé
Positive structural elements
and monocline

FAIRFIELD
BASIN
.4.)

Negative structural elements

HQHV A'lHEl/WM

____&

Faults and fault zones

,

85/?

00 444,0
(5 etc Area of exposed Precambrian
rocks
V; n 100 MILES
awnings: \ l—I—I—I—I—l
f my. rm

\ LARGHNA
* ‘\
\ \

GEORGIA
l

 

 

FIGURE 1,—Generalized structural elements of Kentucky and adjoining areas. Modiﬁed from Sable (1979a).

east-northeastward from the Jessamine dome of central
Kentucky, and extends across eastern Kentucky in the
subsurface. The easterly trending fault systems across
Kentucky are considered to be related to the 38th
parallel lineament (Heyl, 1972), possibly the trace of a
major east-west basement fault. Movement along
major faults shown on figure 1 displaces all Paleozoic
rocks, but locally, earlier movements on these faults
seem to have inﬂuenced at least Late Mississippian
deposition.

Geographic divisions in Kentucky used in this report
are along quadrangle boundaries. Mississippian rocks
crop out mainly in the northeastern, east-central, south-
central, west-central, western part of the central, and
in the western geographic divisions (fig. 3). Mississip-
pian surface exposures have been mapped in more than
360 Kentucky quadrangles (fig. 4 and table 1). In addi-
tion, selected outcrop and borehole sections of Mis-
sissippian rocks were used during report preparation

 

(fig. 5; indexed in Varnes (1979)). The sections
represented range from complete sections of Mississip-
pian strata to partial sections representing rocks of a
single Mississippian series.

Kentucky and surrounding areas were almost con-
tinuously inundated by Mississippian epeiric seas in
which a variety of detrital and chemical sedimentary
rocks were deposited. Terrigenously derived clastics
were mostly deposited in deltas, and were derived from
major source areas east, northeast, and perhaps south
of Kentucky during the Mississippian. Carbonate
deposits accumulated in areas adjoining lobes of detrital
sediments or during times when little detritus was be-
ing transported. Arid climate and restricted circulation
of shallow marine waters during the middle part of
Mississippian time resulted in evaporite precipitation.
In Late Mississippian time rhythmic alternations of
carbonate- and detrital-dominated units characterized
deposition on a relatively stable shelf in western

 

  

INTRODUCTION 5
89° 88° 87° 8'13“ 15° 8‘4“ 313° el2°
l p
. (bf \\ OHIO I \S
INDIANA #3 I; I 45" (

38°
ILLINOIS

37°

 

 

 
 
 

WEST
VIRGINIA

    

 

 

 

 

 

 

 

 

Pennsylvanian rocks

 

Mississippian rocks

9 [37? L TENNESSEE _ i 8’
4 ‘1' ‘ ‘ ‘ ‘ — ‘- - | I | | | l
D 50 100 MILES
EXPLANATION
Quaternary deposits not shown
Intricate outcrop patterns generalized
TK Cretaceous and Tertiary rocks Pre-Mississippian rocks

"""" 0""' Erosional edge of Mississippian rocks
where covered by younger strata

"’ Selected faults and fault zones

 

FIGURE 2.—Distribution of Mississippian. pre-Mississippian, and post-Mississippian rocks exposed in Kentucky.

 

FIGURE 3.—Map of Kentucky showing 7‘/2-minute quadrangle
boundaries and geographic divisions used in this report (C, E, N,
S, W refer, respectively, to central, east, north, south, and west).

Kentucky and to some extent on an unstable shelf in
eastern Kentucky.

The Eastern Interior basin region, which includes
central, west-central, and western Kentucky, contains
the standard type section for the Mississippian System
in North America. The Mississippian System divisions

 

equate with the lower part of the Carboniferous System
of Europe. The Mississippian Series in the United
States include the Kinderhookian (Meek and Worthen,
1861), Osagean (Branner, 1888), Meramecian (Ulrich,
1904), and Chesterian (Worthen, 1860). These four series
have been adopted by the US. Geological Survey and
several State surveys in the Eastern Interior basin. The
Illinois State Geological Survey currently uses a three-
fold series subdivision, the Kinderhookian, Valmeyeran
(combined Osagean and Meramecian), and Chesterian
Series. In south-central and eastern Kentucky, the US.
Geological Survey adheres to an informal two-fold
designation: Lower (Kinderhookian and Osagean age
equivalents), and Upper (Meramecian and Chesterian
equivalents) Mississippian.

Classification of Mississippian and other Paleozoic
strata in the region, other than conventional rock- and
time-stratigraphic designations, are classifications by
Megagroup (Swann and Willman, 1961) and Sequence
(Sloss and others, 1949; Sloss, 1963; ﬁg. 6 of this report).
Although the Megagroup and Sequence concepts are

688.»: Emmancasw :3 how H 293 and H .3 3m. .38an can 332 :sEnmmmmmmmE 5.33 E moﬁwqahcmsv mﬁgosm @33an we maﬁ Novﬁlé 559nm

ms. E 3 ms

m: a: 2m
\ a a: um
am 3 a2 , __

mm SE

.5 i
8

w: a

G mmmwMZZm—H

2, a s a 2 B 5
am F um _ 3 s: m

, 3 mm E mm 3 8 mm 2m ex 3 .m 3

.E
"2.. oz
.1
E 5 a“. E
a

MISSISSIPPIAN ROCKS IN KENTUCKY

<_Z_0~=>

a 3. 3 2 ,m as. 8 a ﬂu _: a 2. a. E a
.5 3 a i i .3 5. m u: , é a .5

e, a a, as a .5 a 5 mm ,, , as 3
.m m 8 3
z z a an 3 a a: Q
) s, 5 . x 202::

a a 8
um i
a
_ 5
<_z_om_> : i .
Emma w z
x
_ 3 3 2, 5 s

E 8
<ZSD7:

 

 

INTRODUCTION

 

.ASTNQ d
.26: $55» US.» 3 .3 $me .8580 98 mayo 8m 35:83 3 N22: hora .839? Saigooo .5360 3033. En» 3.333 .Eﬁwﬁma vies :wmmnmmmmmmﬂz ma macaw

9:60an .mnmE 03:333. 333% Bad 323 525 55c :oEaESwE 8&535 v5 mommy?“ ﬁnﬁﬂmﬂmﬂg «a : mwﬂzﬂx: aﬁcm 65:8 MERE? zxozunevm mo mmzld $5th

mmwmm—ZZN...

(AZ—0:5

9025::

m b
_, S
9.00.. h

= 3.32559. M
32.9.5 1.

Emma , \\

x

<Z<_DZ_

L

(
8:2 9:

omm 0mm 0% 0mm 0mm 05 0mm 0mm

MISSISSIPPIAN ROCKS IN KENTUCKY

TABLE 1.—Kentucky 71/2-minute US. Geological Survey geologic quadrangle (GQ) maps in which Mississippian rocks are exposed
[See figs. 3 and 4 for geographic subdivisions and quadrangle abbreviations]

 

 

 

 

 

Qu dr 1 dr 1 G Y
abbiev‘ilango: Q“:ng e County/State Author Nunflmr publfiezhed
Southwestern Kentucky
Br Briensburg Marshall T.W. Lambert and L.M. 327 1964

MacCary.
FD Fairdealing Marshall, Trigg, Lyon, E.W. Wolfe 320 1964
and Galloway.
Ha-PL Hamlin—Paris Landing Calloway L.V. Blade 498 1966
Hi Hico Calloway and Marshall W.W. Olive 332 1965
LC Little Cypress Livingston, Marshall, and DH. Amos and E.W. 554 1966
McCracken/Ill. Wolfe.
NC New Concord-Buchanan Calloway H.G. Wilshire 313 1964
RC Rushing Creek Calloway, Trigg, and DA. Seeland and H.G. 445 1965
Marshall. Wilshire.
Western Kentucky
Ag Allegre Todd Harry Klemic 446 1965
Al Allensville Todd and Logan/Tenn. Harry Klemic 502 1966
BP Birmingham Point Lyon, Marshall, K.F. Fox, Jr., and W.W. 471 1966
Livingston, and Trigg. Olive.
Bl Blackford Crittenden, Webster. and DH. Amos 873 1970
Union.
Bu Burna Livingston D.H. Amos 1150 1974
Cd Cadiz Trigg K.F. Fox, Jr. 412 1965
Ce Caledonia Tn'gg and Christian G.E. Ulrich and Harry 604 1966
Klemic.
Cl Calhoun McLean and Hopkins W.D. Johnson, Jr., and 1239 1975
A.E. Smith.
CC Calvert City Livingston and Marshall D.H. Amos and W.I. Finch 731 1968
Ca Canton Trigg K.F. Fox, Jr., and DA. 279 1964
Seeland.
CR Cave in Rock Crittenden R.D. Trace 1201 1974
CH Church Hill Christian G.E. Ulrich 556 1966
C0 Cobb Trigg and Caldwell D.A. Seeland 710 1968
Cr Crider Caldwell W.B. Rogers and RD. 1283 1976
Trace.
Cf Crofton Christian and Hopkins T.M. Kehn 1361 1977
Da Dalton Hopkins, Caldwell, Crit- J .E. Palmer 490 1966
tenden, and Webster.
DE Dawson Springs SE Christian and Hopkins R.D. Trace 1365 1977
DW Dawson Springs SW Caldwell and Christian D.E. Hansen 1061 1973
Dx Dixon Webster D.E. Hansen 1293 1976
Dy Dycusburg Crittenden, Livingston, D.H. Amos and W.H. 1149 1974
and Lyon. Hays.
Ed Eddyville Lyon and Caldwell W.B. Rogers 255 1963
El Elkton Todd F.R. Shawe 650 1967
Fe Fenton ’I‘rigg and Marshall R.W. Schnabel and J .S. 317 1964
MacKallor.
Fr Fredonia Caldwell, Crittenden, and W.B. Rogers and W.H. 607 1967

Lyon.

Hays.

INTRODUCTION

9

TABLE 1.—Kentucky 71/2-minute US. Geological Survey geologic quadrangle (GQ) maps in which Mississippian rocks are exposed—Continued
[See figs. 3 and 4 for geographic subdivisions and quadrangle abbreviations]

 

 

 

dr 1 dr 1 GQ Y
ﬁtewﬂio: Q1“inmate‘s e County/State Author Number publeisalried
Western Kentucky—Continued
Gv Glenville McLean and Daviess W.D. Johnson, Jr., and 1046 1972

A.E. Smith.
Go—Br Golconda-Brownfield Livingston D.H. Amos 546 1966
Ga Gracey Trigg and Christian W.H. Nelson and DA. 753 1968
Seeland.
Gr Graham Christian, Hopkins, and T.M. Kehn 765 1968
Muhlenberg.
GR Grand Rivers Lyon and Livingston W.H. Hays 328 1964
Gn Greenville Muhlenberg T.M. Kehn 907 1971
GC-Sh Grove Center- Union J .E. Palmer 1314 1976
Shawneetown
Gu Guthrie Todd/Tenn. Harry Klemic 539 1966
HM Haleys Mill Christian, Muhlenberg, R.D. Trace 1428 1977
and Todd.
Ha Hammacksville Christian and Todd/Tenn. Harry Klemic 540 1966
He Hemdon Christian/Tenn. Harry Klemic 572 1966
HG Honey Grove Christian and Todd Harry Klemic 376 1965
Ho Hopkinsville Christian Harry Klemic 651 1967
J H Johnson Hollow Trigg Harry Klemic, CE. 722 1968
Ulrich and S.L. Moore.
Ke Kelly Christian T.P. Miller 307 1964
Ki Kirkmansville Muhlenberg and Todd T.M. Kehn 1421 1977
La Lamasco Lyon, Trigg, and RD. Sample 608 1967
Caldwell.
Lo Lola Crittenden and RD. Trace 1288 1976
Livingston.
Ma Marion Crittenden and Caldwell R.D. Trace 547 1966
Mo Model Trigg County, Linton W.B. Rogers 409 1965
topographic quadrangle.
Mn Mont Lyon and Trigg P.L. Weis and P.K. 305 1964
Theobald.
Mf Morganfield Union W.D. Johnson, Jr., A.E. 1269 1975
Smith, and G.M. Fairer.
0G Oak Grove Christian/Tenn. Harry Klemic 565 1966
01 Olmstead Todd and Logan G.E. Ulrich 553 1966
On Olney Caldwell and Hopkins R.D. Trace and T.M. 742 1968
Kehn.
PG Pleasant Green Hill Christian W.H. Nelson 321 1964
Pe Pembroke Christian and Todd S.L. Moore 709 1968
PE Princeton East Caldwell R.D. Trace 1032 1972
PW Princeton West Caldwell and Lyon R.D. Sample 385 1965
Re Repton Crittenden D.A. Seeland 754 1968
RS Roaring Spring Trigg and Christian/Tenn. Harry Klemic and GE. 658 1967
Ulrich.
Ro Rosewood Muhlenberg, Todd, and T.M. Kehn 1445 1978

Logan.

10 MISSISSIPPIAN ROCKS IN KENTUCKY

TABLE 1.—Kentucky 71/2-minute US. Geological Survey geologic quadrangle (GQ) maps in. which Mississippian rocks are exposed—Continued

[See figs. 3 and 4 for geographic subdivisions and quadrangle abbreviations]

 

 

 

 

 

$335.53: Quiiiii‘gle OWN/State “the Niger puiﬁiied
Western Kentucky—Continued
Sa Salem Crittenden and RD. Trace 206 1962
Livingston.
Sb Sebree Webster D.E. Hansen 1238 1975
SG Shady Grove Crittenden and Caldwell R.D. Trace and J .E. 880 1971
Palmer.
Sr Sharon Grove Todd and Logan G.E. Ulrich 482 1966
Sh—Ro Shetlerville—Rosiclare Livingston and DH. Amos 400 1965
Crittenden.
Sm Smithland Livingston D.H. Amos 657 1967
Ut Utica McLean, Daviess, and W.D. Johnson, Jr., and 995 1972
Ohio. A.E. Smith.
Wv Waverly Union and Henderson G.M. Fairer 1220 1975
West-central Kentucky
Ad Adairville Logan and Simpson/Tenn. F.R. Shawe 569 1966
A0 Adolphus Allen/Tenn. W.H. Nelson 299 1964
AS Allen Springs Allen and Warren S.L. Moore 285 1963
Al-Db Alton—Derby Meade and Breckinridge D.H. Amos 845 1970
Au Auburn Simpson and Logan H.C. Rainey III 415 1965
BS Bee Spring Edmonson and Grayson Benjamin Gildersleeve 757 1968
BC Big Clifty Grayson and Hardin W C Swadley 192 1962
BP Big Spring Breckinridge, Meade, and W.L. Peterson 261 1964
Hardin.
BN Bowling Green North Warren F.R. Shawe 234 1963
BG Bowling Green South Warren F.R. Shawe 235 1963
Bi Bristow Warren and Edmonson Benjamin Gildersleeve 216 1963
Br Brownsville Edmonson and Warren Benjamin Gildersleeve 411 1965
Ca Caneyville Grayson Benjamin Gildersleeve 1472 1978
and W.D. Johnson, Jr.
Cl Clarkson Grayson E.E. Glick 278 1963
Co—Ct Cloverport-Cannelton Hancock and M.H. Bergendahl 273 1965
Breckinridge/Ind.
Cn Constantine Hardin and Breckinridge E.G. Sable 302 1964
Cu Custer Breckinridge and Hardin D.H. Amos 1367 1977
De Dennis Logan H.C. Rainey III 450 1965
Do Dot Logan/Tenn. FR. Shawe 568 1966
Dr Drake Warren, Simpson, and S.L. Moore 277 1963
Allen.
Du Dundee Ohio G.H. Goudarzi and A.E. 688 1968
Smith.
Dn Dunmor Logan, Muhlenberg, and T.P. Miller 290 1964
Butler.
FR Falls of Rough Grayson, Breckinridge,
and Ohio. W.D. Johnson, Jr. 1360 1977
Fr Franklin Simpson/Tenn. F.R. Shawe 281 1963

INTRODUCTION

11

TABLE l.—Kentucky 71/2-minute US. Geological Survey geologic quadrangle (GQ) maps in which Mississippian rocks are exposed—Continued
[See figs. 3 and 4 for geographic subdivisions and quadrangle abbreviations]

 

 

 

dr 1 dr 1 GQ Y
ﬁievﬁii Quinn}: g e County/State Author Number publeiglried
West-central Kentucky—Continued
Fo Fordsville Hancock, Ohio, and M.H. Bergendahl and 295 1964

Breckinridge. A.E. Smith.
Ga Garfield Breckinridge D.H. Amos 1278 1976
GD Glen Dean Breckinridge and G.H. Goudarzi 836 1970
Hancock.
Gu Guston Meade and Breckinridge J .E. Palmer 1481 1978
Hd Hadley Warren and Butler H.C. Rainey III 237 1963
Ha Hardinsburg Breckinridge D.H. Amos 1232 1975
HF Hickory Flat Simpson and Allen/Tenn. S.L. Moore 420 1965
Ho Homer Logan Benjamin Gildersleeve 549 1966
Ht Horton Ohio W.D. Johnson, Jr. 915 1971
Ir Irvington Breckinridge and Meade D.H. Amos 1331 1976
Ki Kingswood Breckinridge D.H. Amos 1447 1978
Lw Leavenworth Meade D.H. Amos 941 1971
Le Leitchfield Grayson Benjamin Gildersleeve 1316 197 8
Lb Lewisburg Logan H.C. Rainey III and RC. 830 1969
Miller.
Lo Lodiburg Breckinridge and Meade R.K. Hose, E.G. Sable, 193 1963
and DC. Hedlund.
Mc McDaniels Grayson and W.D. Johnson, Jr. 1473 1978
Breckinridge.
Ma Madrid Hardin, Breckinridge, and W.D. Johnson, Jr. 1482 1978
Grayson.
Mt Mattingly Breckinridge/Ind. L.D. Clark and MD. 361 1965
Crittenden, Jr.
Me Meador Warren, Allen, and W.H. Nelson 288 1963
Barren.
Mo Morgantown Butler and Warren Benjamin Gildersleeve 1040 1972
NA-Ma New Amsterdam— Meade/Ind. D.H. Amos 990 1972
Mauckport
NR Nolin Reservoir Edmonson and Grayson Benjamin Gildersleeve 895 1971
01 Olaton Ohio, Grayson, and W.D. Johnson, Jr. and 687 1968
Breckinridge. A.E. Smith.
Pe Petroleum Allen/Tenn. W.B. Myers 352 1964
PR Pleasant Ridge Ohio and Daviess G.H. Goudarzi and A.E. 766 1968
Smith.
Po Polkville Warren and Allen Benjamin Gildersleeve 194 1962
PM Prices Mill Simpson and Logan/Tenn. F.R. Shawe and H.C. 449 1965
Rainey III.
Qu Quality Butler and Logan Benjamin Gildersleeve 673 1968
Re Ready Edmonson, Grayson, and Benjamin Gildersleeve 1263 1975
Butler.
Rd Reedyville Butler, Edmonson, and F.R. Shawe 520 1966
Warren.
Rh Rhoda Edmonson Harry Klemic 219 1963
Ri Riverside Butler and Warren F.R. Shawe 736 1968
R0 Rochester Butler, Ohio, and DE. Hansen 1171 1974

Muhlenberg.

12

MISSISSIPPIAN ROCKS IN KENTUCKY

TABLE 1.—Kentucky 71/2-minute US. Geological Survey geologic quadrangle (GQ) maps in which Mississippian rocks are exposed—Continued
[See figs. 3 and 4 for geographic subdivisions and quadrangle abbreviations]

 

 

 

 

 

$333335. Tiﬂé‘g‘e (MW/SW Nth“ NIEnQber 1.1.3133“
West-central Kentucky—Continued

Rc Rockfield Warren, Logan, and H.C. Rainey III 309 1964
Simpson.

Rm Rome Breckinridge M.D. Crittenden, Jr., and 362 1965

R.K. Hose.

Rs Rosine Ohio, Grayson, and W.D. Johnson, Jr. 928 1971
Butler.

Ru Russellville Logan R.C. Miller 714 1968

Sc Scottsville Allen K.B. Ketner 184 1962

SG Smiths Grove Warren, Edmonson, and PW. Richards 357 1964
Barren.

SL Spring Lick Ohio, Butler, and Benjamin Gildersleeve 1475 1978
Grayson. and W.D. Johnson, Jr.

Su Sugar Grove Butler, Warren, and T.P. Miller 225 1963
Logan.

SH South Hill Butler and Ohio S.L. Moore 1180 1974

SU South Union Logan, Warren, and Harry Klemic 275 1963
Simpson.

Wo Woodburn Simpson and Warren F.R. Shawe 280 1963

Central Kentucky

Br Bradfordsville Taylor and Marion S.L. Moore 1386 1977

Be Bradfordsville NE Marion, Casey, and S.L. Moore 1396 1977
Taylor.

Bo Brooks Bullitt and Jefferson R.C. Kepferle 961 1972

Ce Cecilia Hardin R.C. Kepferle 263 1963

C0 Colesburg Hardin and Bullitt R.C. Kepferle 602 1967

Cr Cravens Bullitt and Nelson W.L. Peterson 737 1968

El Elizabethtown Hardin and Larue R.C. Kepferle 559 1966

Ei Ellisburg Casey S.L. Moore 1397 1977

F1 Flaherty Hardin and Meade W C Swadley 229 1963

FK Fort Knox Jefferson, Meade, Hardin, R.C. Kepferle and E.G. 1375 1977
and Bullitt. Sable.

GS Gravel Switch Boyle, Marion, and Casey S.L. Moore 1506 1978

HG Halls Gap Lincoln G.W. Weir 1009 1972

Ha Hammonville Larue and Hart F.B. Moore 1051 1972

Hi Hibernia Green, Larue, and Taylor S.L. Moore 1352 1976

Ho Hodgenville Larue and Nelson RE. Moore 749 1968

Hw Howardstown Larue, Nelson, and R.C. Kepferle 505 1966
Marion.

HV Howe Valley Hardin R.C. Kepferle 232 1963

Hu Hustonville Casey and Lincoln R.Q. Lewis, Sr., and A.R. 916 1971

Taylor.

J C Junction City Boyle, Lincoln, and Casey L.D. Harris 981 1972

LE Lebanon East Marion S.L. Moore 1508 1978

LJ Lebanon Junction Bullitt, Nelson, and W.L. Peterson 603 1967
Hardin.

LW Lebanon West Marion S.L. Moore 1509 1978

Lo Loretto Nelson, Marion, and W.L. Peterson 1034 1972

Washington.

INTRODUCTION

13

TABLE 1.—Kentucky 71/2-minute US. Geological Survey geologic quadrangle (GQ) maps in which Mississippian rocks are exposed—Continued
[See figs. 3 and 4 for geographic subdivisions and quadrangle abbreviations]

 

 

 

 

 

dr 1 dr 1 GQ Ye
ﬁiewﬁgo; Quanarfelg e County/State Author Number published
Central Kentucky—Continued
Le Louisville East Jefferson R.C. Kepferle 1203 1974
LW-Ln Louisville West and Jefferson R.C. Kepferle 1202 1974
Lanesville
Mg Magnolia Green, Hart, and Lame F.B. Moore 1280 1975
Mi Millerstown Grayson, Hart, and RE. Moore 417 1965
Hardin.
Ne Nelsonville Larue, Hardin, and W.L. Peterson 564 1966
Nelson.
NH New Haven Nelson and Larue W.L. Peterson 506 1966
Pa Parksville Boyle and Casey S.L. Moore 1494 1978
PP Pitts Point Bullitt and Hardin R.C. Kepferle 1376 1977
Ra Raywick Marion, Nelson, and R.C. Kepferle 1048 1973
Larue.
RH-La Rock Haven—Laconia Meade/Ind. C.F. Withington and E.G. 780 1969
Sable.
Sa Saloma Taylor, Marion, and S.L. Moore 1351 1976
Larue.
Sm Samuels Bullitt and Nelson R.C. Kepferle 824 1969
Sh Shepherdsville Bullitt R.C. Kepferle 740 1968
So Sonora Hardin and Larue F.B. Moore 492 1965
Sp Spurlington Marion and Taylor S.L. Moore 1181 1974
St Stanford Lincoln, Boyle, and RR. Shawe and RB. 1137 1974
Garrard. Wigley.
Su Summit Hardin and Grayson F.B. Moore 298 1964
To Tonieville Larue and Hardin RE. Moore 560 1966
Up Upton Hart, Hardin, and Larue RE. Moore 1000 1972
VS-Ko Valley Station— Jefferson and Bullitt R.C. Kepferle 962 1972
Kosmosdale
VG Vine Grove Hardin and Meade R.C. Kepferle 645 1967
South-central Kentucky
Al Albany Clinton and R.Q. Lewis, Sr., and RE. 550 1966
Cumberland/Tenn. Thaden.
Am Amandaville Cumberland, Adair, and A.R. Taylor 186 1962
Russell.
An Ano Pulaski and Laurel H.K. Stager 171 1962
As Austin Allen and Barren S.L. Moore 173 1961
Ba—ON Barthell-Oneida North McCreary J .B. Pomerene 314 1964
Bf-BW Bell Farm and Barthell McCreary and Wayne, J .H. Smith 1496 1978
Southwest Scott/Tenn.
Bl Billows Pulaski, Rockcastle, and N .L. Hatch, Jr. 228 1963
Laurel.
BF Blacks Ferry Monroe and Cumberland Richard Van Horn and 803 1969
W.R. Griffitts.
Bo Bobtown Pulaski R.Q. Lewis, ‘Sr., A.R. 1102 1973
Taylor, and G.W. Weir.
Br Breeding Adair, Cumberland, and A.R. Taylor 287 1964

Metcalfe.

14

MISSISSIPPIAN ROCKS IN KENTUCKY

TABLE 1.—Kentu.cky 71/2-minute US. Geological Survey geologic quadrangle (GQ) maps in which Mississippian rocks are exposed—Continued
[See figs. 3 and 4 for geographic subdivisions and quadrangle abbreviations]

 

 

 

Qu dr 1 dr 1 G Y
abbievila‘go: Quins: g e County (State Author Nunilier published
South-central Kentucky—Continued
Bu Burkesville Cumberland J .M. Cattermole 220 1963
En Burnside Pulaski, McCreary, and A.R. Taylor, R.Q. Lewis, 1253 1975

Wayne. Sr., and J .H. Smith.
Ca Campbellsville Taylor and Adair A.R. Taylor 448 1965
CV Cane Valley Adair, Taylor, and Green C.H. Maxwell and W.B. 369 1964
Turner.
Cn Canmer Hart R.C. Miller 816 1969
Ce Center Green, Hart, Metcalf and R.C. Miller and S.L. 693 1967
Barren. Moore.
Cl Clementsville Casey and Adair A.R. Taylor and R.Q. 1033 1972
Lewis, Sr.
Co Columbia Adair R.Q. Lewis, Sr., and RE. 249 1963
Thaden.
Cp Coopersville McCreary and Wayne R.Q. Lewis, Sr., and A.R. 1315 1976
Taylor.
Cr Creelsboro Russell, Adair, Clinton, R.E. Thaden and R.Q. 204 1963
and Cumberland. Lewis, Sr.
CR Cub Run Hart, Grayson, and CA. Sandberg and CG. 386 1965
Edmonson. Bowles.
CC Cumberland City Clinton, Wayne, and R.Q. Lewis, Sr., and RE. 475 1965
Russell. Thaden.
De Dehner Pulaski R.Q. Lewis, Sr. 909 1971
Du Dubre Cumberland, Monroe, and R.Q. Lewis, Sr. 676 1967
Metcalfe.
Dn Dunnville Adair, Casey, and CH. Maxwell 367 1965
Russell.
Dy Dykes Pulaski J.H. Smith 1197 1974
EF East Fork Metcalfe, Adair, and J .M. Cattermole 413 1965
Green.
Ed Edmonton Metcalfe J .M. Cattermole 523 1966
El Eli Russell, Pulaski, and BE. Thaden and R.Q. 393 1965
Casey. Lewis, Sr.
Eu Eubank Lincoln, Casey, and R.Q. Lewis, Sr., A.R. 1096 1973
Pulaski. Taylor, and G.W. Weir.
Ex Exie Green S.L. Moore 752 1968
Fa Faubush Pulaski and Russell R.E. Thaden and R.Q. 802 1969
Lewis, Sr.
FR Fountain Run Allen, Monroe, and Warren Hamilton 254 1963
Barren/Tenn.
Fr Frazer Pulaski and Wayne R.Q. Lewis, Sr. 1223 1975
Fe Freedom Monroe and Barren S.L. Moore 349 1964
F0 Frogue Cumberland and R.Q. Lewis, Sr. 675 1967
Clinton/Tenn.
Ga Gamaliel Monroe D.E. Trimble 253 1963
GN Glasgow North Barren D.D. Haynes 339 1964
GS Glasgow South Barren S.L. Moore and R.C. Miller 416 1965

INTRODUCTION

15

TABLE 1.—Kentucky 71/2-mirmte US. Geological Survey geologic quadrangle (GQ) maps in which Mississippian rocks are exposed—Continued
[See figs. 3 and 4 for geographic subdivisions and quadrangle abbreviations]

 

 

 

dr 1 dr 1 GQ Y
gamma: Quinn: g e County /State Author Number publizhed
South-central Kentucky—Continued
Gr Gradyville Adair, Green, and A.R. Taylor 233 1963

Metcalfe.
Gb Greensburg Green and Taylor A.R. Taylor, S.J. Luft, 739 1968
and R.Q. Lewis, Sr.
Gs Gresham Green, Adair, and Taylor A.R. Taylor 421 1965
Ha Hail McCreary and Pulaski J .H. Smith, J .B. 1058 1973
Pomerene, and R.G.
Ping.
Hi Hiseville Barren and Metcalfe D.D. Haynes 401 1965
Ho Holland Allen/Tenn. W.H. Nelson 174 1962
HC Horse Cave Barren and Hart D.D. Haynes 558 1966
Hu Hudgins Green and Hart R.G. Miller and S.L. 834 1969
Moore.
Jb J abez Russell and Wayne R.E. Thaden and R.Q. 483 1966
Lewis, Sr.
J a Jamestown Russell and Wayne R.E. Thaden and R.Q. 182 1962
Lewis, Sr.
Kn Kniﬂey Adair C.H. Maxwell 294 1964
Li Liberty Casey A.R. Taylor and R.Q. 946 1971
Lewis, Sr.
Lu Lucas Barren and Aﬂen D.D. Haynes 251 1963
MC Mammoth Cave Edmonson, Hart, and DD. Haynes 351 1964
Barren.
Ma Mannsville Taylor, Adair. and Casey A.R. Taylor 562 1966
Mt Maretburg Rockcastle, Pulaski, and 8.0. Schlanger 338 1965
Lincoln.
MS Mill Springs Wayne. Pulaski, and R.Q. Lewis, Sr. 1057 1972
Russell.
Mi Mintonville Casey and Pulaski R.Q. Lewis, Sr., and A.R. 1198 1974
Taylor.
Mo Monticello Wayne A.R. Taylor 1319 1976
Mp Montpelier Adair and Russell R.Q. Lewis, Sr., and RE. 337 1964
Thaden.
MV Mount Vernon Rockcastle S.O. Schlanger and G.W. 902 1971
Weir.
Mu Munfordville Hart S.L. Moore 1055 1973
Ne Nevelsville McCreary and Wayne J .H. Smith 1326 1976
Pa Park Hart, Barren, and S.L. Moore and DD. 634 1967
Metcalfe. Haynes.
PC Park City Barren, Edmonson, and DD. Haynes 183 1962
Warren.
Pr-SP Parmleysville and Sharp Wayne and A.R. Taylor 1405 1977
Place McCreary/Tenn.
Pn Parnell Wayne R.Q. Lewis, Sr., and SJ. 861 1970
Luft.
Ph Phil Casey, Russell, and CH. Maxwell 395 1965

Pulaski.

16

MISSISSIPPIAN ROCKS IN KENTUCKY

TABLE 1.—Kentucky 71/2-minute US. Geological Survey geologic quadrangle (GQ) maps in which Mississippian rocks are exposed—Continued
[See figs. 3 and 4 for geographic subdivisions and quadrangle abbreviations]

 

 

 

 

 

ﬁﬂi‘ﬁi @232” “WW/SW Authm N33,... puiﬁiied
South-central Kentucky—Continued
Po-PM Powersburg and Pall Wayne and Clinton/Tenn. R.Q. Lewis, Sr. 1377 1977
Mall
RS Russell Springs Russell and Adair R.Q. Lewis, Sr., and RE. 383 1965
Thaden.
Sa-Mv Savage and Moodyville Clinton and Wayne R.Q. Lewis, Sr. 1318 1976
Sy Sawyer Whitley, McCreary, W.P. Puffett 179 1962
Laurel, and Pulaski.
SH Science Hill Pulaski, Casey, and AR. Taylor and R.Q. 1105 1973
Lincoln. Lewis, Sr.
Sh Shopville Pulaski and Rockcastle N .L. Hatch, Jr. 282 1964
So Somerset Pulaski R.Q. Lewis, Sr. 1196 1974
SL Sulphur Lick Monroe, Metcalfe, and L.D. Harris 323 1964
Barren.
SW Sulphur Well Metcalfe and Green J .M. Cattermole 555 1966
SS Summer Shade Metcalfe and Barren W.J. Hail, Jr. 308 1964
Su Summersville Green S.L. Moore 870 1970
TH Temple Hill Barren S.L. Moore and RC. 402 1965
Miller.
To-UH Tompkinsville and Union Monroe, Clay/Tenn. I.J. Witkind 937 1971
Hill
Tr ’Ilracy Barren, Monroe, and S.L. Moore 217 1963
Allen.
Ve-Cl Vemon—Celina Monroe and Cumberland R.Q. Lewis, Sr. 966 1972
Wa Waterview Cumberland and Metcalfe J .M. Cattermole 286 1963
WD Wolf Creek Dam Clinton, Russell, and R.Q. Lewis, Sr., and RE. 177 1962
Cumberland. Thaden.
Wo Woodstock Lincoln, Pulaski, and G.W. Weir and SD. 776 1969
Rockcastle. Schlanger.
Yo Yosemite Casey and Lincoln A.R. Taylor and R.Q. 910 1971
Lewis, Sr.
East-central Kentucky
Al Alcorn Jackson, Estill, and CL. Rice 963 1972
Madison.
Ba Bangor Rowan, Morgan, Menifee, D.K. Hylbert and J .C. 947 1971
and Bath. Philley.
Be Berea Madison, Rockcastle, and G.W. Weir 649 1967
Garrard.
Bn Bemstadt Laurel and Rockcastle N .L. Hatch, Jr. 202 1963
Bh Bighill Madison, Jackson, and G.W. Weir, K.Y. Lee, and 900 1971
Rockcastle. P.E. Cassity.
Br Brodhead Rockcastle, Lincoln, and J .L. Gualtieri 662 1967
Garrard.
CC Clay City Powell and Estill G.C. Simmons 663 1967
Ch Cobhill Estill, Lee, and Powell D.C. Haney 1347 1976
Co Colfax Bath and Fleming R.C. McDowell 1332 1976
CO Crab Orchard Lincoln J .L. Gualtieri 571 1967

INTRODUCTION 1 7

TABLE 1.—Kentucky 71/2-minute US. Geological Survey geologic quadrangle (GQ) maps in which Mississippian rocks are exposed—Continued
[See figs. 3 and 4 for geographic subdivisions and quadrangle abbreviations]

 

 

 

dr 1 Qu dr 1 G Y
ﬁgevilago: almazg e County/State Author Nuifber published
East-central Kentucky—Continued
Ez Ezel Morgan and Menifee G.N. Pipiringos, SC. 721 1968

Bergman, and VA.
Trent.
Fa Farmers Rowan, Fleming, and R.C. McDowell 1236 1975
Bath.
Fr Frenchburg Menifee and Powell H.P. Hoge 1390 1977
He Hedges Clark D.F.B. Black 1235 1975
Hi Heidelberg Lee, Estill, and Owsley D.F.B. Black 1340 1977
Iv Irvine Estill H.P. Hoge, P.B. Wigley, 1285 1976
and F.R. Shawe.
Jo Johnetta Rockcastle and Jackson J .L. Gualtieri 685 1968
La Lancaster Lincoln and Garrard G.W. Weir 888 1971
Le Leighton Estill, Jackson, and Lee D.C. Haney and CL. 1495 1978
Rice.
Lv Levee Montgomery, Clark, and R.C. McDowell 1478 1978
Powell.
Li Livingston Rockcastle, Laurel, and W.R. Brown and M.J. 1179 1974
Jackson. Osolnik.
Mc McKee Jackson and Owsley G.W. Weir and MD. 1125 1973
Mumma.
Me Means Menifee, Montgomery, G.W. Weir 1324 1976
and Powell.
Mo Morehead Rowan H.P. Hoge and J .R. 1022 1972
Chaplin.
01 Olympia Bath and Menifee R.C. McDowell and G.W. 1406 1977
Weir.
PL Paint Lick Garrard, Madison, and G.W. Weir 800 1969
Lincoln.
Pa Palmer Estill, Clark, Madison, G.C. Simmons 613 1967
and Powell.
Pn Panola Estill and Madison R.C. Greene 686 1968
Pt Parrot Jackson and Laurel D.F. Crowder 236 1963
P0 Pomeroyton Wolfe, Menifee, and G.W. Weir and P.W. 1184 1974
Morgan. Richards.
Pr Preston Bath and Montgomery G.W. Weir and R.C. 1334 1976
McDowell.
SL Salt Lick Rowan, Bath, and J .C. Philley 1499 1978
Menifee.
Sa Sandgap Jackson J .L. Gualtieri 1100 1973
Sc Scranton Menifee D.C. Haney and N .C. 1488 1978
Hester.
Sl Slade Powell, Menifee, and G.W. Weir 1183 1974
Wolfe.
St Stanton Powell and Estill G.W. Weir 1182 1974
Ty Tyner Jackson, Clay, and Laurel G.L. Snyder 247 1963
Wi Wildie Garrard and Rockcastle J .L. Gualtieri 684 1968
Za Zachariah Estill, Lee, Powell, and D.F.B. Black 1452 1978

Wolfe.

18

MISSISSIPPIAN ROCKS IN KENTUCKY

TABLE 1.——Kentucky 71/2-minute US. Geological Survey geologic quadrangle (GQ) maps in which Mississippian necks are exposed—Continued
[See figs. 3 and 4 for geographic subdivisions and quadrangle abbreviations]

 

 

 

 

 

Qu dr 1 dr 1 G Y
abbgeviiiftgio: Quins? e County/State Author Nun?ber pubﬁzhed
Northeastern Kentucky
Br Brushart Greenup and Lewis C.S. Denny 324 1964
Bu Burtonville Fleming and Lewis R.H. Morris 396 1965
Ch Charters Lewis R.H. Morris 293 1965
Cc-BV Concord-Buena Vista Lewis R.H. Morris 525 1966
Cr Cranston Rowan, Fleming, and J .C. Philley, D.K. 1212 1975

Lewis. Hylbert, and HP. Hoge.
Fl Flemingsburg Fleming and Mason J .H. Peck 837 1969
Fr Friendship Lewis and Greenup R.L. Erickson 526 1966
Ga-PR Garrison and Pond Run Lewis and Greenup, J .R. Chaplin and GE. 1490 1978
Scioto/Ohio. Mason.
Gr Grahn Carter K.J. Englund 1262 1976
Gs Grayson Carter C.L. Whittington and 640 1967
J .C. Ferm.
Gp-Ir Greenup—Ironton Greenup and Boyd Ernest Dobrovolny, J .C. 532 1966
Ferm, and SO. Eroskay.
HG Head of Grassy Lewis R.H. Morris 484 1966
Hi Hillsboro Fleming and Bath J .W. Mytton and RC. 876 1970
McDowell.
Lo Load Greenup J .A. Sharps 519 1966
MI Manchester Islands Lewis J .H. Peck and KL. 581 1966
Pierce.
01 Oldtown Greenup and Carter C.L. Whittington and 353 1965
J .C. Ferm.
OH Olive Hill Carter K.J. Englund and J .F. 1270 1975
Windolph, Jr.
PL Plummers Landing Fleming and Rowan R.C. McDowell, J .H. 964 1971
Peck, and J .W. Mytton.
Po-Wh Portsmouth- Greenup. Scioto/Ohio R.A. Sheppard 312 1964
Wheelersburg-New
Boston
So Soldier Carter, Lewis, and J .C. Philley, D.K. Hylbert, 1233 1975
Rowan. and HP. Hoge.
St Stricklett Lewis, Fleming, and RH. Morris 394 1965
Rowan.
To Tollesboro Lewis and Fleming J .H. Peck 661 1967
TV Tygarts Valley Greenup and Carter R.A. Sheppard 289 1964
Va Vanceburg Lewis/Ohio R.H. Morris and KL. 598 1967
Pierce.
We Wesleyville Carter and Lewis J .C. Philley and J .R. 1305 1976
Chaplin.
Eastern Kentucky
Au Ault Elliott, Carter, and AD. DeLaney and K.J. 1066 1973
Rowan. Englund.
Br Bruin Elliott and Carter K.J. Englund and AD. 522 1966
DeLaney.
Ha Haldeman Rowan, Carter, and SH. Patterson and J .W. 169 1961

Elliott.

Hosterman.

INTRODUCTION

19

TABLE 1.—-Kentucky 7V2-minute US. Geological Survey geobgic quadrangle (GQ) maps in which Mississippian rocks are exposed-Continued
[See figs. 3 and 4 for geographic subdivisions and quadrangle abbreviations]

 

 

 

 

 

dr 1 dr 1 GQ Y
gamma: Gluing? e County/ State Author Number publizlied
Eastern Kentucky—Continued
Is Isonville Elliott, Morgan. and K.J. Englund and AD. 501 1966

Lawrence. DeLaney.
SH Sandy Hook Elliott and Morgan K.J. Englund and AD. 521 1966
DeLaney.
WL West Liberty Morgan K.J. England, J .W. 589 1967
Huddle, and AD.
DeLaney.
Wr Wrigley Morgan, Rowan, and J .W. Hosterman, SH. 170 1961
Elliott. Patterson, and J .W.
Huddle.
Southeastern Kentucky
Ba Balkan Bell and Harlan A.J. Froelich and J .F. 1127 1973
Tazelaar.
Be-Ap Benham and Appalachia Harlan and Letcher, A.J. Froelich and ED. 1059 1973
Lee/V a. Stone.
Bl Bledsoe Leslie and Harlan Bela Csjetey, Jr. 889 1971
EC Elkhom City—Harman Pike/Va. D.C. Alvord and R.L. 951 1972
Miller.
Ev Evarts Harlan J .F. Tazelaar and W.L. 914 1974
Newell.
Ew Ewing Harlan and Bell/Va. K.J. Englund and others 172 1961
Fr-Ea Frakes—Eagen Bell and Whitley W.L. Newell 1249 1975
HS Hubbard Springs Harlan J .F. Tazelaar and W.L. 914 1974
Newell.
JE Jellico East Whitley C.L. Rice and W.L. Newell 1264 1976
He—Cw Hellier and Clintwood Pike/V a. DC. Alvord 950 1971
Ha Harlan Harlan A.J. Froelich and E.J. 1015 1972
McKay.
He Helton Leslie, Harlan, and Bell D.D. Rice 1227 1975
JE Jenkins East Pike and Letcher D.E. Wolcott 1210 1974
JW Jenkins West Letcher and Pike/V a. C.L. Rice 1126 1973
Kj-FR Kayjay and Fork Ridge Bell and Knox. Campbell C.L. Rice and E.K. 1505 1978
and Claibome/Tenn. Maughan.
Lo Louellen Harlan. Letcher, and A.J. Froelich 1060 1973
Perry.
MN Middlesboro North Bell K.J. Englund. J .B. Roen, 300 1964
and AD. DeLaney.
MS Middlesboro South Bell/Tenn. and Va. K.J. Englund 301 1964
No Nolansburg Harlan, Letcher. and Bela Csjetey, Jr. 868 1970
Leslie.
Pi Pineville Bell and Knox A.J. Froelich and J .F. 1129 1974
Tazelaar.
RH Rose Hill Harlan E.K. Maughan and J.F. 1121 1973
Tazelaar.
R0 Roxana Harlan and Letcher E.K. Maughan 1299 1976
Va Varilla Bell and Harlan/V a. K.J. Englund, E.R. Landis, 190 1963
and H.L. Smith. '
WC Wallins Creek Harlan and Bell A.J. Froelich 1010 1972
Wh-FG Whitesburg and Flat Gap Letcher/V a. 0L. Rice and DE. Wolcott 1119 1973

 

20 MISSISSIPPIAN ROCKS IN KENTUCKY

 

 

 

 

 

 

 

 

 

 

 

   

 

 

 

 

 

 

 

 

 

 

 

 

 

         

 

 

 
  
    

 

 

 

 

 

 

 

 

   

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

2 E
WEST OR ILLINOIS KENTUCKY SERIES )u—J I“
SOUTH u, D
> o
a) Lu
(I)
z' 31;;
2 U)
a 28
Sub-Absaroka
unconformi
ty CHESTER EXPLANATION
E .. ..
I ‘~ . -_ E Sandstone
n.
St'e.Gia'rie:/.i-(;\'/e,I I I I I I ; I MERAMEC g
St. Louis, Salem, Ullin I L-L (7; 5
unconfo m' '_— ‘9
_’:Y _ M_A_MMOLH E 3 Shale and mudstone
I 1 FT 1 'I—[ CAIVEI Tl—I | OSAGE U)
3 I H I I I I I ‘1—1— 5 E
__ _¥ _ '_ _-I=
— — — KINDERHOOK “mm“
UPPER Dolomite
2
1‘
Z
MIDDLE g
Sub-Kaskaskia Sub-Kaskaskia I B A
uncorliformity unconformity I g
D.
I LOWER E)
o.
a
I I [ I I I 1 EL
I [TiIIIIIIIIIII 8
I 7 I I L I CAYUGAN- z z
I I ' I I .Hqu/m/YI , I r I NIAGARAN s <
I 1 l I I I [I I I I I I ﬂ g 8
#1 I I I I I I I l I d it
4 I l I I I I I I . U) —
VA'II'IIIII'I‘I'I4v'f :- ‘~ I.
I I I I I I I IiI l Exe=~_ .. ALEXANDRIAN

 

 

FIGURE 6.—Pre-Pennsylvanian—post—Ordovician Megagroups (capitalized italic) and Sequences in the Eastern Interior basin and adjoining
areas (modified from Swarm and Willman, 1961).

valuable in establishing broad lithogenetic, tectonic, and
time frameworks within which Mississippian events can
be placed, this report discusses Mississippian history
in detail that requires conventional, specific rock and
time divisions.

The following discussion on stratigraphy of the
Mississippian System in Kentucky stresses lithostratig-
raphy, and includes discussion of depositional environ-
ments, and paleotectonic implications.

Interstate and intrastate correlations of the Missis-
sippian strata of Kentucky based on reports by Pryor
and Sable (1974) and by Rice and others (1979) are
shown in ﬁgures 7 and 8. Although regional correlations
are principally lithostratigraphic, based on field map-
ping, microfauna and macrofauna have been useful both
in a lithostratigraphic sense and as tools in defining
time-unit boundaries. (See section, “Paleontology.”)

 

Representative sections of Mississippian rocks and
an interpretative cross section showing vertical and
lateral relationships constitute figure 9.

PHYSIOGRAPHY, OUTCROP, AND
SCENIC FEATURES

Mississippian strata underlie four of Kentucky’s six
principal physiographic regions: The Knobs, Mississip-
pian Plateau, Western Coal Field, and Cumberland
Plateau (fig. 10). The Knobs region is a narrow, arcuate
belt of conical hills around the periphery of the Blue
Grass Region of central Kentucky, a lowland underlain
by early Paleozoic rocks. The Knobs are erosional rem-
nants, composed mostly of Lower Mississippian
(Osagean) shale and siltstone, standing along the front

INTRODUCTION 2 1

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

  

 

 
 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

Eastern Interior Basin Appalachian Basin
E: g lllinois Kentucky
"7 a:
>‘ an East Central and SC, EC, NE S th
to Southeastern Western West Central 30"”1 Central Northeastern (Ettensohn and Wiggly/1235,3331)
(as mapped) others, 1984)
. Pennington
Pennington *Paragon Formation
Formation Formation 2
: Upper
'E Bangor member
‘3 (See fig. 8) Limestone
a:
J: Hartselle *Several
0 Formation “’35:; members
w (See fig. 8)
”Kidder :
Limestone 2
g Member 3
'Leyc‘as Lgrgiestone 'Levias ‘3 E
em r o Limestone 0 ‘1’
._ Renault Limestone E g Member .§ 5 I 5 5 Tfﬂgegﬁbder :19
g Aux Vases 3 g Rosiclare . 7.3 E 3;; 8 Greenbrier
:3 Sandstone 5 $ Sandstone Ste.Genevreve ‘6: Ste. Genevieve g Ste. Genevieve g ._E_ Formation
q g Member Limestone g Limestone Limestone c -J
Ste. 24 ‘Fredonia E Member Member ‘5 g
c Genevieve m Limestone g I g E
.3 Limestone Member l E 5
o * 2
E 3%; Uppg; I Lower member
._ - .. mem r _
a: .St- L°”'s 3 3; St. Louis Limestone
2 Limestone ... E sr. Louis Member
W— Lower Limestone
E -‘ member I
E
... Salem *Renfro
«n Salem . Salem
.3 Limestone leestone and Member
_v_: Warsaw
2 Ullin LWarstaW Harrodsburg Formations I
Limestone __"12531‘2._,_mean_r _______ ______|.. _____ _._q ______
Fort Payne Fort Payne ’Muldraugh I Fort Pa ne
Formation Formation Member Egﬁfl‘ngg S l Cher);
a: ______I—
Floyds :nob J‘Wilcblie E
: Be em er O ..
o Nada
'ﬁ Holtsclaw I’Halls Gap : Member
I: E Siltstone Member *3
8 .8 8
g :: *Na ncy *Cowbell
o is; Member ’Nancy Member Grainger
_ no Member Formation
Q)
3 *Kenwood *Na ncy
3 Siltstone Member
Member
Springville New Providence New Providence ”Farmers
Shale Shale shale Member J Member
________ s
5 Ch Rockford E112}?
.- outeau Li to
‘3 Limestone (locarl'?es reggnt) >‘
g V P g g (l; Berea g
B . “‘13.: Sandstone o a
1, (Maury'Formation 230) g;
.E Hannibal Hannibal equwalent) Bedford 33 -=
5‘ Shale Shale Shale E m
_ '—"’ _____________ "— o — —
o =
> a New Albany Chattanooga New Albany Chattanooga -
.82 Group Shale Shale Shale 31“ng

 

 

 

 

 

 

 

 

 

FIGURE 7.—Stratigraphic nomenclature of the Mississippian System in Kentucky and Illinois. Asterisk (*) indicates type section in Kentucky.
Modified from Pryor and Sable (1974) and Rice and others (1979).

of the Cumberland or Pottsville Escarpment across Muldraugh Hill is a limestone-capped east-facing es-
east-central Kentucky and along the front of Muldraugh carpment forming the east border of the Mississippian
Hill across west-central and south-central Kentucky. Plateau region.

22

MISSISSIPPIAN ROCKS IN KENTUCKY

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

  

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

E Eastern Interior Basin I Appalachian Basin
‘0
a .2 East Central
> a: Western West Central South Central and Southeastern
<0 w Northeastern (Pine Mountain)
llllllllllllllllll lllll ll'll '(illlll lllllllll llll; \
Grove Church \
Shale \
Kinkaid _? \
Limestone LIttle
Stone Gap \ South Central, East Central,
Degonia Membe; \ and Northeastern
Sandstone ; (Ettensohn and others, 1984)
2 \
Clore Tc 5 .
Limestone 3 kg '2 c Pennington
2 E 9.0 Formation
Palestine 3'2 3 2 g 0 Upper shale member
Sandstone =|L g g c g
M d 5” 33 .g a
Limeegidne ‘ E (E Limestone member
7 .9. 8 .
Waltersburg Stony Gap 5 3 Clastlc member
Sandstone Sandstone a ‘3
V' Merr71ber ; g
Ienna o. a
. - 0 Lower dark
Limestone 7 Upper ‘ shale member
— member
E g Sandstone ' *Poppin Rock Member
Sf '3 ”Glen Dean
:1”; ; Limestone
g Lg) ’Hardinsburg B ”Madmiﬁtirfnm
'- angor
E Sandstone Limestone t:
a *Ramey Creek
Haney .. Member
Limestone Member §
Golconda I339 Clifty Hartselle 0 g 3 *Tygarts Creek
Formation Sandstone Member Formation o 2 333 Member
Beech Creek g 2 <3 E .
Limestone a; g " lg ’Armstrong HIII
Member 3 Upper ,g Lower Member
Cypress Elwren .. 3 member 2' member “Cave Branch Bed
Sandstone Sandstone 2 3 g E
. g E E g : *Mill Knob Member
Reelsvrlle o E 3 o (u
Limestone g g o g z E
._ _‘ Q!
Paint Creek ”Sample 5 ‘5 g g 3 'Warix Run Member
Limestone Sandstone I: E 9 3 3
(Shale) 9 LE 9 g ,\r
Beaver Bend 5 1‘ /
Limestone 5 3 /
1:
’Bethel Mooretown 52 /
Sandstone Formation ‘ /
Renault Paoli /
Limestone Limestone

 

 

 

 

FIGURE 8.—Stratigraphic nomenclature of the Chesterian Series and equivalent strata in Kentucky. Asterisk (*) indicates type section in
Kentucky. Based on Pryor and Sable (1974).

The Mississippian Plateau region is a broad, arcuate
belt around the Western Coal Field. It extends eastward
to the Pottsville Escarpment where the Mississippian
outcrop becomes a narrow belt across east-central and
northeastern Kentucky. The region consists of two prin-
cipal parts divided by the Dripping Springs Escarp-
ment, an east- and south-facing ridge generally capped
by the Big Clifty Sandstone Member of the Golconda
Formation (Chesterian). The outer, eastern part is a
broad plain of low relief, locally with karst development

where underlain by St. Louis and Ste. Genevieve
Limestones (Meramecian); the western portion of this
plain is commonly called the Pennyroyal or Pennyrile.
The inner, western part of the Mississippian Plateau
region is a dissected upland of moderate relief underlain
by limestone, sandstone, and shale (Chesterian), local-
ly called the Sand Hills country.

The Western Coal Field, part of the Eastern Interior
basin, is a gently rolling to hilly upland underlain by
Pennsylvanian (Morrowan-Virgilian) shale, sandstone,

EAST

WE ST

A"

supIuap —

qlnouisuod

Limos

EEEI

 

 

 

 

  
    
 

FEET
100

23138 —

UBBJE) Buumog

   

IUD MILES

 

  

 

 

FORMATION

LEE

Ban I or Limestone FURMA TIUN

artselle Formation

SIUI'IE

 

I Monteagle Lime

 

 

55%

 

 

 

 

 

 

E
“III

     

     
     
 
  
   

II??-

min!" “I“ “\
Ian-I i JIII‘I'IIIIII'“
I\i\\ H‘IIIIIIIIIHII ‘

 

 

    
      
    
 

0/770

      

   

IIII I

KnIery
Sandstone
Member

Cane Valley
Limestone

Member

    
 

 

 

 

 

INTRODUCTION
E a a a c.
L I I 1 I l J .5. E
NUILVWHDi NUlSNINNJd 3N01§3WIT NVWMEN IN:I H39NIVH9 g g '5 E
,. .‘ ‘ “I li° . E a = 2: E .2
. N c = .— w .t.‘
' ~ W I:.;-.~'~ % .2 g g ‘E g
1: ._
, I I l5 s 3 g E. 1: S
I / II 122 B 3: 3 .2”: '3
/ II ”-6 2 m 0 Lu 0 O
I , M Le a
I / a: ’I/ + I
/ EII I
I / ﬁII ‘I' I
I / t I + I
,/ .
/ =
1 ymsm'unv- I m H I E {E E
II 3 = _:
II '33 Z E 3'? 137°
9, ﬁ-
2 = 8 a w 2 we»:
*- 3 .8 3 85 85 =t-E
‘2‘ E e 3.: g a a? 3.3 E
o o u u -— E = E: I.
S '5 '3 E E 5.": 3: 0.9
Q. Q Q (I) < < O
X
LLI |

   

.2?
ﬂ
.1:
(I!
m
5
O
.5
E
=
f:
m

3,2. Sandstone
Siltstone
Shale

 

 

_L .1. _n_ Calcareous shale
Limestone

_L__|_.|
_L__.L_._I.
I
I
I
I

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

     

South
A’ Portsmouth

  
 

Bowling Green

New Providence Member

limestone

 

 

UJSWSW

ENOLSJWII

 

 

awoisawn I Moukum

 

 

 

. _. I . 2H vmuuaua smm lS waws MVSHVM mm '13
h g':_ 332-} {.31 :5: '._':j':.';- 'j-ﬁ-j; '81 mm awoisawn I anmsawn I 51 NDIiVWHOi
-..-. .-,-- :-:-‘ --. ~3N39 '31 smm ‘13 waws mm 3mm: 13
:1 w m :7 ‘3 4 In 2 ‘5 < - z x I z r—
— c: g z: I z u: 4 n: r: m w m 5.1 A
§§§§§§§3§ SEEEEEE
2 _, E E > % z E 3 g m ,_ gm m
< w 13: a:
m : u: 5 E 8 ‘-’ E g
‘g‘ >5 § l s
I I T I I I I I |
a g g L ; 1
a E’ m S g

 

 

 

23

FIGURE 9.—Mississippian rock units and relationships in Kentucky. All scales approximate; modified from Rice and others (1979, p. F6).

24

  

/”/.y;,«,P/ raging»? ,
*S I sis I‘P P)
U 50
L J I I I l

 
 

  

MISSISSIPPIAN ROCKS IN KENTUCKY

 
   
    
     
  

“A
100 MILES ~. 24" e9“
I ON

 

FIGURE 10.—Physiographic diagram of Kentucky (modified from Lobeck, 1929).

and economically important coal. The resistant thick
Caseyville Formation of Pennsylvanian (Morrowan) age
forms ridges and cliffs along parts of the coal field
border, and these make up a subdued and intermittent
topographic escarpment in that area.

In eastern Kentucky, the Cumberland Plateau is
an intricately dissected upland of V-shaped valleys
underlain by Pennsylvanian (Morrowan-Virgilian) shale,
sandstone, and coal. It is bordered on the west by
the Pottsville Escarpment, a northwest-facing ridge
made up largely of Mississippian limestones and silt-
stones, and capped by sandstone of the Upper Missis-
sippian and Lower Pennsylvanian Lee Formation.
Relief increases abruptly in southeasternmost Ken-
tucky. Pine and Cumberland Mountains, two northeast-
trending ridges with relief as much as 2,200 ft, consist
largely of thick Mississippian clastic and carbonate
rocks capped by very thick Lower Pennsylvanian
(Morrowan) sandstone and underlain by Devonian
organic shale.

Weathering and erosion of Mississippian rocks in
Kentucky have produced a variety of scenic features of
interest to both laymen and geologists (McFarlan,
1958). The Mississippian Plateau of western Kentucky
is an internationally known classic karst region, with
a broad sinkhole plain and extensive systems of under-
ground drainage and caverns, including the Mammoth
Cave—Flint Ridge cave system (Livesay, 1953), devel-
oped in the St. Louis, Ste. Genevieve, and Girkin Lime-
stones. In eastern Kentucky, limestone caves are locally
developed in the Monteagle and Newman Limestones

 

and Slade Formation along and northwest of the Potts-
ville Escarpment. These include the Carter Caves
(McGrain, 1954) in northeastern Kentucky and the
Sloans Valley cave system (Malott and McGrain, 1977)
near the southern border of the State.

The conical hills of The Knobs standing above the
outer Blue Grass Region lowlands form a strikingly
unique example of an erosional landscape, remnants of
the uplands left behind the retreating Muldraugh Hill
and Pottsville escarpments (McGrain, 1967).

Mississippian rocks are well exposed in many road-
cuts along highways in Kentucky such as the Bluegrass,
Western Kentucky, Cumberland, and Mountain Park-
ways, U.S. Interstate 65 south of Louisville, U.S. Inter-
state 75 south of Berea, and Interstate 64 east of the
Licking River. These and other exposures are described
in several field conference guidebook articles, for exam-
ple, Weir (1970); Ohio and Kentucky Geological Soci-
eties (1968); Sable and Peterson (1966); Smith and
others (1967); Lewis (1963); Ferm and others (1971);
Dever and others (1977); Lewis and Potter (1978); and
Ettensohn and Dever (1979). Locations of selected ex-
posures of Mississippian strata and information about
them are shown in figure 11, and keyed to quadrangle
maps and referenced publication information in table 2.

DEVONIAN AND MISSISSIPPIAN ROCKS

Units discussed in this section are those which
underlie known Early Mississippian strata and are:

25

DEVONIAN AND MISSISSIPPIAN ROCKS

.ﬁ 033 E 2:533 3 v93“— Nvlﬁ 333:5 man—Same .850 v5. 63:9. .3533: 305% .bqosuqavﬁ E maven gin—BEES: «a $58an goo—om we mnoﬁauoaléa EDGE

 

L!

K“ m4 54 Hr \\

       

       

 

          

ﬁ/JJ 02E ”3thva
w é , ., ,
N. ﬁg _, :
Vﬂﬂ‘f\(\

<

/
COC»®
O > ~

  

   

 

OLODmCAvBO Q

\ 1/
Zr(h\\
\

t\

 

$22 8— am a

26

MISSISSIPPIAN ROCKS IN KENTUCKY

TABLE 2.—Selected exposures of Mississippian rocks in Kentucky

[Numbers of sections keyed to map. ﬁg. 11]

 

Stratigraphic
units exposed

US. Geological Survey
Geologic Quadrangle Maps

Location and access

Features

Pertinent references

 

Northeastern Kentucky

 

 

 

1. Lee Portsmouth GQ—312 South Portsmouth section. Gradational contacts. Nada Calvert, Bernhagen, and
Borden (Sheppard, 1964); South on US. Hwy. 23, red and green shales. others (1968).
Nada Friendship GQ—526 southeast on Ky. Hwy. 7.
Cowbell (Erickson, 1966). west on blacktop, north
Nancy at first junction.

2. Borden Vanceburg GQ—598 Apple Tree Hill section, Ky. Bedding features Ohio Geological Society—
Nancy (Morris and Pierce. 1967). Hwy. 59. Zoaphycas (Taonurus) Geological Society of
Farmers fucoid in Borden. Kentucky (1968. p. 56-57).

Sunbury
Berea
Bedford
3. Lee and Breathitt Grahn GQ-1262 (Englund, Numerous roadcuts and Detrital rocks of Bedford, Philley (1970); Farm and
Paragon 1976); Olive Hill GQ-1270 outcrops along Interstate Sunbury. Borden, Carter others (1971); Ettensohn
Carter Caves (Englund and Windolph. Hwy. 64 between mile- Caves, Paragon. Complex and Dever (1979,
Slade 1975); Soldier GQ—1233 post 132 (east of Licking carbonaterock sequence p. 84-162); Chaplin (1980);
Borden (Philley and others, 1975); River) and US. Hwy. 60 in Renfro and Slade. Up— Ettensohn (1981).
Renfro Cranston GQ-1212, interchange. per and lower systemic
Nada (Philley and others, 1975); boundaries. Nature of
Cowbell Morehead GQ-1022 (Hoge upper boundary (grada-
Nancy and Chaplin, 1972); tional vs. erosional) is
Farmers Farmers GQ-1236 controversial.
Sunbury (McDowell, 1975).
Bedford
Ohio
East-central and southeastern Kentucky
4. Lee and Breathitt Clay City GQ-663 (Simmons, Mountain Parkway from Solutiorrcollapse features in Dever (1971).
Paragon 1967); Stanton GQ-1182 Interchange with Ky. Slade Formation filled
Slade (Weir, 1974a); Slade Hwy. 1057 eastward to with shale and sandstone
Borden GQ-1183 (Weir, 1974b). 2.1 mi east of Slade of lower tongue of
Renfro interchange with Ky. Breathitt. Representative
Nada Hwy. 11. Slade and Breathitt.
Cowbell Complete Slade section is
Nancy present along the
New Albany parkway.
5. Lee and Breathitt Stanton GQ—1182 (Weir, Numerous roadcuts and Erosional features and Dever and others (1977).
Paragon 1974a); Means GQ-1324 adjacent quarry sections depositional patterns in
Slade (Weir, 1976); Frenchburg along Ky. Hwys. 15, 213, Slade and Paragon related
Borden GQ—1390 (Hogs, 1977b); 36, 1274. 801, and US. to intra-Mississippian
Renfro Scranton GQ—1488 Hwy. 460. tectonism.
Nada (Haney and Hester,
Cowbell 1978); Ezel GQ-721 (Pipi-
Nancy ringos and others, 1968);
Farmers Bangor GQ-947 (Hylbert
New Albany and Philley, 1971); Salt
(also Sunbury Lick GQ-1499 (Philley,
and Bedford) 1978).
6. Lee Frakes-Eagan GQ-1249 Roadcnts along Ky. Hwy. Fort Payne Chert underlain
Pennington (Newell, 1975). 1595. by Floyds Knob
Newman glauconite.
Fort Payne
Grainger
7. Pennington Bledsoe GQ-889 Roadcuts and quarry along See quadrangle map Hauser and others (1957.
Newman (Csejtey, 1971). U.S. Hwy. 421, north side columnar section. p. 13—16); Wilpolt and
Grainger of Pine Mountain. Marden (1959, p. 622-625).
Chattanooga
8. Lee Louellen GQ—1060 Hurricane Gap section. See quadrangle map Wilpolt and Marden (1959,
Pennington (Froelich, 1978). Roadcuts and quarries columnar section. p. 631-636).
Newman along Ky. Hwy.160.

Grainger

DEVONIAN AND MISSISSIPPIAN ROCKS

TABLE 2.—Selected exposures of Mississippian rocks in Kentucky—Continued
[Numbers of sections keyed to map, ﬁg. 11]

27

 

Stratigraphic
units exposed

US. Geological Survey
Geologic Quadrangle Maps

Location and access

Features

 

Pertinent references

 

 

 

E...‘ ‘ ‘and .. . K 'y—r‘ ,
9. Lee Whitesburg-Flat Gap Roadcuts and quarry along See quadrangle map
Pennington GQ—1119 (Rice and US. Hwy. 119, north side columnar section.
Newman Wolcott, 1973). of Pine Mountain.
Grainger
10. Pennington Jenkins West GQ-1126 Pound Gap section. Wilpolt and Marden (1959.
Newman (Rice, 1973). Roadcuts and quarry p. 645—648).
Grainger along U.S. Hwy. 23.
11- Newman Hellier-Clintwood GQ-950 Secondary road and quarry See quadrangle map Smith and others (1967.
Grainger (Alvord. 1971). along Mountain Branch; columnar section. p. 11—13).
Sunbury south side of Ky. Hwy.
Berea 197.
Ohio
12. Lee Elkhom City—Harman Secondary road along Blue See quadrangle map Smith and others (1967,
Pennington GQ—951 (Alvord and Head Branch and quarry columnar section. p. 14—16).
Newman Miller, 1972). near head of Rough
Branch; south side of Ky.
Hwy. 197.
Southern central and south-central Kentucky
13- Salem Halls Gap GQ—1009 (Weir, Halls Gap Section U.S. Typical Borden of this area Weir (1970, p. 33-35,
Borden 1972). Hwy. 27 at Halls Gap. overlain by Salem. 43-44).
Muldraugh
Floyds Knob
Nancy
Halls Gap
14. Paragon Wildie GQ-684 (Gualtieri. Renfro Valley interchange. Renfro type section and up Weir (1970, p. 38—40,
Slade 1968a); Mount Vernon Interstate Hwy. 75 and per Borden. Glauconite in 45—46); Ettensohn and
Borden GQ-902 (Schlanger and US. Hwy. 25. Roadcuts Wildie. Representative Dever (1979, p. 170-181).
Renfro Weir, 1971). along 4 miles of US. Newman.
Wildie Hwy. 25 north from inter-
(Floyds Knob) change and roadcuts along
Halls Gap 1.5 mi of Interstate Hwy.
Nancy 75 south from interchange.
15. Fort Payne Mannsville GQ-562 (Taylor. Ky. Hwy. 76, north of bridge Representative Fort Payne; Sedimentation Seminar
1966). across Baker Branch. Floyds Knob exposure. (1972).
16. Science Hill Maretburg GQ-338 Roadcuts in Science Hill Lewis and Taylor (1979).
Sandstone Member (Schlanger. 1965). along Ky. Highway 70,
Rockcastle County. Prob-
ably same as section 33
of Lewis and Taylor (1979).
17. Fort Payne Cane Valley GQ-369 Fisher Bend Bluff north of Knifley type section. Sedimentation Seminar
Knifley (Maxwell and Turner, Green River Reservoir. (1972) (Section 1).
1964). Difﬁcult access.
18. Fort Payne Cane Valley GQ-369 Quarry along Butler Cane Valley type section. Kepferle and Lewis (1974).
Cane Valley (Maxwell and Taylor, Branch on Ky. Hwy. 55.
1964).
19. Monteagle Delmar GQ—909 (Lewis, Cumberland Parkway east Bedded-appearing chert in Dever and Moody (1979b).
St. Louis 1971a). of Fishing Creek (Lake Muldraugh. Large-scale

Salem and Warsaw
Borden
Muldraugh
Floyds Knob
Nancy

Cumberland); about 8 mi
west of West Somerset.
Quarries in Monteagle
Limestone on Smith and
Hale Knobs south of Ky.
Hwy. 80.

foreset beds in Borden.
Stromatolitic beds in St.
Louis. Relict-evaporite
features in lower St. Louis.
Salem and Warsaw unit
calcarenite and dolomite.

28 MISSISSIPPIAN ROCKS IN KENTUCKY

TABLE 2.—Selected exposures of Mississippian rocks in Kentucky—Continued
[Numbers of sections keyed to map, fig. 11]

 

 

 

ﬁgﬁggﬁ Goaligigegifgmzps Location and access Features Pertinent references
Southern central and south-central Kentucky—Continued
20. Fort Payne Creelsboro GQ-204 (Thaden West side of Wolf Creek Representative Fort Payne. Lewis and Potter (1978).
and Lewis, 1973); Wolf Dam and roadcuts along
Creek Dam GQ-177 Ky. Hwy. 1730.
(Lewis and Thaden, 1962).
21. Salem and Warsaw unit Jamestown GQ-182 (Thaden Shores of Lake Cumberland Carbonate mud mounds Lewis and Potter (1978).

 

 

Fort Payne and Lewis, 1962); J abez near Lake Cumberland and crinoidal limestone
Cane Valley GQ-483 (Thaden and State Park. bodies in Fort Payne.
Lewis, 1966). J abez Sandstone Member
of Fort Payne.
22. Monteagle Frazer GQ-1223 (Lewis, Ky. Hwy. 790, south of Kidder Limestone Member Lewis (1971b).

Kidder 1975). Kidder. type section; Hartselle
Formation composed of
shale.

23. Breathitt Burnside GQ—1253 (Taylor Roadcuts along 2 mi of Ettensohn and Chesnut
Paragon and others, 1975). US. Hwy. 27 north of (1979a, 1979b).
Bangor Sloans Valley. Strunk
Hartselle Construction Co. quarry,

Monteagle east side of US. Hwy.
Kidder 27, south of Tatesville.
24. Hartselle Cumberland City GQ-475 Roadcuts east of Cartwright Ste. Genevieve completely
Monteagle (Lewis and Thaden. 1965). along Ky. Hwy. 90 and exposed.
Kidder along secondary road
Ste. Genevieve leading southward to
St. Louis Poplar Mountain.
25A. Monteagle Albany GQ-550 (Lewis and Wago quarry. east of Ky. Lewis and Potter (1978,
Kidder Thaden, 1966). Hwy. 639, north of Wago p. 39—41).
Ste. Genevieve 3/4 mi.
St. Louis
253. Monteagle Albany GQ—550 (Lewis and Huddleston quarries, east Complete, or nearly com— McFarlan and Walker
Kidder Thaden. 1966). side of US. Hwy. 127, plete Kidder and Ste. (1956, p. 13).
Ste. Genevieve 2.5 mi north of Albany. Genevieve sections
near western limit of
Appalachian basin.
250. Hartselle Albany GQ-550 (Lewis and Gaddie Shamrock, Inc.
Monteagle Thaden, 1966). quarry, north side of Ky.
Kidder Hwy. 1590, west of
Ste. Genevieve Albany 1.5 mi.
Western, central, and west-central Kentucky

26. Elwren New Amsterdam—Mauckport Battletown quarry east of Best exposures of Ste. Stokley and McFarlan
Reelsville GQ—990 (Amos, 1972). Ky. Hwy. 228, north of Genevieve in region. Aux (1952, p. 64—68).
Sample-Mooretown Battletown V2 mi. Vases(?) equivalent.

Paoli
Ste. Genevieve

27. St. Louis Rock Haven-Laconia US. Hwy. 60 at Grahamton Rarely exposed St. Louis

Salem GQ—780 (Withington and section.

Sable, 1969).

Upper part of Salem and lower part of St. Louis are well exposed in at least two other localities:
Vulcan Materials Co. Brandenburg quarry, east side of Ky. Hwy. 933. Rock Haven—Lacenia GQ-780 (Withington and Sable, 1969)
Roadcuts along Ky. Hwy. 1638 west of Doe Run Mill. Rock Haven—Laconia and Guston GQ’s. Guston GQ-1481 (Palmer, 1978)

28. St. Louis Fort Knox GQ-1375 Round Hollow section, U.S. Shaly Salem. Reference Sable, Kepferle, and
Salem (Kepferle and Sable, Hwy. 31W north of section of Harrodsburg. Peterson (1966).
Harrodsburg 1977). Muldraugh. Muldraugh with crinoidal
Borden limestone beds.

DEVONIAN AND MISSISSIPPIAN ROCKS

TABLE 2.—Selected exposures of Mississippian rocks in Kentucky—Continued

[Numbers of sections keyed to map. ﬁg. 11]

29

 

 

 

fﬁéﬁgggsh; Gegligigesrﬁzlngslmgps Location and access Features Pertinent references
Western, central, and west-central Kentucky—Continued
29A. Harrodsburg Elizabethtown GQ-559 Bluegrass Parkway about Large roadcuts.
Borden (Kepferle. 1966b). 4 to 6 mi east of
Muldraugh Elizabethtown.
Floyds Knob
Nancy

293. Harrodsburg Elizabethtown GQ-559 Vulcan Materials Co. Striking foreset bedding in Kepferle, Peterson, and

Muldraugh (Kepferle, 1966b). Elizabethtown quarry Muldraugh. Sable (1964, p. 31—32).
north of U.S. Hwy. 62,
23/4 mi northeast of
Elizabethtown.

290. Salem Elizabethtown GQ—559 Tunnel Hill section; Classic localities. Butts (1917, 1922);
Harrodsburg (Kepferle, 1966b); Coles- Louisville and Nashville Stockdale (1939).
Muldraugh burg GQ—602 (Kepferle. Railroad cuts northeast

1967). of Tunnel Hill.
30. Chester units Summit GQ—298 (Moore, U.S. Hwy. 62. 1/2—1 mi east Fair exposures. USGS core

Paoli to Big Clifty

1964a).

of Summit.

hole here (Moore, 1964b).

Big Clifty asphaltic sand-

stone in quarries of
Summit.

 

Western Kentucky

 

31.

32.

33.

34.

35.

36.

37.

Big Clifty to Glen Dean

Leitchfield

Caseyville
Leitchfield

Renault
Ste. Genevieve

Tar Springs to Cypress

Salem
Warsaw
Fort Payne

Fort Payne

Big Clifty GQ—192
(Swadley, 1962).

Leitchfield GQ-l316
(Gildersleeve, 1978).

Caneyville GQ-1472
(Gildersleeve and
Johnson. 1978).

Burns GQ-1150 (Amos,
1974).

Olney GQ-792 (Trace and
Kehn, 1968).

Grand Rivers GQ-328
(Hays, 1964); Calvert City
GQ-731 (Amos and Finch,
1968).

Grand Rivers GQ-328
(Hays, 1964); Birmingham
Point GQ-471 (Fox and
Olive, 1966); Fenton
GQ-311 (Schnabel and
MacKallor, 1964);
Fairdealing GQ-320
(Wolfe. 1964).

Scattered cuts along U.S.
Hwy. 62 and Western
Kentucky Parkway.

Western Kentucky Parkway;
first cuts west of Leitch-
field interchange.

Western Kentucky Parkway.

Barrett quarry U.S. Hwy.
60; 5 mi north of Smith-
land, then 11/: mi south-
east on secondary road.

Western Kentucky Parkway
3V2—4l/z mi east of
Princeton.

Reed Crushed Stone Co.
quarry U.S. Highway 62
at Lake City.

Best exposures are along
Kentucky Lake shoreline.
Residuum and poor ex-
posures along and near
U.S. Hwy. 68.

Geologic map does not
show Western Kentucky
Parkway. Fossiliferous
Glen Dean in railroad
cuts at West Clifty.

Seldom exposed poorly
resistant section; partly
overgrown and slumped.

Excellent Pennsylvanian-
Mississippian contact
exposures.

Generally only residuum
exposed; chert content
2060-40 percent.

An excellent exposure of Fort Payne (interbedded chart and limestone) is present on the east shore of
Lake Barkley in a roadcut along Ky. Hwy. 295, west of Old Kuttawa.

Eddyville GQ-255 (Rogers, 1963).

Vincent (1975, p. 50—51. 64).

Williamson and McGrain
(1979, p. 34—35).

Williamson and McGrain
(1979, p. 33-34); Whaley
and others (1979,
p.42-44)

Dever and McGrain (1969,
p.100-109L

Trace (1981).

Dever and McGrain (1969,
p.48—57)

30

MISSISSIPPIAN ROCKS IN KENTUCKY

TABLE 2,—Selected exposures of Mississippian rocks in Kentucky—Continued

[Numbers of sections keyed to map, fig. 11]

 

 

 

Stratigraphic US. Geological Survey . .
units exposed Geologic Quadrangle Maps Location and access Features Pertinent references
Western Kentucky—Continued
38. Renault Princeton East GQ—1032 Kentucky Stone Co. Fredonia, Rosiclare, Levias, Dever and McGrain (1969.

39.

Ste. Genevieve

Chester series

(Trace, 1972).

Princeton East GQ—1032
(Trace, 1972).

40. Salem Canton GQ—279 (Fox and
Warsaw Seeland, 1964).

41. Bethel Hopkinsville GQ—651
Renault (Klemic, 1967).

Ste. Genevieve

42. 'I‘radewater Morgantown GQ-1040
Caseville (Gildersleeve, 1972);
Leitchfield Sugar Grove GQ-225
Glen Dean (Miller, 1963); Hadley
Hardinsburg GQ-237 (Rainey, 1963);
Golconda Rockfield GQ—309
Girkin (Rainey, 1964).

Princeton quarry via Ky.
Hwy. 91, 2% mi southeast
of Princeton, Ky.

Walches Cut. Illinois Cen-
tral Railroad, via Ky.
Hwy. 91 and secondary
roads, about 4 mi east-
southeast of Princeton,
Ky.

Kentucky Stone Co. Canton
quarry US. Hwy. 68, 1
mile east of Canton.

Christian Quarries quarry,
east side of Hopkinsville,
south of US. Hwy. 68.

Green River Parkway, from
interchange with US.
Hwy. 231 at Bowling
Green, northward to
interchange with US.
Hwy. 231 at Morgantown.

and Renault distinguished.

Classic locality. Dipping
beds, in part faulted.
Renault to Kinkaid.

Core drill site (Warsaw to
Chattanooga). See quad-
rangle map columnar
section.

Good exposures of
Chesterian units,
particularly Beech Creek
Member of the Golconda
Formation.

p. 118—133).

Dever and McGrain (1969,
p. 27—41 and fig. 6).

Dever and McGrain (1969,
p. 168—175).

Ste. Genevieve

 

(1) of known Devonian age (Ohio Shale and in part,
Chattanooga Shale), (2) known to span the Devonian-
Mississippian systemic boundary (New Albany Shale
and in part Chattanooga Shale), and (3) of equivocal age
relationships in which investigators are not in total
agreement about the position of the boundary (Bedford
Shale and Berea Sandstone).

NEW ALBANY, CHATTANOOGA, AND
OHIO SHALES

Rocks of Middle and Late Devonian age which under-
lie Early Mississippian strata in Kentucky are mainly
“black shale” units consisting of organically rich, ﬁssile
shale and thin-bedded siltstone. Three nomenclatural
units, the New Albany, Chattanooga, and Ohio Shales,
are laterally continuous with one another; their names
have been extended into Kentucky from Indiana, Ten-
nessee, and Ohio respectively. Nomenclature] bound-
aries of these units cropping out in Kentucky arbitrarily
coincide with mapped 71/2-minute quadrangle bound-
aries. In general, the name New Albany Shale is used
across central and into east-central Kentucky; the

 

Chattanooga is present in western, southwestern, south-
ern west-central, south-central, and southeastern Ken-
tucky; and the Ohio Shale is restricted to the exposure
belts in northeastern and east-central areas.

Maximum thicknesses of the Ohio, Chattanooga, and
New Albany Shales, the lower part of the Knobs Mega-
group (Swann and Willman, 1961), are more than 440
ft in western Kentucky (Schwalb and Potter, 1978), and
more than 1,700 ft in extreme eastern Kentucky
(Fulton, 1979; fig. 12 of this report). The units thin to
less than 40 ft thick along the Cincirmati arch, which
was a relatively positive area during the Middle and
Late Devonian; they are known to be absent in at least
two locations in south-central Kentucky, where Missis-
sippian rocks overlie Ordovician and Silurian strata
(Potter, 1978). The two correspondingly negative areas
of thick accumulations on either side of the arch are the
Moorman syncline in western Kentucky, and the west-
ern ﬂank of the Appalachian basin depositional trough
in eastern Kentucky. Thickness variations of the Devo-
nian shale units along the Rough Creek fault system
north of the Moorman syncline indicate pre-
Mississippian tectonic activity there (Schwalb and
Potter, 1978).

DEVONIAN AND MISSISSIPPIAN ROCKS

 

31

 

88° 85°

 

 

  

FOREST CITY '
BASIN

  

40° 1.1

KANSAS
I mesons:
l
35°—
g
.91) OZARK UPLIFT
~~~~a
OKLA. (MOE
360 M»: M“ N
EXPLANATION

—700— lsopach of thickness of Upper
Devonian rocks. in feet—Dashed
where approximately located or

restored
—0_ Zero limit of Upper Devonian
rocks
34° ‘ MEESI‘PESHWE
U 100 200 mm
W;

 

l l l

 

[1' I: ,’ | MICHIGAN
““1" “‘ “ BASIN

    
    
   
  

 

 

 

FIGURE 12.—Thickness of Upper Devonian rocks and their relationship to major structural elements in the east-central United States (from
Sable, 1979; Schwalb and Potter, 1978; Potter, 1978; Fulton, 1979, de Witt and McGrew, 1979).

BEDFORD SHALE AND BEREA SANDSTONE

The Bedford Shale and Berea Sandstone (N ewberry,
1870) overlie the Ohio and Chattanooga Shales in
eastern Kentucky with apparent conformity. Together,
they form a wedge which thins southwestward in east-
central Kentucky (McDowell and Weir, 1977). The Berea
is mappable from the Ohio border as far south as
Stricklett and Head of Grassy quadrangles (Morris,
1965b; 1966a); the Bedford extends farther southward
and during mapping was recognized as far as the vicini-
ty of the Means quadrangle (Weir, 1976). In the sub-
surface of eastern Kentucky, the Bedford has been
traced as far south as Knox County in southeastern
Kentucky (Kepferle and others, 1978) and pinches out
to the southwest.

 

The Bedford Shale is principally gray to greenish-gray
claystone with pyritic nodules and thin siltstone layers.
The overlying and intertonguing Berea Sandstone is
light-gray to buff, very fine to fine grained subgray-
wacke sandstone. Oscillation ripple marks trending
generally N. 50° W. to due west (Morris, 1965a, 1966b;
Morris and Pierce, 1967), crossbedding, and penecon-
temporaneous deformation features are common.
Grains are predominantly subangular quartz with minor
chert, feldspar, and rock fragments. Combined max-
imum thickness of Bedford and Berea strata is less than
200 ft in eastern and southeastern Kentucky (fig. 52),
which indicates two minor depocenters corresponding
to the Red Bedford and Virginia-Carolina deltas of
Pepper and others (1954). Thickness trends indicate a
general west to southwest paleoslope.

32

A classic regional analysis of these units in Ohio and
West Virginia was given by Pepper and others (1954),
but extended only into northeasternmost Kentucky.
The Bedford and Berea are discrete units in some areas
of northeastern Kentucky, but elsewhere they inter-
tongue and the Berea is in part a submarine channel
fill in the Bedford (Morris, 1966a; Morris and Pierce,
1967; fig. 13 of this report). They thin and apparently
grade into dark shales of the uppermost Chattanooga
and New Albany Shales (Elam, 1981). Overlying thin
Sunbury equivalents extend southwestward as the up-
permost beds of the Chattanooga into south-central
Kentucky and northern Tennessee (Elam, 1981).

In northeastern Kentucky, lithofacies and thickness
distribution of sandstone-dominated lobes (de Witt and
McGrew, 1979; fig. 52 of this report) suggest westerly
and southwesterly transport directions for the Bedford
and Berea sediments from source areas east and north-
east of the present Appalachian Mountains (Pepper and
others, 1954). The sandstones were derived probably
from earlier Paleozoic detrital rocks (Pepper and others,
1954, p. 91, 95). The ripple marks of N. 50° W. to due
west trend in the Berea of this area and throughout a
large adjacent area of southern Ohio have been ascribed
to northeasterly prevailing winds (Bucher, 1919; Pepper
and others, 1954, p. 91) or shoreline coastal control
(Hyde, 1911). Crossbedding and sandstone-filled chan-
nels characterize the sequence, and indicate deposition
in a shallow-water deltaic complex. Sole markings in
some sandstones suggest that they were deposited by
density currents, probably at the delta front (Wilson,
1950; Rich and Wilson, 1950).

 

MISSISSIPPIAN ROCKS IN KENTUCKY

In southeastern Kentucky, the Bedford and Berea are
mappable units southward to the Roxana (Maughan,
1976) and Benham (Froelich and Stone, 1973) quad-
rangles. In the eastern Kentucky subsurface, they thin
to a combined thickness of less than 10 ft along a north-
trending line from Leslie to Wolfe Counties.

DEVONIAN-MISSISSIPPIAN
SYSTEMIC BOUNDARY

In Kentucky, southern Indiana, and Tennessee, the
Devonian-Mississippian boundary can be determined
fairly closely although its position in northeastern and
southeastern Kentucky has been strongly debated. The
uppermost part of the New Albany Shale in its type area
of southern Indiana spans the boundary (Huddle, 1934;
Campbell, 1946; Lineback, 1968a). Beds of Kinder-
hookian age are represented in the top approximately
3 ft of the New Albany where the New Albany totals
about 100 ft thick. The base of the lowest Kinder-
hookian bed coincides closely with a subtle lithologic
break. In the upper 3 ft, which comprise the Underwood,
Henryville, and Jacobs Chapel beds (Campbell, 1946)
of the uppermost New Albany, conodonts are of Kinder-
hookian age. A similar section occurs in Kentucky in
the Louisville West quadrangle (Kepferle, 1974), but
in the adjacent Brooks quadrangle to the southeast
(Kepferle, 1972b), conodonts in the upper 4 ft of the
New Albany are of probable Late Devonian age. All
conodont collections from the uppermost beds \of the
New Albany of central Kentucky farther south either

 

 

 

 

 

 

 

. 0‘ . 6‘
,3“in $3 \

‘6

38 °30'

FIGURE 13.—Generalized relationships between Berea Sandstone (M Db) and Bedford Shale (M Dbd) in Vanceburg quadrangle, northeastern
Kentucky (Morris and Pierce, 1967). Berea Sandstone thickens eastward, fills channel cut in Ohio Shale (Do) as shown in panel X -X '.

Vertical exaggeration about X60.

ROCKS OF KINDERHOOKIAN AND EARLY OSAGEAN AGE 33

indicate a probable Late Devonian age or are non-
diagnostic (J .W. Huddle, written commun., 1967).

The position of the Devonian-Mississippian boundary
in northeastern Kentucky is extrapolated from ex-
posures in Ohio, and was interpreted by de Witt (1970),
using conodont and ﬂoral evidence, to lie within the
basal few feet of the Bedford-Berea detrital wedge that
separates the Sunbury Shale of Mississippian (Kinder-
hookian) age from the underlying Late Devonian Ohio
Shale. More recent spore evidence (Eames, 1978) indi-
cates that the boundary lies between the Berea-Bedford
and the Sunbury Shale in northeastern Ohio. The
systemic boundary in northeastern Kentucky is also
considered by Sandberg (1981) to lie at the base of the
Sunbury Shale.

Conodonts from beds in the Sunbury Shale, the up-
permost part of the New Albany Shale, and basal beds
of the Fort Payne Formation in east-central, central,
and south-central Kentucky are shown in tables 3—6 of
this report. Their distribution indicates that the upper-
most beds of the New Albany in east-central Kentucky
at least as far south as the Berea and Bighill quad-
rangles are of probable Sunbury (Kinderhookian) age.
Southward in Kentucky, conodonts from the uppermost
beds of the New Albany are nondiagnostic in places; and
beds of the basal Borden and Fort Payne Formations
(Maury Formation equivalent) contain conodonts of
Kinderhookian age in some localities and Late Devonian
age in others (Hass, 1956). Very slow deposition and
gaps in conodont assemblages seem to have character-
ized the Devonian-Mississippian boundary conditions
in south-central and parts of central Kentucky. In some
areas, a hiatus above the uppermost beds of the New
Albany Shale is indicated by conodonts of probable up-
per Burlington (Osagean) age in the basal beds of the
Borden Formation (Rexroad and Scott, 1964).

The evidence suggests that deposition was generally
continuous across the Devonian-Mississippian bound-
ary in the northern part of northeastern and central
Kentucky, but that during this time nondeposition or
very slow deposition, periodically interrupted by sub-
marine scour, characterized south-central Kentucky
(Conkin and Conkin, 1979).

MISSISSIPPIAN ROCKS

ROCKS OF KINDERHOOKIAN AND EARLY
OSAGEAN AGE (EARLY MISSISSIPPIAN)

Earliest Mississippian (Kinderhookian) rocks in Ken-
tucky are widespread, but they are relatively thin and
constitute a small part of the total Mississippian rock
volume. In eastern Kentucky, the strata which succeed

 

the Bedford-Berea deltaic wedge are terrigenous organic
clastics. During Kinderhookian time a thin distinctive
green claystone unit was deposited across southern
Kentucky, and a thin unit of carbonate rocks with in-
terbedded mudstones was deposited in west-central and
western Kentucky. These units are, respectively, the
Sunbury Shale, the equivalent of the Maury Formation
of Tennessee, and the Rockford Limestone (fig. 7).

SUNBURY SHALE

The Sunbury Shale (Hicks, 1878) provides an easily
recognizable horizon and is extensively used as a struc-
tural datum; it consists of fissile organic-rich claystone
and siltstone similar to that in the older Ohio and Chat-
tanooga Shales. It is 10—25 ft thick in northeastern Ken-
tucky (fig. 14) and as much as 55 ft thick along Pine
Mountain, southeastern Kentucky. Both the upper and
lower contacts appear to be conformable. The unit is
readily distinguishable by its radioactivity profile (Et-
tensohn, 1979b; Ettensohn and others, 1979). Still far-
ther south and west, beds in the uppermost part of the
New Albany and Chattanooga Shales have been recog-
nized as Sunbury equivalents by their contained con-
odonts and by geophysical log studies (Elam, 1981). The
conodonts, principally Siphonodella spp., occur in the
top few inches of the New Albany Shale in the Berea
(Weir, 1967) and Bighill (Weir and others, 1971)
quadrangles.

Except for southwestward thinning, the Sunbury
Shale exhibits no definitive clues to the transport direc-
tion of its sediment. Following progradation of Bedford-
Berea sediments under shallow aerobic conditions, the
Sunbury accumulated in an anaerobic, probable deep-
water environment during a marine transgression, prob-
ably from the same general eastern provenance and
under conditions similar to those which produced the
Devonian-Mississippian black Shales. Lineback (1970)
considered that the black Shales of the eastern midcon-
tinent were deposited in shallow water with circulation
restricted by the presence of an algal ﬂoatant. Deep-
water environments for these black Shales were favored
by Rich (1951) and by Elam (1981).

MAURY FORMATION EQUIVALENT

South and west from about the vicinity of Berea, a
few inches to about a foot of greenish claystone with
phosphatic nodules, glauconite, and lag concentrates
containing conodonts, unidentified fossil fragments, and
fish remains directly overlies the New Albany or Chat-
tanooga Shale (Elam, 1981; fig. 15 of this report). This
lithic and positional equivalent of the Maury Formation
of Tennessee (Safford and Killebrew, 1900) is also

34 MISSISSIPPIAN ROCKS IN KENTUCKY

 

FIGURE 14.—Sunbury Shale (above vehicle) overlying the Bedford Shale (mostly talus-covered slope) and underlying the Farmers Member
of the Borden Formation (tabular light-hued beds). Milepost 132.6, Interstate Highway 64, Farmers quadrangle, Rowan County.

reported in basal beds of the New Providence Shale
Member of the Borden and Fort Payne Formations in
many quadrangles in south-central and central Ken-
tucky. These include Eli, Faubush, Delmer (Thaden and
Lewis, 1965a; 1969; Lewis, 1971a) and as far west as
Dubre quadrangle (Lewis, 1967) in south-central Ken-
tucky, and Louisville West, Shepherdsville, and
Howardstown (Kepferle, 1974; 1968; 1966a) in western
central Kentucky. Similar greenish claystone above
black shale is reported in many well descriptions in
southern and western Kentucky by Freeman (1951,
1953).

Conodonts from several outcrops in southern Ken-
tucky and Tennessee suggest that the Maury is a time-
equivalent of most of the Bedford, the Berea, and the
Sunbury (Collinson and others, 1962, p. 13), and that
it locally contains elements younger than the Sunbury
(Hass, 1956, p. 23). In these areas, the Maury may repre-
sent generally uninterrupted deposition during most or
all of Kinderhookian time (Conant and Swanson, 1961,
p. 67). In others, as along the Cincinnati arch in the
western part of central Kentucky, Maury lithic equiva-
lents contain a mixed conodont assemblage of Devo-
nian, Kinderhookian, and Osagean ages, and are
interpreted as a lag concentrate at an erosional hiatus,
which possibly extended through Kinderhookian and
early Osagean times (Rexroad and Scott, 1964). This in-
terpretation supports the concept of very slow deposi-
tion and essentially starved basin conditions over much

 

of Kentucky and Tennessee during earliest Mississip-
pian time.

The widespread, thin claystone of the Maury Forma-
tion and its lithic equivalents in the basal Borden and
Fort Payne probably were deposited remote from source
areas (Conant and Swanson, 1961, p. 68), at a slow rate,
in a low-energy environment favorable to the precipita-
tion of phosphate. The concentration of phosphate
nodules in the Maury equivalents appears to represent
a lag concentrate from which most of the fine detrital
material was winnowed by submarine currents; this en-
vironment is also suggested by mixed conodont “lag”
assemblages near the Cincinnati arch. South of Ken-
tucky, thin sandstones occur in the Maury along the
western margin of the Nashville Dome and suggest local
source areas (Conant and Swanson, 1961, p. 53). The
normally subjacent Chattanooga Shale is also locally
absent there, indicating local uplift or local scour and
resultant deposition along the arch. The abundant
phosphatic nodules in the Maury Formation equivalent
may indicate that southern Kentucky and Tennessee
occupied a broad shelf north of a deep basin from which
upwelling ocean currents contributed nutrients such as
phosphate. Phosphatic nodules have also been reported
in the underlying New Albany Shale and in the overly-
ing New Providence Shale Member of the Borden For-
mation, indicating that conditions for the precipitation
of phosphate existed periodically from Late Devonian
into Osagean time.

ROCKS OF KINDERHOOKIAN AND EARLY OSAGEAN AGE 35

 

FIGURE 15,—Maury Formation equivalent (Mme). 0.8 ft thick,
underlying Fort Payne Formation (pr) and overlying Chattanooga
Shale (Dc). East side of Kentucky Highway 61, 4.7 mi south of junc-
tion of Kentucky Highways 90 and 61, east side of Burkesville.
Frogue quadrangle, Cumberland County.

ROCKFORD LIMESTONE

In much of Indiana and adjacent parts of central and
west-central Kentucky, earliest Mississippian strata are
represented by the thin but widespread Rockford
Limestone and thin shale beds which are superjacent
to or in the uppermost few feet of the New Albany
Shale. The Rockford, first named the Rockford
Goniatite bed by Meek and Worthen (1861), was given
formational status by Kindle (1899). The Rockford, a
gray to greenish-gray, micritic, glauconitic limestone or
dolomite, was reported to contain both Kinderhookian
and Osagean conodont elements in Indiana (Rexroad
and Scott, 1964) but Kinderhookian foraminifera in In-
diana and northwestern central Kentucky (Conkin and
Conkin, 1979). In northwestern central Kentucky, the

 

Rockford is locally present south of the Ohio River in
the Louisville West quadrangle (Kepferle, 1974) and was
reported farther south in the Brooks quadrangle (Con-
kin and Conkin, 1979, p. 57). In this area, it is as much
as 3 ft thick and directly overlies New Albany Shale.
Southward, the Rockford grades to shale (Conkin and
Conkin, 1979, p. 57), and thin carbonate strata in the
position of the Rockford are also reported to the south-
west in well logs of northern Meade and Breckinridge
Counties. In the Louisville West quadrangle, the Rock-
ford apparently is correlative with lithologies of the
Maury equivalent.

In western Kentucky, some well records in counties
along the Ohio River indicate that the Rockford Lime-
stone (Chouteau Limestone of Illinois usage) is present
there. In the Canton quadrangle, Trigg County, a thin
limestone within the unit referred by Fox and Seeland
(1964) to the New Providence Shale may also represent
the Rockford. If this limestone is the Rockford, the 34 ft
of shale between it and the older Chattanooga Shale is
probably equivalent to the Kinderhookian Hannibal
Shale of Illinois.

Carbonate rocks of the Rockford Limestone in In-
diana and parts of northern Kentucky were mostly
deposited in a notably low energy environment, on a sea
floor having little relief. A high iron content and a low
proportion of terrigenous detritus in Indiana may in-
dicate a low land area to the east along the Cincinnati
arch; dolomitic beds may suggest an intratidal deposi-
tional environment, but the enclosing sediments in-
dicate a subtidal marine environment.

The Rockford Limestone in northwestern central Ken-
tucky seems to be coeval with the Maury Formation
equivalent farther south. The Maury equivalent may en-
compass several small-scale disconformities ranging
from Late Devonian to Osagean age, expressed by
superposition of and mixing of lag concentrates in a
starved shelf or basin environment. Landward, such
small-scale disconformities might be expressed individ-
ually by separate detrital units containing lag con-
centrates, as suggested by Conkin and Conkin (1979)
to be present at the Devonian-Mississippian and
Kinderhookian-Osagean boundaries.

KINDERHOOKIAN-OSAGEAN UNIT RELATIONSHIPS

Fine detrital rocks, mostly of the New Providence and
Nancy Members (Osagean) of the Borden Formation,
overlie Sunbury and New Albany Shales in northeast-
ern, east-central, central, and west-central Kentucky.
Judging from the widespread persistence of the thin
Sunbury, the Maury equivalent, and the Rockford units,
little hiatus preceded deposition of Borden sediments
except locally along the western side of the Cincinnati

36 MISSISSIPPIAN ROCKS IN KENTUCKY

arch. There, an unconformity at the base of the Fort
Payne Formation in the Petroleum quadrangle, south-
ern west-central Kentucky, was reported by Myers
(1964). The Rockford Limestone is also locally absent
in the southern Indiana outcrop belt (Lineback, 1964),
and shallow channels filled by Borden strata cut into
the New Albany Shale in northwest-central Kentucky
(R.C. Kepferle, oral commun., 1967). Scour prior to or
during deposition of Borden Formation sediments is
therefore indicated in those areas.

In much of the region southwest of the limit of rocks
assigned to the Borden, in southern and western Ken-
tucky and western Tennessee, mudstones in the New
Providence Shale Member of the Fort Payne Formation
rest with apparent conformity on the Maury Formation
and its lithic equivalents in places, but abrupt contacts
also exist between overlying Fort Payne siliceous rocks
and the Maury. These were considered by Conant and
Swanson (1961, p. 68) to be due to sudden changes in
depositional environments. Conodont faunas collected
at scattered localities in southern Kentucky, however,
indicate that basal beds of the Fort Payne may vary
in age from place to place (J .W. Huddle, written com-
mun., 1967), and an obscure nondepositional or sub-
marine erosional hiatus may therefore be reﬂected
between Maury and Fort Payne strata. Such relations
would be expected for a distal, sediment-starved area
basinward of deltaic deposits which were subsequent-
ly overlain by transgressive carbonate strata of the Fort
Payne.

PALEOTECTONIC IMPLICATIONS

Because no known physical evidence points to inter-
systemic unconformities of much magnitude, structural
elements during Kinderhookian time in Kentucky were
probably of extremely low amplitude. The region was
mildly negative, except for the Cincinnati arch, which
may have been a barrier to westward dispersal of Late
Devonian and Kinderhookian clastic sediments.

If a shelf-trough depositional and tectonic model is
invoked, the widespread thin deposits of the Rockford
and Chouteau limestones in Kentucky could have ac-
cumulated on the southern part of a broad stable shelf.
Maury Formation equivalents accumulated either on
the southern extension of this shelf, or in a negative ele-
ment of larger magnitude which adjoined the shelf to
the south. The northward and westward change from
these phosphatebearing Maury equivalent rocks to car-
bonate rocks of the Rockford Limestone may suggest
northward transition from a subsiding oceanic basin,
the Ouachita trough. Alternatively, considering a
trough model without transition to a shelf, a eustatic

 

rise in sea level could have initiated or extended starved
basin conditions. The Rockford Limestone in such a
model might represent a deep-water carbonate deposit
(R.C. Kepferle, oral commun., 1985).

ROCKS OF MOSTLY OSAGEAN AGE

Rocks of Osagean age are comparatively thicker,
lithologically more diverse, and exhibit more complex
interrelationships than earlier Mississippian rocks (ﬁgs.
7 and 9). They comprise two marine rock assemblages
which are dominant in specific areas:

1. Older terrigenous detrital clastics (Borden Forma-
tion, New Providence Shale Member, Grainger Forma-
tion, and basal clastic parts of Fort Payne Formation)
deposited as prograding deltaic wedges. These are most-
ly confined to the eastern and central parts of Kentucky,
but their thinner distal equivalents extend throughout
much of the southern and western parts of the State.
They are generally bracketed in Kentucky by underly-
ing “black shale” units and Maury Formation lithic
equivalents and by overlying dominantly carbonate-
rock units.

2. Younger siliceous, argillaceous, dolomitic, and
crinoidal carbonate-rock units with local sandy ter-
rigenous clastics (most of the Fort Payne Formation,
Muldraugh Member of the Borden Formation and
Renfro Member of the Slade Formation), which extend
from east-central Kentucky across southern and west-
ern Kentucky and Tennessee into southeastern Illinois
and southwestern Indiana.

During the 1960—1978 geologic mapping program, the
concept of deltaic deposition for Borden and Fort Payne
terrigenous clastic rocks in Kentucky (Borden delta)
was developed (Peterson and Kepferle, 1970), a concept
that has stimulated related research by individuals and
groups; see for instance Sedimentation Seminar (1972);
Whitehead (1976); Benson (1976); Kepferle (1972a,
1977a); and Kepferle and Lewis (1974).

Thickness and directional transport components of
sandstone units of the Borden and Fort Payne are
shown in figure 16. Distribution of distinctive lithic
units, dark shales, crinoidal limestones, and greenish
mudstones, which make up significant rock volumes in
some areas, is shown in ﬁgure 54 and discussed on pages
51 and 100—101.

BORDEN FORMATION
In eastern and central parts of Kentucky, units

of Osagean age are dominated by terrigenously de-
rived detrital rocks in the lower part, overlain by

88° 88°

ROCKS 0F MOSTLY OSAGEAN AGE

87" 88° 85°

37

82°

 

l

*"‘i
L e Z _ 4
Area depicting selected
sandstone units
1 Farmers Siltstone

39" _ EXPL‘ANATION

l )

lsopach thickness in feet—
Thickness limits given where
not depicted by isopachs

<—

Member (Moore and

Clarke. 1970) and “Weir”

sandstone of subsurface

terminology

38° _ 2 “Rockcastle freestone”
(Weir, 1970)

3 Jabez and Kniﬂey Sand-
stone Members (Kepferle,
Lewis, and Taylor, 1972; '
Kepferle and others,

1980)
4 Kenwood Siltstone :3,
.3. Member (Kepferle 1977) _‘

Sediment transport directions
as given in references
listed for areas 1—4

 

 

37° —

 

 

 

 

 

58 l88 MlLES

 

FIGURE 16.—Distribution (stipple pattern). thickness, and sedimentary transport directions of selected sandstone units in the Borden and
Fort Payne Formations in Kentucky. Base includes selected faults and fault zones.

carbonate-rich beds in the upper part. The Borden For-
mation (Weir and others, 1966), in part correlative with
the Cuyahoga and Logan Formations of Ohio (Hyde,
1915), the Borden Group of Indiana (Cumings, 1922),
and the Pocono Group of West Virginia, makes up all
or most of Osagean age-equivalents in northeastern
east-central, central, and parts of south-central Ken-
tucky. The Borden Formation consists of 10 members.
The terrigenous clastic part of the Borden in Kentucky
typically consists of lower units of gray and green
mudstone and siltstone (Nancy and New Providence
Members) containing planar-bedded sandstone
(Farmers and Kenwood Members) and lenses of
crinoidal limestone. Sandstones and siltstones are com-
posed largely of subangular to subround quartz and
feldspar grains with quartz and calcite cement; they are
mostly subgraywackes originating in a provenance of
largely felsic plutonic rocks (Kepferle, 1977a). Tongues
and lenses of siltstone in the middle part of the Borden
in Kentucky grade northward and eastward into thicker
and more abundant units of coarser siltstone and sand-
stone in Ohio and West Virginia. In Kentucky these
tongues and lenses include the Cowbell, Halls Gap, and
Holtsclaw Members, and thinner, less extensive units
such as the Gum Sulphur Bed of the Nancy Member
and “Rockcastle freestone” of the Wildie Member. The
detrital units form a elastic deltaic wé’dge which thins
southwestward, and in large exposures they exhibit

 

large-scale foreset bedding with westerly dip com-
ponents (fig. 24). The uppermost units, the Wildie and
Nada Members, are composed mostly of shale, including
variegated varieties. Contacts of elastic units within the
Borden Formation are generally gradational and do not
show evidence of significant depositional or erosional
breaks, except for thin glauconitic beds that are inter-
preted to mark a depositional hiatus or an interval of
slow sedimentation.

As mapped, the uppermost, dominantly carbonate
rock divisions of the Borden included the Renfro
Member and the Muldraugh Member. The Renfro
Member has been reassigned as the basal member of
the Slade Formation by Ettensohn and others (1984).
It is, however, discussed here with Borden Formation
rocks with which it is in part equivalent. The Renfro
includes dolomitic limestone and dolomite in northeast-
ern, east-central, eastern, and south-central Kentucky.
The Muldraugh Member, a thicker unit of light-colored
cherty and silty dolomite and limestone, crops out in
central and south-central Kentucky. The carbonate rock
units are commonly separated from underlying elastic
strata by thin glauconitic siltstone and limestone beds
such as the Floyds Knob Bed of the Muldraugh Member
in central Kentucky, and by thin glauconitic siltstone
beds in northeastern, east-central, south-central, and
west-central Kentucky. The Muldraugh is overlain by
the Harrodsburg Limestone or the Salem and Warsaw

38

MISSISSIPPIAN ROCKS IN KENTUCKY

 

SW NE
FEET
400
Harrodsburg Limestone (crinoldal biorudite) 14f
///ew $114-1

 

   
 

   
   

_—_—~_—:_1—_—_‘*:_—____—ﬁ—_—;::r‘_—_ Shale Member —‘_‘—_ —_—_—__—f——_:_—_‘__ _ A __

   

  
      
  
   

- .2 .. r. all-4 _..._...._._L
...,’,”_‘_;7_‘ Muldraugh Member 43:2,?” 495‘4:_:J_—O_.
--(do|o§iltite and calcgiltite))-,Hﬁ/,j. _ ”AL—L", 3%

  
 
  

L". ﬁNancy Membe '
..’__;(gray'to greeﬂjhale a—n_d silt

. ,’

 

 

New Albany Shale (fissile organic shale)

 

 

lllLL)

l 2 MlLES
l |

FIGURE 17.-Restored cross section of the Borden Formation delta “front” in Howardstown quadrangle (Kepferle, 1966a), north-central
Kentucky, showing relationships of Borden units. Datum, top of New Albany Shale. Modiﬁed from Kepferle, 1966a.

Formations unit. Westward thickening of the carbonate
rock units and reciprocal thinning of the underlying
detrital units are characteristic internal relationships
in the Borden. Contacts between these rocks are along
generally southwestward-sloping planar surfaces (Peter-
son and Kepferle, 1970; figs. 9 and 17 of this report).

The southwest-thinning terrigenous elastic wedge of
the Borden is an upward-coarsening succession of beds,
representing westward progradation of the Borden delta.
As indicated in the cross section (fig. 9), two pulses of
clastic deposition culminated in deposition of the
Farmers Member and the Cowbell Member; shales of the
Nancy and New Providence Members represent distal
prodelta sediments of the actively prograding deltaic
system, followed by cessation of delta encroachment
during which time the glauconite-bearing beds in the
Nada Member and Floyds Knob Bed were deposited.

FARMERS MEMBER

In northeastern Kentucky, the Farmers Member is
the basal unit of the Borden Formation (fig. 14). The
Farmers includes the Vanceburg Sandstone Member of

Hyde (1915), and was termed the Farmers Siltstone
Member of the New Providence Formation by Stockdale
(1939). The Farmers in its type area is 60 percent planar-
bedded subgraywacke sandstone and 40 percent shale.
It is more than 200 ft thick in northeastern Kentucky,
thins southward into the Olympia quadrangle
(McDowell and Weir, 1977) and extends into the Clay
City quadrangle (Simmons, 1967; fig. 18 of this report).
It is mapped with the Nancy Member of the Borden
Formation in the Preston, Means, and Levee quad-
rangles (Weir and McDowell, 1976; Weir, 1976;
McDowell, 1978), where it consists of the Clay City
(sandstone) and underlying Henley (shale) beds.
Directional features, principally sole markings, indi-
cate westward paleocurrent transport directions of the
Farmers (Rich and Wilson, 1950; Rich, 1951; Wilson,
1950; fig. 16 of this report). The member is interpreted
by Moore and Clarke (1970) to be a turbidite deposit.
The “Weir Sand,” a driller’s term for several subsurface
sandstones in eastern Kentucky and adjoining States,
is in about the same stratigraphic position as the
Farmers (Pepper and others, 1954). Of the two apparent
south-trending lobes of sandstone in northeastern

 

SW

ROCKS OF MOSTLY OSAGEAN AGE 39

Olympia quadrangle NE

 

SYSTEM

SERIES

FORMATION, MEMBER, AND BED

SERIES
SYSTEM

 

Nancy Member of
Borden Formation

 

MISSISSIPPIAN
Lower

Clay City Bed

   

l

 
    

Farmers Member
of Borden
Formation

Lower
MISSISSIPPIAN

 
   
   

    
 

 

Henley Bed l

 

 

 

 

 

/ x ___
__7___7_ datum Sunbury Shale l Bedford Shale Berea Sandstone 7 7
l
a l
3 : .
Z n- a)
< 3 | o.
—- NewAlbany . a. z
E g Shale : OhloShale 3 S
LI>J 33 | g g
a E l w >
.— _ Lu
2 | E o
l s
l

 

 

 

 

FIGURE 18.—Generalized stratigraphic diagram of lower part of the Borden Formation and underlying units in northeastern Kentucky
(from McDowell and others, 1981).

Kentucky shown in figure 16, the western lobe
represents the Farmers, strictly speaking, and the
eastern lobe, more than 100 ft thick, is a “Weir Sand,”
which may be a subsurface continuation of the Farmers.

NANCY AND NEW PROVIDENCE MEMBERS

The widespread Nancy Member, from about 150 to
300 ft thick, represents distal foreset and bottomset
strata of the Borden delta. For the most part poorly
resistant greenish-weathering gray silty shale and
siltstone, the Nancy gradationally intertongues with the
distal portions of several resistant siltstone-dominant
members and beds (Cowbell, Holtsclaw, Roundstone,
Conway Cut) (fig. 19). The Nancy Member also contains
the Gum Sulphur Bed, a lentil of resistant siltstone in
the Nancy shales (fig. 20) that emphasizes the elastic
wedge character of the lower Borden. In central Ken-
tucky, beds equivalent to the lower part of the Nancy,
mostly greenish-weathering clay shale with sideritic
claystone nodules and lenses, are termed the New Prov-
idence Shale Member of the Borden Formation
(Kepferle, 1971). The New Providence Member in the
Louisville West quadrangle (Kepferle, 1974), 120 to

250 ft thick, encompasses the Kenwood Siltstone
Member, a unit similar to the Farmers Member but
thinner, ﬁner grained, and less extensive. The Kenwood,

 

FIGURE 19.—Borden Formation; contact of Cowbell Member (resist-
ant unit) and underlying poorly resistant Nancy Member. US.
Highway 25 at Boone Gap, Berea quadrangle, Rockcastle County.

 

40 MISSISSIPPIAN ROCKS IN KENTUCKY

 

 

 

FEET
300

200

100

 

3 4'1 E13 MILES

 

 

 

 

Location of sections

 

FIGURE 20.—Sections showing relationships of siltstone units in the Borden Formation in the area between Brodhead and
Bighill, east-central Kentucky (from Weir and others, 1966).

an illite subarkose to illite-arkose, was studied intensive-
ly by Kepferle (1972a, 1977a), who concluded it to be
a pro-delta turbidite apron deposit. Directional elements
in the Kenwood indicate west-southwest sediment
transport directions (fig. 16).

The New Providence and Nancy locally contain
crinoidal limestone lenses and concentrations of
fossils at many localities, including the classic Button
Mould Knob fauna locality of Butts (1917, p. 11—17)
and the Coral Ridge fauna, both studied by Conkin
(1957), Kammer (1982), and Gordon and Mason (1985).
Relationships of these clastic units in the Borden
of central to south-central Kentucky are shown in
figure 21A, B.

In western central Kentucky the Nancy and New
Providence are overlain abruptly by the Muldraugh
Member (fig. 22) or they inter-tongue with siltstone beds
of the Holtsclaw Member. In westernmost Kentucky,
the New Providence Shale is given formational rank as
a unit between the Fort Payne Formation and the New
Albany Shale, as used in the Briensburg quadrangle
(Lambert and MacCary, 1964); it is equivalent to the
Springville Shale (Savage, 1920) of Illinois, and possibly
to the basal fine—grained green shale of the Fort Payne
above the Maury Formation equivalent in south-central
Kentucky.

 

COWBELL, HALLS GAP, AND HOLTSCLAW MEMBERS

Resistant siltstone members of the Borden, made up
of dominantly gray subgraywacke siltstone, generally
with indistinct bedding because of extensive bioturba-
tion, include the Cowbell Member in northeastern and
east—central Kentucky (fig. 23), the Halls Gap Member
in central, south-central and east-central Kentucky, and
the Holtsclaw Siltstone Member in western central Ken-
tucky. These mapped units, generally less than 100 ft
but as much as 250 ft thick, include large-scale deltaic
foreset bedding; and they intertongue with the Nancy
Member. The Halls Gap Member exhibits these foreset
beds in exposures such as the railroad cut at Kings
Mountain in the Halls Gap quadrangle (Weir, 1972; fig.
24 of this report). These beds dip in westerly and south-
westerly directions at angles of generally less than 5 °.
The Cowbell and Halls Gap Members thin or grade to
extinction westward. The Halls Gap extends as far west
as the Saloma and Raywick quadrangles (S.L. Moore,
1976; Kepferle, 1973).

NADA AND WILDIE MEMBERS

Two ﬁne-grained clastic units in the upper part of the
Borden in east-central and northeastern Kentucky are

ROCKS OF MOSTLY OSAGEAN AGE

 

 

 

 

 

 

 

 

 
 
 

    

 
 
 

 

 

 

 

 

 

 

 

 

 

 

B

 

 

 

 

 

 

41

 

KOSMOSDALE VALLEY STATION 338%?
QUADRANGLE QUADRANGLE RANGLE FEET
. 500
Harrodsburg Limestone
Muldraugh Member
— — 400
Holtsclaw Siltstone Member
c
— é — 300
g Nancy Member
8
— E, -- 200
E
o
m Kenwood Siltstone Member 100
New Providence Shale Member
datum\ New Albany Shale
' o
A o 1 2 3 4 5 MILES
l l | | J
SHEPHERDSVILLE
BROOKS / I CRAVENS 1 NEW HAVEN l fPUHLINGTOlIV IYOSEMITE
datum FEET
§ Muldraugh Member -300
/
\ / Halls Gap Member
\ 2___
\ — 200
Nancy Member
Holtsclaw
Siltstone /\
Member
New Providence Shale Member _100
Siltstone — 0
Member /
0 10 20 30 40 MILES
l’J l l l I l

FIGURE 21.—Relationships of the elastic units of the Borden Formation. A, As mapped in Kosmosdale, Valley Station, and
Brooks quadrangles. central Kentucky (from Kepferle, 1971). B, As mapped from southern Jefferson County to Casey County,
central and south-central Kentucky (from Kepferle, 1971). Dashed line, relationships uncertain.

42 MISSISSIPPIAN ROCKS IN KENTUCKY

 

FIGURE 22.——Nancy Member (at vehicle level) and overlying
Muldraugh Member of the Borden Formation. Milepost 97.8, Inter-
state Highway 65, 3.5 mi north of Elizabethtown exit, Colesburg
quadrangle, Hardin County.

the Nada and Wildie Members (fig. 9). The Nada (fig.
25), 30—65 ft thick, is predominantly variegated (olive
gray, grayish red, grayish purple), clayshale and silt
shale with siltstone and glauconitic siltstone; it extends
from northeastern Kentucky southeastward through
the Berea quadrangle (Weir and others, 1966, p. F15)
and the northern part of the J ohnetta quadrangle
(Gualtieri, 1968b). In northeastern Kentucky the Nada
locally contains beds of crinoidal limestone, as in the
Cranston quadrangle (Philley and others, 1974).
Glauconite—rich layers at the base and top or within the
member occur south of the Berea quadrangle (Weir,
1967), and a glauconite layer occurs a few feet below
the top of the Nada farther north. The lower part of the
Nada is interpreted to grade southwestward into the
Nancy Member and the siltstone of the Halls Gap
Member, and the upper part into greenish siltstone and
shale of the Wildie Member.

The Wildie Member contains very minor amounts of
variegated, mostly greenish shale, common beds of
glauconitic siltstone, and locally phosphatic nodules at
its top, base, or both (Brown and Osolnik, 1974;

 

Gualtieri, 1967a; Schlanger and Weir, 1971). In general,
reddish hues in these units diminish southward along
the eastern Kentucky outcrop belt.

The Nada and Wildie Members are both overlain by
the Renfro Member (fig. 25). The Wildie merges
westward with the Halls Gap Member (Weir, 1972) or
may thin to become the Floyds Knob Bed (Whitehead,
1976). In earlier mapping, the Nada was considered part
of the Muldraugh Member, as in the Haldeman
quadrangle (Patterson and Hosterman, 1961). An infor-
mal unit of limited extent in the Wildie Member is the
“Rockcastle freestone,” in which paleocurrent features
indicate westward directions of transport (fig. 16), and
bedding features suggest that it is a turbidite (Weir,
1970, p. 39).

RENFRO MEMBER

Carbonate rock units named the Renfro and
Muldraugh Members (Weir and others, 1966) were
placed in the uppermost part of the Borden. These units
were considered part of the Borden until 1984, when the
Renfro was placed in the basal part of the Slade For-
mation, overlying the Borden (Ettensohn and others,
1984).

The Renfro Member (Schlanger, 1965) of east-central,
northeastern, and south-central Kentucky, with its sub-
surface and southeastern Kentucky equivalents, is a
widespread unit throughout much of the eastern part
of the State. It consists dominantly of aphanitic to ﬁne-
ly crystalline, argillaceous dolomite and dolomitic lime-
stone with local interbeds of gray micritic limestone.
The dolomite in part represents dolomitized calcarenite
with few relict grains. At some localities, siliceous car-
bonate rocks, probably Muldraugh Member equiva-
lents, have been included in the basal Renfro, and St.
Louis Limestone micritic lithologies have been included
in the upper part of the Renfro. The unit weathers to
distinctive yellow and orange hues. Macrofossils are
very sparse. Grayish-green shale beds are relatively
common, as are arenaceous and glauconitic grains, par-
ticularly in east-central Kentucky. In most places the
Renfro conformably overlies glauconitic siltstone beds
of the Nada or Wildie Members, and in the Woodstock
quadrangle where the Wildie pinches out, the Renfro
rests on the Halls Gap Member of the Borden Forma-
tion (Weir and others, 1966). In northeastern Kentucky,
the Renfro is thin, and was included with the Newman
Limestone on some geologic quadrangle maps, for ex-
ample Brushart (Denny, 1964) and Olive Hill (Englund
and Windolph, 1975). Renfro-type dolomite lithologies
can also be recognized in the basal Newman Limestone
(“Greenbrier” or “Big Lime”) in many boreholes of
eastern and southeastern Kentucky. They also occur

ROCKS OF MOSTLY OSAGEAN AGE 43

 

FIGURE 23.—Cowbell Member of Borden Formation (siltstone and shale). Milepost 145, Interstate Highway 64. Cranston quadrangle,
Rowan County.

sporadically along the Pine Mountain outcrop belt,
southeastern Kentucky, verified by Dever at three sites:
the Burdine quarry north of Jenkins, in Hurricane Gap
near Cumberland, and in an abandoned quarry at the
head of Limestone Branch in Frakes quadrangle.
Renfro-type dolomite is also present on Interstate
Highway I-75 south of Jellico, Tenn. (Sedimentation
Seminar, 1981) and at Cumberland Gap, Va. Subsurface
studies of the “Big Lime” (lower and middle Newman
Limestone) in eastern Kentucky (Pear, 1980) and
southeastern Kentucky (Hetherington, 1981) indicate
that dolomite facies possibly equivalent to Renfro
dolomite are intermittently present in eastern Kentucky
and more persistent in southeastern Kentucky in the
basal and lower parts of thick sections of the “Big
Lime.” These possible Renfro equivalents reach a
thickness of more than 30 ft, and contain petroleum and
gas reservoir beds.

The upper contact of the Renfro according to Weir
and others (1966) is a conspicuous diastem. Inter-
tonguing of Renfro dolomitic strata with the overlying
St. Louis dense micritic limestone has been reported in
the Wildie (Gualtieri, 1968a) and Brodhead (Gualtieri,

 

1967b) quadrangles, but subsequent examination of
these areas by Dever and others (197 9d) suggests that
the reported intertonguing may simply reﬂect the fact
that discrete bodies of Renfro-er dolomite occur in the
St. Louis Limestone Member of the Newman, and con-
versely, interbeds of micritic limestone like that in the
St. Louis occur in the Renfro. Such occurrences appear
to indicate ﬂuctuations in depositional environments
during Renfro and subsequent St. Louis time without
representing true intertonguing along contemporaneous
time boundaries. Subsurface studies in eastern and
southeastern Kentucky (Pear, 1980; Hetherington, 1981)
suggest that basal dolomitic beds of the “Big Lime”
were deposited in submarine channels cut into the
underlying Grainger sediments or in other submarine
topographic lows. In the east-central Kentucky outcrop
belt, the Renfro thins northeastward; the overlying St.
Louis Limestone Member maintains a relatively con-
stant thickness. The sharp contact between the Renfro
and the St. Louis in east-central Kentucky may repre-
sent a hiatus, but observations in recent years suggest
that it is a diagenetic contact between dolomitized and
undolomitized limestone. Lobate patterns shown by

44 MISSISSIPPIAN ROCKS IN KENTUCKY

 

FIGURE 24.—Halls Gap (dark) and Muldraugh (light) Members of Borden Formation, looking west. Inclined bedding along front of Borden
deltaic sediments in center of photograph. Railroad cut northwest of Kings Mountain. Halls Gap quadrangle, Lincoln County.

isopach maps of the underlying detrital deposits in east-
central and south-central Kentucky suggest that the
topography on the top of the Grainger may be mainly
depositional rather than the result of erosion.

The major part of the Renfro is reported to merge
southwestward with the lower and middle parts of the
St. Louis Limestone of south-central Kentucky (Lewis
and Taylor, 1979; Dever and Moody, 1979a; Dever,
McGrain, and Moody, 1979). Laterally, the lower and
middle Renfro merges westward with the Muldraugh
Member of the Borden Formation and with the Salem
and Warsaw Limestones in the vicinity of the Crab
Orchard and Brodhead quadrangles (Gualtieri, 1967a,
1967b); rock types in this interval in the Crab Orchard
quadrangle in particular contain lithologies typical of
these units and also of the Harrodsburg Limestone.
Because Renfro lithologies characterize the unit west-
ward to the Halls Gap quadrangle, an arbitrary west
limit of the unit was placed at the Halls Gap—Crab
Orchard quadrangle boundary (Weir, 1972).

 

Fossils in the Renfro are mainly corals, crinoids,
and brachiopods. Butts (1922) reported on collections
in the present Brodhead quadrangle from Renfro
beds which he termed Warsaw Formation, and the
fauna was reported by him to have both Osagean
and Meramecian affinities. Foraminifera from the
middle part of the Renfro in the Bighill quadrangle
(Weir and others, 1971), were reported by BA. Skipp
(written commun., 1966) to include endothyrid species
generally considered to be of Meramecian age. Con-
odonts from the Renfro in the Wildie quadrangle
(J.W. Huddle, written commun., 1966) indicate that
the collections seem to be representative of the
Gnathodus texanus—Apatognathus zone (early Val-
meyeran or Osagean) of Collinson and others (1962)
and also of their Taphrognathus van‘ans—Apatognathus
zone (late Valmeyeran or Meramecian). Thus the
Renfro would seem to represent late Osagean to
early Meramecian ages, in agreement with Butts’
conclusions.

ROCKS OF MOSTLY OSAGEAN AGE 45

 

 

FIGURE 25.—Nada Member of Borden Formation (Mbn) and overlying Slade Formation consisting of Renfro Member (Msr) and limestone
and dolomite mostly of the St. Louis and Holly Fork (Ettensohn and others. 1984) Members (Mssh). Milepost 146.2, Interstate Highway

64, Cranston quadrangle, Rowan County.

MULDRAUGH MEMBER

The Muldraugh Member of the Borden Formation
(Weir and others, 1966), is 50 to 100 ft thick in southern
central and south-central Kentucky and as much as 300
ft thick in the western central Kentucky outcrop; it is
the main unit in the uppermost Borden in these areas
and in adjoining Indiana. Highly resistant, it is
dominantly an olive gray, siliceous dolomitic siltstone
(dolosiltite) or silty dolomite (fig. 26). Silica and car-
bonate (dolomite and calcite) are main matrix minerals
and constitute as much as 85 percent of the rock.
Quartz-lined geodes and concretions are locally
common. Fine-grained components are highly biotur-
bated. Crinoidal calcirudite lenses and patches of
skeletal limestone composed dominantly of crinoid and
bryozoan debris (fig. 27) are also common. Large crinoid
stem fragments reaching to more than an inch in
diameter are similar to those in the Fort Payne Forma-
tion and in limy lenses of Borden clastic units. The size
and the robust character of brachiopod shells and
crinoid columnal fragments are distinctive features of
these units in Kentucky, and are not typical of younger
Mississippian units.

 

The lower contact of the Muldraugh Member with the
Nancy, Halls Gap, and Holtsclaw Members is common-
ly abrupt and marked by one or more thin glauconitic
beds that have been previously ascribed to the Floyds
Knob Formation by Stockdale (1939). These contact
features suggest a depositional hiatus, followed by
sudden resumption of deposition of mixed clastic and
carbonate strata.

FLOYDS KNOB BED

The Floyds Knob Formation of Stockdale (1931,
1939), commonly a thin unit of glauconite or limestone
and glauconite, was designated the Floyds Knob Bed
of the Muldraugh Member of the Borden Formation in
western central Kentucky by Kepferle (1977a). There,
it is a widespread unit and an important stratigraphic
marker within the Borden. In that area, for example,
mapping of the abrupt southwestward stratigraphic
drop of the Floyds Knob Bed relative to the base of the
Borden enabled Peterson and Kepferle (1970) to
recognize a northwest-southeast-trending foreset slope
of the deltaic front which marks the southwestern limit
of thick Borden detrital deposits of shale and siltstone.

46 MISSISSIPPIAN ROCKS IN KENTUCKY

 

~ .3‘

a

4. 1

FIGURE 26,—Dolosiltite and calcarenite in Muldraugh Member of the Borden Formation. Milepost 96.8, Interstate Highway 65, 2.5 mi north
of Elizabethtown exit, Elizabethtown quadrangle, Hardin County.

To the east, in east-central and northeastern Kentucky,
two or more seams of glauconite occur in the upper part
of, or at the top of, Borden terrigenous clastic strata.
Southward, in south-central Kentucky, glauconite is
also common in the uppermost part of the New Prov-
idence Member of the Fort Payne Formation. The strati-
graphic intervals represented by these occurrences are
here considered to be approximately correlative with the
Floyds Knob Bed of central Kentucky, in general agree-
ment with the model proposed by Whitehead (1976,
1978b); and they are herein referred to as the Floyds
Knob. Rare exposures in the southwestern part of Pine
Mountain, southeastern Kentucky, show a glauconite
bed, possibly the Floyds Knob, underlying the Fort
Payne Chert.

The Floyds Knob Bed of the basal Muldraugh in
western central Kentucky consists of a single glaucon-
itic layer or two glauconitic layers separated by 5—25 ft
of phosphatic, siliceous silty dolomite, dolosiltite, clayey
siltstone, and, locally, oolitic limestone (Kepferle, 1977a,

 

1979b). The two glauconitic layers converge into one
layer along the Borden delta front (Kepferle, 197 9b). The
Floyds Knob and the overlying lower part of the
Muldraugh in northern central Kentucky and southern
Indiana have been studied intensively by Whitehead
(1976, 1978a). In southwestern central Kentucky, the
Floyds Knob is a glauconitic zone, 5—7 ft thick, with the
glauconite concentrated in seams and disseminated in
the siltstone between seams (Moore, 1977). To the east,
where the unit is included in the uppermost Nancy and
Halls Gap Members, it is 5—10 ft thick and consists of
one to three glauconite seams interbedded with silty
shale, glauconitic siltstone, and, locally, crinoidal lime-
stone; it contains chert and geodes (Lewis and Taylor,
1971; Lewis and others, 1973; Weir, 1972).

The Floyds Knob in the uppermost Nancy Member
grades eastward and northeastward into the Wildie
Member, which consists of as much as 25 ft of shale and
siltstone with persistent seams of commonly phosphatic
glauconitic siltstone and limestone at the top and base

ROCKS OF MOSTLY OSAGEAN AGE 47

 

FIGURE 27,—Upper part of Muldraugh Member of Borden Formation (Mbm) showing westward-thickening calcarenite lenses. Harrodsburg
Limestone (M h) (uppermost light-hued, planar-bedded resistant ledges) at upper left side of photograph. Milepost 3.3, north side of
Bluegrass Parkway. Elizabethtown quadrangle, Hardin County.

of the member. Northeast of the Wildie-Nada boundary,
the equivalent of the Wildie or Floyds Knob forms the
upper part of the Nada Member, with persistent seams
of glauconitic siltstone and limestone occurring at the
top and near the middle of the Nada in southwestern
east-central Kentucky. Discontinuous beds of glaucon-
itic siltstone and limestone also occur in the lower part
of the member (Rice, 1972; Weir and others, 1971).
In east-central and northeastern Kentucky, the
Floyds Knob Formation as deﬁned by Stockdale (1939)
was restricted to the glauconitic seam and associated
limestone which occur at the base of the Wildie and ex-
tend into the middle part of the Nada. Stockdale iden-
tified the Floyds Knob as the most persistent of the
several glauconitic seams present in the upper Borden
of the eastern outcrop belt. Local occurrences of a
glauconitic siltstone near the middle of the Nada were
noted during mapping in central and northeastern east-
central Kentucky, for example in the Stanton and
Olympia quadrangles (Weir, 1974a; McDowell and Weir,
1977). In northeastern Kentucky, a phosphatic,
glauconite-rich seam is present at the base of the Nada
(Philley and others, 1975) and was considered cor-
relative with the glauconitic siltstone near the middle

 

of the Nada in east-central Kentucky (see Stockdale,
1939, pl. 16). A glauconite-rich shale present about 4 ft
below the top of the Nada in exposures along Interstate
Highway 64 in eastern Rowan County and western
Carter County (Chaplin, 1980) may be correlative with
the glauconitic seam at the top of the Nada and Wildie
in southwestern east-central Kentucky.

The areal extent of the Floyds Knob on the foreset
slope of the Borden delta front and in the basin south-
west of the front is not completely known. Peterson and
Kepferle (1970) mapped the glauconite zone down the
delta front in at least part of western central Kentucky
and speculated that its local absence probably resulted
from scouring by currents. In Taylor County, southeast-
ward along the front, a core studied by the Sedimenta-
tion Seminar (1972) showed that the Floyds Knob
extends southwest of the Borden delta front and under-
lies basinal deposits of the Fort Payne Formation. Far-
ther to the southeast, the distribution of the glauconitic
shale in the Delmer quadrangle (Lewis, 1971a), Pulaski
County, indicates that the Floyds Knob is present on
the upper part of the front, but at least locally absent
on the lower part. It appears to merge southwestward
with the Maury Formation equivalent at the base of

48 MISSISSIPPIAN ROCKS IN KENTUCKY

the Fort Payne Formation (R.C. Kepferle, written
commun., 1985).

The Floyds Knob Bed and equivalents are fossilifer—
ous, containing pelmatozoans, brachiopods, bryozoans,
gastropods, pelecypods, cephalopods, trilobites, con-
ulariids, ostracodes, and fish remains (Kepferle, 1979a;
Stockdale, 1939; Weir, 1967; Weir and others, 1971).
Foramin ifera were studied by Conkin (1954, 1960),
and conodont fauna are reported by Weir and others
(1971) and by Whitehead (1978a).

Study of Floyds Knob glauconites by the Sedimen-
tation Seminar (1972) showed that the glauconite occurs
as dark-green microcrystalline pellets, up to 0.5 mm in
diameter, which are concentrated along bedding planes
and in burrows. The ovoid pellet shape and close asso-
ciation with intense bioturbation suggest a fecal-pellet
origin. Concentration of the Floyds Knob glauconites
in a thin, widespread zone suggests accumulation dur-
ing a period of decreased sedimentation.

GRAIN GER FORMATION

The Grainger Formation (Keith, 1895), exposed along
Pine Mountain in southeastern Kentucky, is from 200
to 500 ft thick. Lithologies of the Grainger, largely shale
and siltstone, are strikingly similar to those parts of the
Borden Formation; and the formations probably in-
tergrade or interlense laterally in the subsurface west
of the Pine Mountain fault. The Grainger is character-
istically gray and greenish-gray shale and siltstone
similar to the Nancy and New Providence Members of
the Borden, and commonly contains an upper unit of
resistant gray, greenish-gray, and reddish-gray silt-
stone, shale, and sandstone similar to those strata in
the Nada Member of east-central Kentucky. Southward
the overall grain size of the Grainger strata seems to
diminish, and in Tennessee and Virginia, adjoining the
southernmost part of the southeastern Kentucky out-
crop belt, quadrangle descriptions report variegated
shale as the main or only lithic component of the
Grainger (Englund, 1964b). Locally, cherty beds which
are Fort Payne equivalents are mapped as uppermost
Grainger beds (Englund, Landis, and Smith, 1963;
Englund, 1964b; Englund and others, 1964). The
Grainger is the equivalent of the Nancy Member of
south-central Kentucky and the New Providence Shale
Member of central Kentucky.

The Grainger is overlain by the Newman Limestone
or the Fort Payne Chert. It is separated from the
Newman by a relatively sharp diastem. Basal beds of
the Newman are mostly micritic limestone character-
istic of the St. Louis Limestone Member, but include
local dolomitic limestone and dolomite, similar to those
in the Renfro Member, and also to the Little Valley

 

Limestone of adjoining Virginia (Wallace de Witt, J r.,
oral commun., 1965). The Fort Payne Chert, 0—20 ft
thick, in the southwestern part of the Pine Mountain
outcrop belt (Rice and N ewell, 1975; Newell, 1975; Rice
and Maughan, 1978), is considered to be a thin distal
eastward extension of the Fort Payne Formation. Ex-
posures are largely of chert residuum; in fresh ex-
posures, chert together with dolomitic limestone is
locally interbedded with shale and sandstone (Englund,
1969). Although these beds have been interpreted to in-
tertongue with the Grainger (Englund, 1969), it seems
more likely that the Grainger—Fort Payne contact is a
discrete interface marked by glauconite much the same
as the base of Muldraugh in central Kentucky, as sug-
gested by Whitehead (1976).

In southeastern Kentucky, although glauconitic beds
are not reported in the quadrangles mapped along Pine
Mountain, glauconite is abundant in beds at the top of
the Grainger Formation and in basal beds of the overly-
ing Newman Limestone succession along US. Inter-
state 75 near the Kentucky-Tennessee border in J ellico
West quadrangle. A glauconite bed there was identiﬁed
as the Floyds Knob Bed (Sedimentation Seminar, 1981).
Glauconite has also been identified by Dever at two
locations in Frakes quadrangle. Hasson (1973) reported
a glauconitic zone in the type Grainger area in Ten-
nessee. Whitehead (1976, p. 209—231) considered that
this is part of a single glauconite zone which may con-
sist of one or more beds occurring in the Borden (Floyds
Knob), Grainger, and Fort Payne, and which extends
over a very large region of the eastern and central
United States. Whitehead’s conclusion was that the
glauconitic beds are related to major eustatic sea-level
rise in late Osagean time.

FORT PAYNE FORMATION

Rocks of the Fort Payne Formation (Smith, 1890)
make up most of the interval between the underlying
Chattanooga Shale, Maury Formation equivalent, or
New Albany Shale, and the overlying Warsaw or
Salem and Warsaw undivided, Harrodsburg Limestone,
and Ullin Limestone (Lineback, 1966) in a broad belt en-
compassing southern and western Kentucky and parts
of Tennessee, southern Illinois, and southwestern In-
diana. The Fort Payne and Ullin are interpreted to be
successively younger units that onlap Borden delta
detrital rocks in Illinois and Indiana (Lineback, 1966).
In Kentucky, similar relationships indicate that both
the Muldraugh Member of the Borden Formation and
the Fort Payne are younger than Borden detrital rocks.
The Muldraugh is contemporaneous with at least the
lower part of the Fort Payne, but beds in the upper part
of the Fort Payne in the western part of south-central

ROCKS 0F MOSTLY OSAGEAN AGE 49

Kentucky and in the subsurface of west-central and
western Kentucky may be younger than the
Muldraugh. Interpretations of Fort Payne and
Muldraugh basal relationships in central Kentucky (fig.
28) have ranged from intertonguing to discrete erosional
or nondepositional surfaces.

In Kentucky, the Fort Payne Formation is mostly a
drab gray siliceous dolostone and dolosiltite, crinoidal
limestone and chert, with lesser siltstone and shale, and
locally thick sandstone. It is mapped in south-central,
southern west-central, and western Kentucky. Four for-
mal members, the New Providence Shale Member
(Kepferle and Lewis, 1974), the Cane Valley Limestone
and Kniﬂey Sandstone Members (Kepferle and Lewis,
1974), and the J abez Sandstone Member (Kepferle and
others, 1980), and the informal Beaver Creek limestone
member (Klein, 1974) are exposed in south-central Ken-
tucky. A thin cherty unit ascribed to the Fort Payne

 

 

 

 

 

 

 

  
 

 

 

 

 

 

 

 

 

GREENSBURG QUADRANGLE SALOMA QUADRANGLE RAWVICK QUADRANGLE
(Taylor and others, 1968) (Moore, 1976) (Kepferle, 1973)
St. Louis Limestone St. Louis Limestone
'ﬁ' 1 I'l'_‘ Eroded
‘.1 .1 -‘ .1 cu
C
2
Salem and 8
Warsaw Limestone E
Shale member Salem Limestone g—E/E
:— —i_"— Harrodsburg Harrodsburg
Shale :LL— Limestone Limestone
n;
C :m Halls Gap
9 L' -I—'-- Muldraugh Member
E imestone ._ Member
'5
"L C Muldraugh _,__ ‘
3:) E ,9 Member
1: g Gap ‘5
3 H— Member LL
5 E
Siltstone g '9
and 0 8
shale 1::
Not exposed
. New Providence
Nseriggragﬁggf Shale Member
‘1‘

(Rice and Newell, 1975; Newell, 1975; Rice and
Maughan, 1978) is recognized in the Pine Mountain area
of southeastern Kentucky.

Most of the Fort Payne in south-central Kentucky
consists of generally uniform dolosiltites that are olive
gray to medium gray, very fine grained, spally, and
slightly argillaceous (fig. 29); their silicification ranges
from irregular patches to almost complete replacement
of patches of crinoidal and bryozoan debris. Evidence
of intense bioturbation is common in south-central Ken-
tucky. In a matrix of generally fine dolomite rhombs
and clay minerals (illite, chlorite, and kaolinite), the
framework is mostly quartz silt and sand grains, and
minor glauconite pellets and fossil fragments. Many
dolomite rhombs also appear to be detrital. Geodes and
blebs are locally common, containing mostly quartz,
gypsum, and anhydrite with minor sulfide minerals.
Most of the Fort Payne in south-central Kentucky

 

 

 

 

 

     
 
     
   

 

 

 

 

 

 

 

 

FIGURE 28.—Generalized diagram showing interpretations of relationships of Salem-Warsaw, Salem, Harrodsburg, Borden, and Fort Payne
Formations in three quadrangles in central and south-central Kentucky (from S.L. Moore, 1976).

 

 

FIGURE 29.-Fort Payne Formation in roadcut along launching ramp
road, Lake Cumberland State Park, Jamestown quadrangle, Russell
County.

closely resembles the Muldraugh Member of the Borden
Formation but is somewhat darker hued. It ranges frOm
about 200 to more than 300 ft thick in south-central
Kentucky.

In western and southwestern Kentucky the Fort
Payne is thicker (to more than 600 ft), darker, more
siliceous and cherty (as much as 40 percent chert), is
interbedded with limestone (Fox and Olive, 1966) rather
than dolomite, and is generally more planar bedded than
the ﬁghter hued, highly bioturbated succession in south-
central Kentucky. The westernmost Kentucky Fort
Payne is interpreted to represent a basinal facies and
the central Kentucky unit to represent platform slope
and platform facies, roughly similar to the interpreta-
tion of deposition on the Borden platform by Benson
(1976).

The lowermost unit of the Fort Payne Formation in
south-central Kentucky, the New Providence Shale
Member (Kepferle and Lewis, 1974), thins westward
from 150 to a few feet thick. Lithologically like the New

 

MISSISSIPPIAN ROCKS IN KENTUCKY

Providence Member of the Borden Formation and
commonly including basal phosphatic claystone iden-
tical to the Maury Formation equivalent, it represents
the distal foreslope and bottomset beds of the Borden-
Grainger deltaic deposits. A glauconite-rich bed, rare-
ly reported in south-central Kentucky geologic
quadrangles, at least locally separates the siliceous Fort
Payne beds from the underlying New Providence
Member. This is correlated with the Floyds Knob Bed
of western central Kentucky.

Within the dominant siliceous and dolomitic rock
types in the south-central Kentucky outcrop area, the
Fort Payne contains elongate lenses of light-hued
crinoidal calcirudite limestone with characteristic ac-
companying green-tinged mudstone. The largest of
these lenses, the Cane Valley Limestone Member, is as
much as 150 ft thick. It lies west of southeast-trending
lenses of sandstone more than 250 ft thick, the Kniﬂey
Sandstone Member (Kepferle and Lewis, 1974). The
Beaver Creek limestone member of the Fort Payne
(Klein, 1974) is another calcirudite unit which is locally
a petroleum reservoir. These limestone and sandstone
lenses occur parallel to and west of the southeast-
trending Borden delta front (fig. 54). They are inter-
preted to represent two types of carbonate banks, as
mud mounds and as skeletal sands (Lewis and Potter,
1978). The Cane Valley and Kniﬂey Members have been
intensively studied (Sedimentation Seminar, 1972).

Beds in the Cane Valley Limestone Member inter-
finger with Fort Payne dolosiltites and locally inter-
finger eastward with the Kniﬂey Sandstone Member.
Limestones are dominantly biorudites, coarse-grained
skeletal bryozoan and pelmatozoan limestones with
minor quartzose sand grains in a matrix of sparry
calcite and minor microcrystalline dolomite. Partial
replacement by silica is common. Depositional dips are
mostly southwestward but are locally bimodal.
Greenish-gray mudstone occurs as conspicuous inter-
beds and as matrix between fossil fragments. Other
largely northwest trending related limestone lenses as
much as 100 ft thick exposed in the Fort Payne of south-
central Kentucky are composed largely of echinodermal
debris and include large crinoid stem segments more
than an inch in diameter (Taylor, 1962, 1964; Thaden
and others, 1961). The Cane Valley has been interpreted
to be a slope edge or platform edge shoal bank deposit
into which fossil debris on the deltaic platform to the
east was transported, partly winnowed, and concen-
trated (Sedimentation Seminar, 1972, p. 20; Lewis
and Potter, 1978). It is considered mostly younger than
the Knifley Sandstone Member, and formed during an
interval when coarse elastic terrigenous sedimentation
was weak. Cane Valley depositional types, mostly
occupying the middle part of the Fort Payne, extend

ROCKS OF MOSTLY OSAGEAN AGE 51

northwestward in the subsurface and may have accumu-
lated as a deeper water submarine bank or fan deposit.

The Beaver Creek limestone member, similar
lithologically to the Cane Valley Member, occurs in the
lower part of the Fort Payne down-dip from the Cane
Valley Limestone Member bodies. It consists of coarse-
ly crinoidal biohermal-type deposits of varying grain
size and bedding thicknesses. Klein (1974) interpreted
the deposits to be a series of submarine fans deposited
by ﬂow mechanisms in approximately 300 ft of water
at the toe of a prograding carbonate platform.

Sandstone in the Kniﬂey Member is fine-grained,
silty, argillaceous subgraywacke with angular grains,
dominantly an illite-sublitharenite (Kepferle, 1977a).
Mineral composition is about 30 percent clay and 6 per-
cent dolomite matrix within a framework of 90 percent
quartz, 5 percent feldspar and micas, rock fragments,
glauconite, and opaque minerals. West to southwest
depositional dips are dominant (fig. 16); grain size
generally increases upward, but the bedding in the
Kniﬂey is poorly defined because of intense bioturba-
tion. The Kniﬂey has been interpreted to be a slope
break shoal deposit (Sedimentation Seminar, 1972). The
J abez Sandstone Member of the Fort Payne (Kepferle
and others, 1980) is a body similar to and along strike
with the Knifley in south-central Kentucky.

Subsurface well logs indicate that a dark shale facies
occurs in the Fort Payne—Borden succession along a
northwest-trending belt from about 5 to 15 mi wide and
80 mi long in the subsurface of the northern part of
western and west-central Kentucky (fig. 54). This belt
lies west of the Borden delta front and northeast of the
crinoidal limestone—greenish shale facies, along strike
of the Kniﬂey Sandstone. West of the dark shale facies
are dark siliceous basinal Fort Payne rocks. The shale
is generally described (Freeman, 1951, p. 167, 230) as
dark-gray to dark-brown limy shale or shaly limestone
in contrast to the hard siliceous shale and limestone that
are characteristic of the Fort Payne. Because the dark
shale facies lies along strike of the Kniﬂey Sandstone
Member, the shale may represent fines that were win-
nowed from Kniﬂey sands or which passed over or
around Kniﬂey sand deposits and settled west of and
marginal to the Borden delta slope break. Transporta-
tion agents may have been northwest-ﬂowing currents
along the delta slope.

High carbonate and silica content is characteristic of
Fort Payne rocks, but origin of the silica content is
uncertain. Analyses of the Fort Payne in south-central
Kentucky and Illinois show that quartz of silt and clay
size is an important although variable constituent of
these rocks (Sedimentation Seminar, 197 2, p. 4;
Lineback, 1966, p. 23). Silica content in the Fort Payne
has also been ascribed to chemical precipitation or

 

replacement during deposition and diagenesis or to
secondary causes such as weathering (Bassler, 1932, p.
154—155). Sponge spicules are also abundant consti-
tuents of some of these rocks (Gutschik, 1954).

STRATA OVERLYING BORDEN AND
FORT PAYNE FORMATIONS

Borden and Fort Payne strata, herein treated as in-
cluding the Renfro Member, are overlain conformably
by units mapped as Warsaw Limestone, Salem and
Warsaw Formations undivided, and Harrodsburg
Limestone in south-central and western Kentucky.
Isolated outliers of the Salem and Warsaw Formations
unit above more continuous Borden and Fort Payne
strata across the Cumberland saddle in south-central
Kentucky permit correlations west of the saddle with
those in eastern Kentucky. In eastern Kentucky, the
Salem and Warsaw are at least in part continuous with
strata of the Renfro Member, exposed along the western
margins of the Appalachian basin. In northeastern Ken-
tucky these units are absent because Pennsylvanian
rocks unconformably overlie Borden Formation. Fort
Payne rocks are locally overlain by Cretaceous and Ter-
tiary beds along the margins of the Mississippi Embay-
ment in western Kentucky (Olive, 1965; Wilshire, 1964).

The Grainger Formation and Fort Payne Chert are
overlain abruptly by lowermost beds of the Newman
Limestone along parts of the Pine Mountain outcrop
belt. In places, the basal beds of the Newman are
dolomitic limestones which are like those in the Renfro
Member of the Slade Formation in east-central Ken-
tucky; and they are here provisionally correlated with
the Renfro.

DEPOSITIONAL HISTORY OF THE
BORDEN AND FORT PAYNE FORMATIONS

That the Borden Siltstone of Illinois and the elastic
units of the Borden Formation in Kentucky and Indiana
were both deposited in a deltaic framework (Frund,
1953; Swann and others, 1965; Lineback, 1966; Weir and
others, 1966; Peterson and Kepferle, 1970) is well
established. According to Lineback (1966), crinoidal car-
bonate banks in western Illinois grew to heights of
200—300 ft above the sea ﬂoor of a deep-water basin
located in central and southern Illinois. As the Borden
deltaic complex advanced westward into Illinois, its
southward deflection by the carbonate banks on the
west determined the direction of growth of the long,
tongue-shaped Borden delta there. Lineback (1966) con-
cluded that sediment at the foot of the delta was
deposited in water depths exceeding 600 ft, and inter-
preted several sandstone bodies in Illinois to be

52 MISSISSIPPIAN ROCKS IN KENTUCKY

turbidites deposited largely on the pro-delta plain
(Lineback, 1968b). After active delta growth ceased,
dark siliceous carbonate rocks of the Fort Payne were
deposited in the adjoining deep-water basin and on
foreset slopes of the delta, partly filling depressions ad-
jacent to the delta. Although deltaic structure in the
Fort Payne has not been proven, convex-upward pro-
files of upper surfaces of the unit in Illinois where it is
thick (Lineback, 1966, p. 26-27) and some current struc-
tures suggest deposition of very fine detritals in en-
vironments also favorable to carbonate accumulation.
Progradation of the Borden delta into Kentucky,
however, probably preceded that of the Borden delta
in Illinois, as indicated by the biostratigraphic work of
Gordon and Mason (1985).

After deposition of the basinal Fort Payne, according
to Lineback (1966, 1969), an irregular submarine topog-
raphy in part with starved basin conditions was left in
southern Illinois. There, deep narrow depressions in the
sea ﬂoor were filled by Muldraugh-like sediments of the
Ullin Limestone, which overlapped the Fort Payne and
eventually onlapped terrigenous clastics of the Borden
delta. Cross-stratified fossil-fragmental carbonate
sediments (Harrodsburg) in the upper part of the Ullin
Limestone were deposited on a shallow-water platform
on the Borden delta, and filled in the adjoining depres-
sions, resulting in shallow water throughout the area.

Deltaic deposition of Borden terrigenous strata in
Kentucky was probably similar to that just described
for Illinois, except that the delta in Kentucky advanced
as a broad depositional shelf. Evidence includes wester-
ly and southwesterly depositional dips in Borden
detrital rocks in south-central and central Kentucky
(Weir and others, 1966; Kepferle, 1968), and abrupt
reciprocal thickness relationships between westward-
thinning Borden siltstone and shale and overlying
Muldraugh Member carbonates, which are separated by
a discrete depositional delta-front interface (Kepferle,
1966b; Peterson, 1966; Peterson and Kepferle, 1970).
The depth of water probably did not exceed 375 ft in
northern west-central Kentucky (Kepferle, 1977a, p. 38),
and restricting crinoid banks west of the encroaching
deltaic sequence were absent or else minor in thickness
and extent. Probable turbidite sandstone and siltstone
bodies include, in approximate decreasing age order, the
Farmers, the Kenwood, and the “Rockcastle freestone.”
Their emplacement was possibly triggered by seismic
shock (Kepferle, 1977a, p. 40). Discontinuous crinoidal
biostromal limestones, such as the Cane Valley Lime-
stone Member and Beaver Creek limestone member,
and the elongate bar-like sandstone bodies such as the
Kniﬂey Sandstone Member, lie parallel and marginal
to the main mass of Borden detrital rocks in west-
central and south-central Kentucky. Carbonate skeletal

 

material was washed from platform areas of the delta
into deeper water where it accumulated in bar-er and
fan deposits along the delta front, possibly on a shelf-
break on the delta slope and on the delta toe adjoining
the basin.

According to Hannan (1975), Fort Payne rocks in
south-central Kentucky represent deposits by gravity
ﬂow mechanisms along the southwest-facing clinoform
slope of a carbonate platform edge. The deposits in-
clude toe, slope, shallow water, and evaporite facies
of a southwestward-prograding carbonate platform
which succeeded the more actively prograding Borden
delta.

The Cincinnati arch area, “Cincinnatia” of Pepper and
others (1954), was probably submergent during most
or all of Early Mississippian time. Thicknesses of
Borden strata and their depositional vector directions
do not show evidence that such a barrier was present
during Borden—Fort Payne deposition (Whitehead,
1976, p. 269—270). It has been suggested that the Cin-
cinnati arch was slightly emergent during earliest
Mississippian time and even contributed minor
amounts of fine-grained sediment initially to the basal
Borden (Sable, 1970). That phosphatic nodule accumula-
tions likely caused by upwelling along a single shelf
were present both east and west of the arch (R.C.
Kepferle, written commun., 1985) argues against the
presence of an extensive emergent arch during any part
of Early Mississippian time.

The protoquartzites or subgraywackes prevalent in
detrital Borden units (Potter and Pryor, 1961) contrast
strongly with the clean orthoquartzite characteristic of
Upper Mississippian sandstones; Borden rocks are rela-
tively immature, and their mineralogy reﬂects a source
terrane of metamorphic and felsic plutonic rocks. Major
source areas for Borden sandstones have been inter-
preted as being east and north of the Appalachian basin
(Potter and Pryor, 1961; Walker, 1962), in the Canadian
Shield (Potter and Pryor, 1961), and possibly in north-
ernmost Canada (Swann and others, 1965, p. 15).
Whitehead (197 6, p. 267) presented evidence for an
easterly source area, and rejected distant northern
Canada sources. The volume and thickness trends of
detrital sediments transported into the Eastern Interior
and Appalachian basins indicative of transport direc-
tions also suggest to us that eastern rather than north-
ern Canada sources were dominant, perhaps located in
or east of the present Piedmont belt or the New England
Acadian mountain systems, and that the Canadian
Shield was not an important source area. Sources east
and northeast of the Appalachian basin contributed a
large quantity of sediment to the eastern United States,
but only the distal portions of this large volume of
detritus reached Kentucky.

ROCKS OF MOSTLY OSAGEAN AGE 53

Location of source areas for the fine detrital com-
ponents of the Fort Payne of southern and western Ken-
tucky is uncertain; if the silica contained in the Fort
Payne is largely clastic detritus, this might indicate dis-
tant topographically low eastern or southern sources.
Perhaps supporting the possibility of southern source
areas is the presence of relatively thick mudstones in
the basal Fort Payne in some places in southwestern
Kentucky and in Tennessee (fig. 9). Paleogeographic
reconstruction for late Borden—Fort Payne time (Sable,
1979a) indicates that shallow marine waters covered
most of the areas, but westward-deepening troughs ex-
isted in western Kentucky (figs. 54, 55, 56; Lineback,
1969, p. 123), later to be filled with the Fort Payne For-
mation, the Harrodsburg Limestone, and the Warsaw
Limestone of Butts (1917).

AGE

In the type Mississippian Mississippi Valley succes-
sion, the Meppen Limestone (Collinson, 1969) and Fern
Glen Limestone are considered to be of earliest Osagean
(Valmeyeran of Illinois) age, the Burlington Limestone
to be of middle Osagean, and the Keokuk Limestone to
represent late Osagean time.

Beds representative of earliest Osagean or Val-
meyeran time (Meppen in the Mississippi Valley (Collin-
son, 1969)) have not been recognized in Kentucky.
Megafauna in the lower part of the Borden of northeast-
ern and east-central Kentucky (New Providence group
of Butts, 1922) were assigned to Fern Glen and possibly
Burlington ages by Butts (1922, p. 50). In central Ken-
tucky, however, megafauna, foraminifera, and con-
odonts in the basal part of the Borden (New Providence
Shale of Butts) are reported as Burlington in age (Butts,
1917, p. 17; Conkin, 1957; Collinson and Scott, 1958;
Rexroad and Scott, 1964). Thus a considerable hiatus
representing early and part of middle Osagean time
seems to be indicated at the base of the Borden Forma-
tion in some Kentucky areas. Such a time gap would
represent the time taken for the Borden delta front, en-
croaching from the east or northeast, to reach central
and south-central Kentucky, time during which starved-
basin conditions existed west of the delta front. During
this time, only thin red and green mudstones in Ken-
tucky, central Indiana, and Ohio were interpreted to
have been shed from Cincinnati arch areas (Sable, 1970),
but even these may represent distal sediments from
sources farther east or northeast.

Fossils in the Muldraugh Member include crinoids,
brachiopods, and bryozoans. Identifications by Macken-
zie Gordon, Jr. (written communs., 1965, 1971) indicate
a late Early Mississippian (late Osagean) age.

 

Conodonts from the Floyds Knob in south-central and
east-central Kentucky (Science Hill, Delmer, Bighill,
Wildie, Mt. Vernon, Maretburg, Halls Gap, and
Yosemite quadrangles) and from the type Floyds Knob
Bed at Floyds Knob, Indiana, were identified by J .W.
Huddle (written communs., 1964, 1965; see tables 3—6).
Gnathodus texanus Roundy is relatively abundant in
those samples, and is representative of the Gnathodus
texanus—Taphrognathus zone (lower Valmeyeran or
Osagean) of Collinson and others (1962, fig. 66). Paleon-
tologic identifications by Whitehead (1976, p. 71—141)
of conodonts from the Muldraugh of Indiana and Ken-
tucky, including the Floyds Knob, confirmed Huddle’s
interpretation of the age of the Floyds Knob in central
Kentucky and Indiana as the Gnathodus texanus—
Taphrognathus zone. Whitehead placed overlying beds
of the Muldraugh in the Taphrognathus varians—
Apatognathus zone of Collinson and others (1962),
which ranges through the Warsaw Limestone and
Salem Limestone in Illinois and according to Nicoll
(1971) into the St. Louis Limestone in Indiana. The
Gnathodus texanus-Taphrognathus zone according to
Nicoll and Rexroad (1975, p. 16) ranges through the
Edwardsville Formation, Ramp Creek Formation
(Stockdale, 1929), and part of the Harrodsburg Lime-
stone in Indiana; and the Taphrognathus varians—
Apatognathus zone extends from the upper part of the
Ramp Creek and Muldraugh Formations (Muldraugh
Member of Kentucky) through the Harrodsburg and
most of the Salem, becoming rare in the St. Louis
Limestone. Thus this evidence suggests that the Floyds
Knob Bed and glauconitic zones of Kentucky are of late
Osagean (Keokuk) age and that the Muldraugh Member
of the Borden Formation is of late Osagean and early
Meramecian age. This is not reﬂected in the correlation
chart (fig. 7), which indicates this boundary to lie above
the Muldraugh Member, pending further verification of
the interpretation discussed herein.

Megafauna in the Fort Payne of south-central, cen-
tral, and west-central Kentucky was reported by Butts
(1922, p. 76—88) to be of late Osagean (Keokuk) age.
Butts obviously included strata now considered to be
in the Muldraugh. Few recent collections have been
made, and none shed more light on specific ages of the
Fort Payne within Kentucky.

PALEOTECTONIC IMPLICATIONS

Parts of the Ozark region, northeastern Missouri, and
western Illinois were uplifted during early Borden time.
Farther east, in Illinois, Indiana, and western Kentucky,
a differentially subsiding basin existed. Still farther
east, a stable, submerged shelf was present in eastern
Kentucky.

54 MISSISSIPPIAN ROCKS IN KENTUCKY

The Cincinnati arch, including the Nashville dome,
was a stable to slightly positive structure that
separated somewhat more rapidly sinking basins. The
LaSalle anticlinal belt of Illinois including its
southeastward extensions in Kentucky was stable to
slightly positive. After possible slight uplift early in
Osagean time, the Ozarks and their marginal areas in
western Illinois and Missouri became a slowly subsiding
platform. North of the area, the Transcontinental and
Wisconsin arches, probably emergent in early Osagean
time, were also slightly negative later in the interval.

A major negative feature in Early Mississippian
(early Osagean) time was a southward-deepening trough
that extended from the Michigan Basin southwestward
across central and southern Illinois and southward
under the present Mississippi Embayment (Pryor and
Sable, 1974, p. 293). This trough may have continued
into the deep Ouachita geosyncline of Arkansas and
Oklahoma.

OSAGEAN-MERAMECIAN SERIES BOUNDARY

The Osagean-Meramecian Series boundary has been
extremely difficult to recognize in the Eastern Interior
basin. Uncertainties of correlation result from disagree-
ments regarding specific ages of megafaunal assem-
blages at or near the boundary, and from imperfect
understanding of the complex depositional relations.
Although many published reports on the western part
of Kentucky and on Indiana, Tennessee, and Missouri
refer to the Osagean and Meramecian Series designa-
tions, the Illinois State Geological Survey has combined
these into a single series, the Valmeyeran. Most geolo-
gists agree that no regional hiatus marks the series
boundary in the Eastern Interior basin, and that deposi-
tion was generally uninterrupted, although in some
areas such as east-central Kentucky, depositional en-
vironments changed markedly.

Osagean-Meramecian Series boundary problems have
been mostly related to the age assignments and correla-
tions of the Warsaw Formation (or Warsaw Shale) near
its type area in western Illinois and in its eastward sub-
surface extension; of rocks termed Warsaw in Kentucky
and Tennessee; and of the Harrodsburg and Salem
Limestones in Indiana and adjoining Kentucky. Shales
and limestones of the Warsaw in western Illinois have
been variously assigned to the Osagean or Meramecian
or both (S. Weller, 1909; Butts, 1922; Moore, 1928; Van
Tuyl, 1925; Laudon, 1948; Wanless, 1957; Weller and
Sutton, 1940; J .M. Weller and others, 1948; Sando and
others, 1969). Subsurface work by Lineback (1966) in-
dicated that shale in the type Warsaw merges eastward
into the Borden Siltstone, thereby suggesting that the

 

shale is of Osagean age. However, Lineback’s cross sec-
tions indicate that part of the Warsaw may descend into
the Borden and part maintain a high position within the
upper Borden. As a discrete unit, the Warsaw is recog-
nizable only in western Illinois and a short distance
eastward into the Bordeh of the Illinois subsurface; thus
far, it has not been possible to establish conclusive
physical continuity or contemporaneity with the Ken-
tucky Mississippian succession.

Butts (1922) first used the name Warsaw in western
Kentucky for rocks overlying the Fort Payne chert or
Holtsclaw Sandstone of Butts (1915), currently the up-
permost Borden detrital rock units in Jefferson COun-
ty, Ky. The name Warsaw remains in current use in
parts of Kentucky and throughout Tennessee. In Ken-
tucky, in addition to this usage, Warsaw has been used
in a restricted sense for rocks corresponding to the “Up
per” Harrodsburg of Indiana (McFarlan, 1943, p. 75).
Because the type Warsaw in Illinois is neither directly
traceable into nor lithologically identical with the War-
saw of Kentucky and Tennessee, both usages of War-
saw in Kentucky and Tennessee, in the opinion of the
present authors, should be abandoned. The 1960—68
mapping in Kentucky confirmed Stockdale’s (1939)
opinion that Butts’ Warsaw unit includes lithologic and
age equivalents of the Salem (Meramecian), Harrods-
burg (Meramecian and Osagean?), and Muldraugh
(Osagean). Examples of varied interpretations in this
stratigraphic interval in central and south-central Ken-
tucky are shown in figure 28. Rocks called Warsaw in
Tennessee, even farther from the type Warsaw in II-
linois, are probably also correlative with part of the
Salem, the Harrodsburg, and the Muldraugh.

In much of east-central and northeastern Kentucky,
the Renfro Member of the Borden and Slade Formations
and Renfro equivalents in the basal part of the Newman
Limestone, as well as younger beds, directly overlie
Borden detrital rocks. In outcrop, the Renfro lies be-
tween Borden detrital rocks and upper St. Louis equiva-
lents in the Slade Formation, and is relatively thin. In
northeastern south-central and southwestern east-
central Kentucky, thin Harrodsburg-like crinoid- and
bryozoan-bearing, light-gray limestone lenses and beds
appear in the middle part of the Renfro and thicken
westward, as the lower and middle parts of the Renfro
grade respectively into the Muldraugh and the Salem
(Weir and others, 1966); the upper part of the Renfro
is absent in this crucial area because of Quaternary ero-
sion. The upper Renfro, however, has been demon-
strated to be equivalent to the lower and middle parts
of the St. Louis Limestone of south-central Kentucky
(Dever and Moody, 1979a).

Nicoll and Rexroad (1975) indicated that the Harrods-
burg Limestone and part of the Salem Limestone in

ROCKS OF LATE OSAGEAN AND MERAMECIAN AGE 55

Indiana contain conodont elements similar to those of
the Warsaw Shale and overlying Salem Limestone of
Illinois. These data suggest that the Harrodsburg-
Salem boundary of Indiana and Kentucky is only slight-
ly younger than the Warsaw-Salem boundary of Illinois.
Lineback’s (1966, p. 14) correlation equating the Western
Illinois Warsaw with part of the Borden Siltstone may
relate only to Illinois stratigraphy. On the basis of con-
odont evidence, the Warsaw Shale equates with the up-
per part of the Muldraugh and lower part of the
Harrodsburg in Indiana because of Taphrognathus
varians—Apatognathus zone forms in these units (Nicoll
and Rexroad, 1975, p. 16).

In the Cincinnati arch area erosion has removed most
of the rocks spanning the Osagean-Meramecian bound-
ary, so that all except the lower units in south-central
Kentucky are separated by at least 80 mi from their
western counterparts. Lithologic and faunal evidence,
however, indicates that all major units were once con-
tinuous at least across the Cumberland saddle in south-
central Kentucky. The extensive yellowish-weathering
carbonate rocks of the Renfro Member of the Slade
Formation and its lithic equivalents in the basal
Newman Limestone span the Osagean-Meramecian
boundary in east-central Kentucky but appear to be a
St. Louis equivalent, and thus are entirely Meramecian
in northeastern Kentucky (Dever, McGrain, and Moody,
1979, fig. 5.4). The Renfro grades westward into the
Salem, Harrodsburg, and Muldraugh, as well as into
the Salem and Warsaw Formations undivided unit in
south-central and central Kentucky (Weir and others,
1966).

N omenclatural history of the Harrodsburg Limestone
in Indiana is complex (Hopkins and Siebenthal, 1897;
Cumings, 1922, p. 493—499; Smith, 1965), and ages of
its faunal elements, which were compared to those in
the Illinois Warsaw and the underlying Keokuk Lime-
stone, have been strongly debated (Laudon, 1948). In
the Indiana outcrop, the Harrodsburg overlies the
Ramp Creek and Muldraugh Formations (usage of
N icoll and Rexroad, 1975), these latter two being equiv-
alent to the Muldraugh Member of the Borden Forma-
tion in Kentucky; the Harrodsburg underlies the Salem
Limestone. Currently, the siliceous and dolomitic car-
bonate rocks of the lower Ramp Creek Member of the
Harrodsburg of Stockdale (1929) are included in the
Muldraugh or Ramp Creek Formations, formerly units
of the Borden Group in Indiana. The younger Leesville
and Guthrie Creek Members (Stockdale, 1929) and an
uppermost unnamed division, the “Upper” Harrods-
burg, dominantly pelmatozoan crinoid- and bryozoan-
bearing skeletal limestone, have been retained in the
Harrodsburg by Nicoll and Rexroad (1975). The base
of the “Upper” Harrodsburg has been considered to be

 

the Osagean-Meramecian Series boundary in Indiana
Geological Survey reports since 1954, although the en-
tire unit had been assigned to the Osagean by others
(Stockdale, 1931, 1939). Nicoll and Rexroad (1975) in-
dicated that the conodont assemblage in the Harrods-
burg Limestone of Indiana and adjacent Kentucky
equates approximately with that of the Warsaw Shale
of Illinois. The Kentucky Harrodsburg is a close litho-
logic and time equivalent of the Indiana Harrodsburg
(Sable and others, 1966). The Kentucky Harrodsburg
as such is the practical mapping unit used in geologic
maps of western central Kentucky such as in the Fort
Knox (Kepferle and Sable, 1977), Colesburg (Kepferle,
1967), and Rock Haven (Withington and Sable, 1969)
quadrangles. '

ROCKS OF LATE OSAGEAN AND
MERAMECIAN AGE

Rocks of late Osagean and Meramecian age in Ken-
tucky are predominantly limestone, dolomite, and
Siltstone that accumulated on intracratonic platforms
and in shallow basins. Sandstone and mudstone, which
form relatively minor constituents in the succession,
mainly occur in south-central and western Kentucky.
Bedded evaporites in west-central Kentucky indicate a
period of shallow restricted seas and aridity during part
of Meramecian time. No regional hiatus is recognized
within the succession.

Principal lithologies in the stratigraphic succession
characteristic of Early to Late Mississippian (mostly
Meramecian) time are, in ascending order, fossil-
fragmental limestones deposited under moderately
high-energy subtidal conditions (Harrodsburg, Warsaw,
Salem, and the Salem and Warsaw Formations unit (fig.
30)); very fine grained chemically or organically
precipitated dolomitic carbonates and evaporites
deposited in quiet-water, subtidal to probable supratidal
environments (Renfro, St. Louis), and oolitic and
bioclastic limestones deposited in shallow-water,
moderately high energy subtidal environments (Ste.
Genevieve). Generalized relationships of these units are
shown in the cross section (fig. 9).

Besides the above units, two additional formational
units of predominantly carbonate rocks are the Mont-
eagle Limestone and the Slade Formation. These repre-
sent equivalents to part of the Meramecian succession
west of the Cincinnati arch as well as Chesterian
equivalents. The Monteagle Limestone, of Tennessee
derivation (Vail, 1959; Stearns, 1963), is restricted to
south-central Kentucky and includes units between
the St. Louis Limestone and the Hartselle Sandstone;
it comprises the Ste. Genevieve and Kidder Limestone

56 MISSISSIPPIAN ROCKS IN KENTUCKY

 

FIGURE 30.—Salem and Warsaw Formations undivided and lower part of St. Louis Limestone. Base of St. Louis shown by arrow. Upper-
most part of Muldraugh Member of Borden Formation (dark beds) at extreme lower left. Milepost 85.6, north side of Cumberland Parkway,

Delmer quadrangle, Pulaski County.

Members (Lewis, 1971b). The Slade Formation (Etten-
sohn, Rice, and others, 1984) in east-central and north-
eastern Kentucky includes the Renfro, St. Louis, and
Ste. Genevieve Members, 10 additional younger
members, and one named bed.

HARRODSBURG LIMESTONE

In Kentucky, the Harrodsburg Limestone (Sable and
others, 1966) consists of light-gray, crossbedded to
planar-bedded, pelmatozoan-bryozoan, relatively pure
biocalcirudite and biocalcarenite. It extends northwest-
ward from central Kentucky through Indiana to cen-
tral and southwestern Illinois. It is equivalent to the
“Upper” Harrodsburg unit of Indiana. Harrodsburg
equivalents are present in the Ullin Limestone of Illinois
and in the Warsaw and subsurface “Big Light” (drillers’
term) of western Kentucky. The Harrodsburg of cen-
tral Kentucky pinches out eastward and southward in
south-central and east-central Kentucky into the Salem
and Warsaw Formations unit, the upper part of the
Muldraugh Member of the Borden Formation (Weir,
1972; Lewis and others, 1973), and laterally correlative
rocks of the Renfro Member (Gualtieri, 1967a; Weir and
Schlanger, 1969). The Harrodsburg averages about 30 ft
thick in its outcrop belt in central Kentucky and
thickens westward in the subsurface to form part of the
more than 500 ft of “Big Light” or Warsaw Limestone
strata in western Kentucky.

 

WARSAW LIMESTONE

The name Warsaw Limestone is used in western Ken-
tucky for a 250 to 500 ft-thick unit of light- to medium-
gray biocalcirudite, biocalcarenite, and dolomitic
limestone which overlies the darker, siliceous beds of
the Fort Payne Formation. It is the “Big Light” of
western Kentucky drillers’ terms. Much of the Warsaw
comprises crossbedded relatively pure limestone like
that in the Harrodsburg Limestone. The lower part con-
sists of, in part, dolomitic and siliceous limestone,
similar lithologically to the Muldraugh Member of the
Borden Formation. The Warsaw, as mapped in western
Kentucky (Fox and Seeland, 1964), is considered to be
equivalent to the Ramp Creek, Muldraugh, and Har-
rodsburg of Indiana, the Ullin and Harrodsburg of II-
linois, and the Warsaw of western Tennessee. A detailed
treatment of the Warsaw in western Kentucky is given
by Trace and Amos (1984).

SALEM LIMESTONE AND
SALEM AND WARSAW FORMATIONS UNIT

The Salem Limestone, named for Salem, Washington
County, Indiana (Cumings, 1901), in Kentucky consists
of conspicuously crossbedded, medium- to coarse-
grained biocalcarenite composed of fossil-fragmental,
pelletal, and minor oolitic limestone interbedded with
micritic dolomitic limestone, calcareous mudstone, and

ROCKS OF LATE OSAGEAN AND MERAMECIAN AGE 57

minor sandstone. Biocalcarenite units range up to sev-
eral tens of feet thick and are dominantly biosparites
with varying amounts of argillaceous matrix. In a few
places they contain units of the well-sorted, winnowed
high-calcium biosparites like that of the Salem building
stone facies in Indiana; in Kentucky this lithology is
limited to very thin beds of endothyrid and
pelmatozoan-bryozoan biocalcarenite. For the most
part, dark argillaceous biocalcarenite and calcareous
mudstone typify the Salem in western central and
south-central areas of Kentucky; dark biocalcarenite
consisting of fossil fragments, endothyrid tests, and
oolitic and pelletal fragments is characteristic of the
Salem map unit in western Kentucky.

The Salem thickens irregularly from generally less
than 100 ft in its Indiana outcrop belt to more than
350 ft in southern Illinois. Strata in the Indiana outcrop
belt, except for the exposureless break across the Ohio
River, are directly traceable into Kentucky (Withington
and Sable, 1969; Kepferle and Sable, 1977).

In central and south-central Kentucky, the Salem
Limestone and Salem and Warsaw Formations unit
range from about 50 to 170 ft thick. In general, the unit
thickens westward, and the areas of greatest thickness
are in northwestern central and southwestern south-
central Kentucky. Meaningful correlation of the Salem
in these areas with strata in western Kentucky, how-
ever, is equivocal, because reliable subsurface informa-
tion across much of west-central and western Kentucky
is widely spaced. Based on a lithologic succession
roughly comparable to that in central and west-central
Kentucky, the Salem interval in western Kentucky, as
mapped in the Fredonia, Crider, and Princeton East
quadrangles, is only 120—130 ft thick. However, if the
Salem is expanded to include strata mapped as the
lower member of the St. Louis Limestone (upper part
of the Illinois Salem) which overlies this interval, the
total Salem thickness in western Kentucky is about
350 ft.

In south-central Kentucky, most of the rocks between
the Fort Payne Formation and the St. Louis Limestone
were mapped in the early stages of the mapping
program as Warsaw Limestone, for example, in the
Austin quadrangle (S.L. Moore, 1961). Later, as it
became apparent that beds equivalent to and litho-
logically like the Salem Limestone of central Kentucky
and southern Indiana were present in part of this unit,
it was designated the Salem and Warsaw Formations
(or Limestones) undivided map unit, as in the Holland
quadrangle (Nelson, 1962). This compound name does
not indicate that Salem and Warsaw beds are “lumped”
for ease of mapping; it designates one mappable unit
which is not practically divisible. The name Warsaw is
retained herein for areas in south-central and western

 

Kentucky because of its long history of usage. As
previously stated, however, the type Warsaw cannot
be directly traced from its type area in western Illinois
into Kentucky, and the lithologies in the widely
separated areas are not closely similar. We therefore
believe that the use of the name “Warsaw” in Kentucky
is inappropriate.

Biocalcarenite, calcareous mudstone, and argillaceous
limestone (Somerset Shale of Butts, 1922; Lewis and
others, 1973) are dominant lithologies in the Salem and
Warsaw Formations unit, but dolomitic limestone and
lesser amounts of sandstone (Garrett Mill Sandstone
of Butts, 1922, p. 107; Science Hill Member of the
Warsaw, Lewis and Taylor, 1979) constitute appreciable
parts in south-central Kentucky. Argillaceous rocks in
the unit there, such as in the Science Hill quadrangle
(Taylor and Lewis, 197 3) and in the Salem farther north
(S.L. Moore, 1976; Withington and Sable, 1969), are also
common in western central Kentucky and southern
Indiana. The Garrett Mill Sandstone Member in the
upper part of the Salem and Warsaw Formations unit,
and the Science Hill Sandstone Member in the basal
part are each less than 50 ft thick. Their distribution
(fig. 31) in the Cumberland saddle area is similar and
also in part overlaps the distribution of the older J abez
and Kniﬂey Sandstone Members of the Fort Payne
Formation. They show two directional distribution com-
ponents, north-northwest, parallel to the J abez-Kniﬂey
sandstone trend, and normal to that, east-northeast.
Apparent age-equivalent strata to the east, west, and
northwest are fine-grained terrigenous clastics with
limy admixture, ranging from calcareous mudstones to
calcareous siltstones. Pebbles occur in the upper part
of the Science Hill Member throughout its outcrop area.
Directional data for the upper part of the Science Hill
(Lewis and Taylor, 1979) indicate that the member was
a shallow subtidal deltaic sand body derived from an
eastern source. More detailed study of these sandstone
units and related fine clastics both in Kentucky and
Tennessee is needed to better understand the period of
transition from delta-dominated elastic deposition to
shelf-controlled carbonate sedimentation phases of
middle Mississippian time.

The basal strata of the Salem and the Salem and
Warsaw Formations unit in Kentucky—biocalcarenite,
argillaceous limestone, or sandstone—in some places
contrast strongly with the lithologies of the under-
lying units, the Harrodsburg, Borden, and Fort Payne;
but the contacts do not seem to indicate a major
depositional hiatus. In western central Kentucky,
the Salem-Harrodsburg contact is generally abrupt,
from relatively pure calcarenite upwards to calcareous
mudstone or shale, but as much as a few feet of argil-
laceous Salem calcarenite are transitional in some

58

89° 88° 87° 83°

MISSISSIPPIAN ROCKS IN KENTUCKY

 

 

l l I I
EXPLANATION

 

Jabez and Kniﬂey Sandston
Members

Science Hill Sandstone
Member

Dominantly calcareous silt-
stone and shale in Salem
and Warsaw Formations
unit

Garrett Mill Sandstone
Member

38°

INDIANA
ll.LlNOlS

37°

 

  
 
 
 
   
   
 
 
  
 
  

CHIC)
e

 

WEST
/\ “BOWL:

 

 

 

(1

100 MILES

 

FIGURE 31.—Approximate distribution of Science Hill Sandstone and

Garrett Mill Sandstone Members of Warsaw Formation and J abez

and Knifley Sandstone Members of Fort Payne Formation in south-central Kentucky.

exposures. In certain south-central Kentucky quad-
rangles, the Salem-Warsaw unit, dominantly argilla-
ceous to sandy biocalcarenite, locally contains beds that
are lithologically similar to the Harrodsburg.

As mapped in western and southwestern Kentucky,
the Salem is commonly restricted to strata containing
coarsely crystalline biocalcarenite with oolitic over-
growths, and pelletal calcarenite with the abundant
foraminifer Globoendothyra (Endothyra) baileyi (Trace,
1974). This definition of the Salem is more restricted
than that used by the Illinois Geological Survey (Baxter
and others, 1963, p. 7). According to Trace and Amos
(1984) the basal Salem contact corresponds to the
base of the middle unit of the Illinois Salem as used
by Baxter and others (1967, p. 9) in the Fluorspar
district adjacent to Kentucky. In Illinois, the Salem in-
cludes younger argillaceous limestone beds bearing
“endothyrids” and “lithostrotionoid”1 colonial corals,
which in western Kentucky are placed in the lower
member of the St. Louis Limestone. In addition, rocks
in the basal unit of the Illinois Salem have been assigned

1“Lithostrotionoid” coral species used as biostratigraphic indicators of
Mississippian strata in Kentucky have been referred to as Lithostrotion
castelnaui Hayasaka, 1936; Lithostrotion canadense Milne-Edwards and
Haime, 1851; and Lithostrotion [Lithostrotionella] proliferum Hall, all from the
St. Louis and Salem Limestones, and to Lithastmtion harmodites Milne-
Edwards and Haime from the Ste. Genevieve Limestone. The ﬁrst three above
have been reassigned to the genus Acrocyathus by Sando (1983). Sando (writ-
ten commun., July 1985) has referred specimens of Lithostration harmodites
to Schoenophyllum aggregatum Simpson, 1900.

 

to the Warsaw Limestone in western Kentucky. Line
back (1972), after subsurface study of the Salem and
St. Louis, indicated that the Salem and St. Louis litholo-
gies in Illinois and Indiana, based on relative abundance
of biocalcarenite versus dolomitic and evaporitic strata
in the Salem and St. Louis, respectively, are intergrada-
tional and intertonguing. The formational boundary,
therefore, has no regional time-stratigraphic signif-
icance. (See section “St. Louis and older unit relation-
ships,” p. 66.)

Both the upper part of the Salem (Kepferle, 1967) and
the basal St. Louis (Kepferle and Sable, 1977) in western
central Kentucky contain thick beds of micritic to cal-
carenitic, dolomitic limestone and dolomite similar to
the “finely granular argillaceous dolomitic limestone”
of the Salem in Indiana (Pinsak, 1957, p. 37 and pl. 5).
This development of dolomitic lithologies is considered
by Sable (1979b) to signal a precursor of the evaporite
environments in the succeeding St. Louis Limestone.
It is interpreted to represent a regressive trend which
resulted in intertidal and supratidal environments
encroaching over a large area of previously subtidal
deposition.

ST. LOUIS LIMESTONE

The St. Louis Limestone (Engelman, 1847) overlies
the Salem and the Salem and Warsaw Formations unit
in western, west-central, central, and south—central

ROCKS OF LATE OSAGEAN AND MERAMECIAN AGE 59

Kentucky. That geologic mapping was difficult because
clearly delineated contacts could scarcely be discerned
suggests that in many areas the contact is gradational
and intertonguing.

During the mapping program, rocks originally
equated with the St. Louis of central and western Ken-
tucky were assigned in east-central and northeastern
Kentucky to two formations (Dever and Moody, 1979a;
Dever, McGrain, and Moody, 1979). The dolomite and
interbedded limestone of the lower part of the St. Louis
were included in the Renfro Member of the Borden
Formation (Weir and others, 1966) together with older
rocks correlative with the Salem-Warsaw unit and
Muldraugh; limestone of the upper St. Louis was desig-
nated as the St. Louis Limestone Member of the
Newman Limestone (Hatch, 1964).

Most of the St. Louis Limestone in Kentucky is
similar to that in adjoining States, and consists of very
ﬁne grained micritic to lutitic carbonate rock with lesser
terrigenous material as dark carbonaceous shale and
grayish-green shale in the lower and middle parts of the

 

formation. Abundantly fossiliferous beds associated
with biocalcarenites and biorudites are common in the
upper St. Louis of eastern and south-central Kentucky.
Chert is a common to abundant accessory and occurs
as irregularly shaped masses, spherical nodules, discon-
tinuous beds, and replacement layers (“scraggy” or
“scraggly” chert) of limestone and dolomite (ﬁg. 32). In
some areas chert occurs in fairly discrete widespread
“zones.”

In west-central and central Kentucky, subsurface
beds of gypsum and anhydrite as much as 15 ft thick,
with associated dolomite, occur within a stratigraphic
interval of as much as 200 ft in the lower part of the
St. Louis; these are related to similar occurrences ex-
tending from west-central Illinois and south-central In-
diana into Kentucky (Saxby and Lamar, 1957;
McGregor, 1954; McGrain and Helton, 1964). The lower
and middle St. Louis of south-central and east-central
Kentucky contains multiple zones of brecciated dolo-
mite associated with quartz nodules, celestite, and
pyrite, which are considered to have formed during

 

FIGURE 32.—Cherty dolomite in upper part of St. Louis Limestone. J .F. Pace Construction Company quarry, Glasgow North quadrangle,
Barren County. Hammer for scale.

60 MISSISSIPPIAN ROCKS IN KENTUCKY

dissolution and replacement of evaporites (Dever and
others, 1978; this report, fig. 33). Beds of carbonaceous
shale containing marine fossils and plant debris,
reported in the lower part of the St. Louis in west-
central Kentucky (R.C. Kepferle, oral commun., 1966;
Withington and Sable, 1969), are in about the same
stratigraphic position as the evaporite beds. Carbonate
strata in this part of the section include laminated
limestone, very probably of stromatolitic origin, and
fine-grained limestone breccia beds.

The St. Louis as mapped shows a general eastward
thinning across the State. On the west side of the Cin-
cirmati arch, average thicknesses range from 475 ft in
western Kentucky, through 300 ft in west-central Ken-
tucky, to 230 ft in western central and south-central
Kentucky. The thinning seems to occur largely in the
interval assigned to the upper member of the St. Louis
north and west of the Caledonia and Johnson Hollow
quadrangles. The “Lithostrotionoid” coral-bearing
lower member averages about 250 ft in thickness in
western, west-central, central, and south-central
Kentucky. In western Kentucky where the St. Louis
map unit is 450—525 ft thick, it includes rocks assigned
to the lower part of the overlying Ste. Genevieve

 

 

Limestone in the outcrop belt east of Trigg County (fig.
34). On the east side of the Cincinnati arch, average
thicknesses of the St. Louis range from 125 ft in south-
central Kentucky to 25 ft in northeastern Kentucky.
Northeastward thinning along the outcrop belt is notice-
able principally in the sequence of dolomite and inter-
bedded limestone of the lower and middle St. Louis and
their equivalents included in the Renfro of east-central
and northeastern Kentucky. This dolomitic sequence
ranges in thickness from 75 ft in Pulaski County to 2 ft
in Greenup County. The distinctive, cherty upper
limestone of the St. Louis maintains a relatively uniform
thickness of 15—25 ft across the outcrop belt.

The St. Louis is absent in parts of northeastern, east-
central, and eastern Kentucky, the result of erosion
associated with Late Mississippian tectonic activity
along the Kentucky River fault system and Waverly
arch (Dever, 1973, 1977, 1980b). Subtidal St. Louis
sediments originally were deposited across the entire
area. Deposition was interrupted by uplift along the
Waverly arch, resulting in subaerial exposure and
vadose diagenesis of the subtidal limestone, but without
large-scale, extensive erosion. Diagenetic fabrics in-
dicative of exposure and vadose diagenesis are present

FIGURE 33.—Dolomite in middle part of St. Louis Limestone. Young man pointing to zone of nodular quartz and brecciated dolomite which
represents dissolved evaporite beds. Burnside Island State Park, Burnside quadrangle, Pulaski County.

ROCKS OF LATE OSAGEAN AND MERAMECIAN AGE 61

 

 

 

 

 

 

 

AREAS WEST AND NORTH OF CHARACTERISTlC FEATURES CALEDONIA QUAD-
N RANGLE AND AREAS
CALEDONIA OUADRA GLE TO EAST AND SOUTH
Levias
m Limestone Light—gray limestone,
: Member some oolitic, elastic, and
.3 very fine grained
d.)
.E
-‘ Rosiclare Sandstone,
3 Sandstone siltstone, and
.g Member Ste limestone
E Genevieve
8 Limestone
.- ,Fedonia mile.
6 imestone . . . _ . Ste.
Member Lisrtrld-s‘lgﬁe Thick oolltlc limestone Genevieve
Limestone
_ ~17); o—f z‘arTaBl' _ ' ' Chert occurs locally
abundant chert
a, Limestone, in part oolitic
E Upper
g member Top of zone of abundant nodular chart
E
3 Cherry limestone, in part oolitic
3
A Lower Top of abundant "lithostrotionoid" corals
(n member
Lower ° . St. Louis
member LTrhleLsfighs‘e Limestone and dolomite, chert common Limestone

 

 

 

 

 

 

FIGURE 34.—Generalized relationships between Ste. Genevieve and St. Louis Limestones as mapped in and near Caledonia quadrangle, western
Kentucky (from Ulrich and Klemic, 1966).

at the top of the St. Louis in the eastern outcrop belt
as far southwestward as northern Rockcastle County.
Erosional removal of the St. Louis and part of the
underlying Borden Formation followed recurrent move-
ment of the Kentucky River fault system during or im-
mediately after Ste. Genevieve time. Eroded areas are
mainly on the north (upthrown) side of the fault system.

The upper limestone unit of the St. Louis, mapped as
St. Louis Limestone Member of the Newman but cur-
rently in the Slade Formation, also is absent to the
south in parts of Estill and Jackson Counties (Rice,
1972; Haney and Rice, 1978); but the dolomite and in-
terbedded limestone representing the lower and middle
St. Louis are present. In adjacent areas, limestone of
the upper unit locally has convolute bedding, exhibiting
ball-and-pillow structure, and projects several feet into
the underlying dolomite.

On the Pine Mountain overthrust block, southeastern
Kentucky, an interval of cherty limestone and dolomite
in the basal Newman Limestone was identified on the
basis of lithology as St. Louis by Butts (1922) and
Hauser and others (1957), and as the Hillsdale Member
of the Greenbrier Limestone of Appalachian basin

 

terminology by Wilpolt and Marden (1959). The domi-
nant lithologies, micrograined limestone, very finely
crystalline dolomite and chert, are typical of the St.
Louis; and the interval commonly is overlain by Ste.
Genevieve-type calcarenite, although distinctive St.
Louis megafossils, such as “lithostrotionoid” corals,
have not been reported. The interval commonly is about
40 ft thick, but in the Pineville area, Hauser and others
(1957) and Butts (1922) assigned 80 and 115 ft, respec-
tively, of basal Newman to the St. Louis.

STE. GENEVIEVE LIMESTONE

The Ste. Genevieve Limestone (Shumard, 1860,
p. 406), the youngest Meramecian formation in Ken-
tucky, is composed principally of carbonate rocks: light-
colored, medium- to coarse-grained, oolitic and bioclastic
calcarenite; light-colored to gray, bioclastic calcirudite;
gray calcilutite; and gray, very ﬁnely crystalline
dolomite. The oolitic and bioclastic calcarenites are
commonly considered to be important lithologic criteria
for recognizing the Ste. Genevieve. Principal constitu-
ents in the bioclastic calcarenite and calcirudite are

 

62 MISSISSIPPIAN ROCKS IN KENTUCKY

pelmatozoans, brachiopods, and bryozoans. Clay shale
and sandstone are relatively minor constituents of the
formation, but they are more common in western Ken-
tucky, where they form named stratigraphic units.
Chert is a relatively minor accessory, occurring as
nodules, irregular masses, and siliceous replacements
of fossiliferous beds such as the Lost River Chert Bed
of Elrod (1899).

The Ste. Genevieve Limestone is a roughly tabular
unit from western central and south-central Kentucky
across west-central and western Kentucky. It averages
from about 180 to 240 ft thick in a distance of about
140 mi, thus suggesting that the rates of deposition and
subsidence were essentially in equilibrium west of the
Cincinnati arch.

In western Kentucky, the Ste. Genevieve averages
about 240 ft thick and ranges to more than 300 ft. There
it is divided into three members, in ascending order, the
Fredonia Limestone Member (Ulrich and Smith, 1905);
Rosiclare Sandstone Member (Ulrich and Smith, 1905);
and Levias Limestone Member (Sutton and Weller,
1932). The Fredonia Limestone Member is character-
ized by oolitic limestone including the drillers’ unit,
“McClosky oolite,” an important oil reservoir in some
areas. The Fredonia ranges from 170 to 260 ft and
averages 205 ft thick. In western Kentucky, 20—30 ft
below the top of the Fredonia, a possible shale and sand-
stone equivalent of the Spar Mountain Sandstone of
Illinois (Tippie, 1945) was noted by Amos (1967, 1974)
and Amos and Hays (1974). The Rosiclare Sandstone
Member, sandstone and green shale with local sandy
limestone, is 5—25 ft thick, averaging about 10 ft, in
western Kentucky. The overlying Levias Limestone
Member, finely crystalline to micritic to oolitic lime-
stone, ranges from 10 to 35 ft thick and averages
25 ft.

The three members are not mapped in the major part
of the Ste. Genevieve outcrop west of the Cincinnati
arch because the Rosiclare Sandstone Member sepa-
rating the Fredonia and Levias Limestone Members
grades into limestone southeast of Caldwell County and
thus does not remain a mappable unit. However, a
number of useful stratigraphic units are recognizable
within the Ste. Genevieve west of the arch, in ascending
order, the Lost River Chert Bed (Shaver and others,
1970); silicified oolitic limestone unit; Schoenophyllum
aggregatum (formerly Lithostrotion (Siphonodendron)
genevievensis) zone; Rosiclare Sandstone equivalent;
and Bryantsville Breccia Bed (Shaver and others, 1970).

The Lost River Chert Bed (Elrod, 1899), whose type
area lies in southern Indiana, is a distinctive unit of one
or more commonly silicified highly fossiliferous lime-
stone beds that ranges from 1 to about 10 ft thick. In
its type area it lies 20 or more feet above the Ste.

 

Genevieve’s base, as defined by oolitic limestone beds
(Malott, 1952, p. 8). When weathered this limestone is
a porous, siliceous coquinoid rock having a chalky
matrix in which the fossil fragments (relatively large
fenestellid bryozoan fronds, and large orthotetid,
productid, and spiriferoid brachiopod whole shells and
fragments) are outlined or filled with stain from bright-
red clay. In Kentucky this lithology is described in
western central Kentucky in the Howe Valley and Rock
Haven quadrangles (Kepferle, 1963a; Withington and
Sable, 1969), and in west-central Kentucky (fig. 35).
Similar rocks are described in several western Kentucky
quadrangles (Sample, 1965; Amos, 1974; Amos and
Hays, 1974). McGrain (1969) also recognized this chert
bed east of the Cincinnati arch in Wayne and Pulaski
Counties (fig. 59), where it serves as a useful strati-
graphic marker as much as 20 ft above the horizon that
separates Ste. Genevieve calcarenitic and oolitic lime-
stone lithology from the underlying lithostrotionoid
coral-bearing micritic limestone of the St. Louis (fig. 36).
In south-central Kentucky it is the uppermost part of
the Horse Cave Member of the St. Louis Limestone of
Pohl (1970). Such a siliceous layer is a valuable map-
ping tool in areas of low relief, poor exposure, and thick
residual soil. It is also recognizable in cuttings and cores
of bore holes (Kepferle and Peterson, 1964). Other
cherty limestone beds which weather to residuum
somewhat similar to that of the Lost River Chert Bed
occur both above (Withington and Sable, 1969) and at
(fig. 37) the Ste. Genevieve—St. Louis contact in central
and south-central Kentucky. The Lost River Chert Bed
is also locally absent in parts of south-central Kentucky,
and thus the other cherty units could be confused with
it during geologic mapping.

The Rosiclare Sandstone Member of the Ste. Gene-
vieve, as mapped in western Kentucky, grades south-
eastward from Caldwell County into a silty to sandy,
peloidal and intraclastic calcarenite. This calcarenite,
although not a mappable unit, has been found to form
a distinct bed, 1—10 ft thick, which can be traced,
through quarry and roadcut sections, across the Ste.
Genevieve outcrop belt from western Kentucky into
west-central, south-central, and central Kentucky where
it is considered to be the Rosiclare equivalent (Dever,
McGrain, Ellsworth, and Moody, 1979; Dever, 1980).
Both the calcarenite bed and the Rosiclare are within
the upper interval of the formation bracketed by the
Schoenophyllum aggregatum (formerly Lithostrotion
(Siphonodendron) genevievensis) zone and Bryantsville
Breccia Bed. Commonly associated with the calcarenite
are micritic crusts and stringers developed during
vadose diagenesis; identical diagenetic features occur
in the Bryantsville Breccia Bed at the top of the Ste.
Genevieve.

ROCKS OF LATE OSAGEAN AND MERAMECIAN AGE

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

2 v: FORMATION
E ‘é‘ AND LITHULOGV "1mg? 0 E s c R i P r l o N
a “6:“ MEMBER
'5‘
‘ - Limestone, shale, and chart breccia: Limestone, light- to medium-gray, finely to medium crystalline; few beds are sublithographic; medium to thick bedded; oolitic beds
, , , 15-25 local in upper part, some beds locally ‘ ' ' ' Shale, ' g“ g t. ' l , ay, in part ‘ thin' ‘ ‘ .Chert breccia, medium- to light-gray veryfina
1 T ' ' grained chert fragments in sparse to abundant matrix of light-gray sublitliogrepliic limestone; large chert fragments as much as 1 font in diameter consist of
well-cemented small angular cliert fragments of various shades of gray; at top of unit, locally absent; bed is the Bryentsvilla Breccia of Melett (1952,
l p. B).
J
t . " g.‘ , Limestone and sandstone: Limestone, very light gray to medium-gray, finely to coarsely crystalline, excellent abundant and well-developed white to light-gray oolites
‘ _I' i I' ‘l «‘73 60-80 common in upper half of unit in zones as thick as 15 feet; medium to thick bedded, less commonly thin bedded; crossbedding in coarsely crystalline beds;
‘ scattered gray fine-grained chart nodules that weather chalky white at and near top of unit; few slightly sandy beds about middle, dolomitic beds near top and
base; limestone breccie or intraforrnational conglomerate locally in upper part; Plalycrini'tes penicillus Meek and Worthen common. Sandstone, grey, very
calcareous, medium- to thick-bedded, in upper part; locally absent. Unit generally well exposed throughout map area.
135-195
1 i; i L r 15.20 Shale and limestone: Shale, dark—green to greenish-gray, slighty calcareous. Limestone, medium-gray, finely crystalline, medium- to thick-bedded,
It I t J t
Y 1' n . 1 . r
y I Tl
. r y l
|
ﬁfﬁ~ Limestone and dolomite: Limestone, very light gray to medium-gray, sublithograpliic to medium-crystalline; oolitic beds in upper and lower thirds; medium to thick
‘ , ‘ . ‘ .‘ 50-70 bedded, less commonly thin bedded and laminated; light-gray, fine-grained nodules of chart 15 to 25 feet above base; much of upper half of unit is dolomitic.
Dolomite, gray, finely crystalline, medium- to thick-bedded, near middle of unit.
2
:4 i I . 1 i
< Lost Riverl?) *J J ' , _
l ’, ,‘g,”’,r‘ ', . Limestone and chart: Limestone, very light ten to light-yellowrsh-gray, medium- to fine crystalline, medium- to thick-bedded; ebundem fenestellid bryozeans end
Chartof ,,;,, 514 . _ ‘
_ Elrod (1899) ’ ' ' brachiopods lespecialfy Orthotetes sp.) on bedding plane surfaces. Eliert, light- to medium-gray, commonly grey mottled, very fine grained, medium beds to
I J [I r [I 1 '1 thin elongated lenses, wavy bedding plane surfaces. (in upland surfaces unit weathers to angular residual slabs and blocks with Iimonite-stained irregular
m E I ' r‘ r ‘ 1 surfaces. Well exposed in bluffs along Ohio River.
2 ﬂ 'f ' * 20-35
g . limestone, light- to medium»gray, medium-crystalline to oolitic, thick to medium-bedded, few thin beds. Contact between Ste. Genevieve and St. Louis Limestones
n' s has been defined by other workers as approximately at the bottom of this unit 20 to 35 feet below the base of the Lost Rivarl‘!) Chart (Ray and others, 1946;
_ E Malott, 1952i
a: E c: I I I r
2; g L J_
m 2
E
_ g
: Limestone, dolomite, and chert, Limestone, brownish- to medium-gray, medium- to finely crystalline, commonly dolomitic in upper half; lower half commonly
w "’ 130. argilleceous; scattered beds have light-gray to light-brown chert nodules; thick to medium bedded, few thin beds. Dolomite, medium- to dark-gray, finely
zoo crystalline, thick- to mediumvbedded, few beds argillaceous; in upper part. Chart, light-gray to light-tannish-gray, very fine grained. medium-bedded, wavy
m bedding plane surfaces; grades laterally to lenses; at top of unit,
5
TI l‘K | r l‘
t t t
l I l I
i‘r'l'i 340-375
‘7 —l _
I,' i Limestone end gypsumlll: Limestone, medium— to dark- to brownish-gray, fine-grained to sublitliographic, tliick- to medium-bedded; some thin to shaly bedded to
T‘Z;T:_‘ laminated argilleceous limestone near middle; in part dolomitic; “Lithostrotion” spp. and Syringopora 5p. present; petroliferous odor common when
I I ‘ If 60-70 broken. Gypsum not present at surface in map area but penetrated by core drill hole in the adjacent Guston quadrangle [Kepferle and Peterson, l962l; large
1T1 ' J ; springs in bluff about 0.8 mile east of Brandenburg (Gallaher, 1964) are at the same stratigraphic positions as bedded gypsum encountered in the drill
I T I I. hole.
‘ . 1 . ‘l’
T I l x1 I l
L i .
II J l‘ l
f 1' l i
J 1
I; r "
I l
r; f ‘
r
r—Tt
1 E] r
.V L _ Limestone, medium- to dark- to brownish-gray; some light olive gray in lower part; fine grained to sublithographic; few beds in lower half medium grained organic
J'flégﬁ: 55.30 detrital; thick to medium bedded, few thin to laminated beds near middle; dolomitic at base and near top; commonly argilleceous near middle, scattered gray
cliert nodules about middle and near base; few small spiriferid brachiopods; strong petroliferous odor when broken,
Salem Limestone, light-brownish-gray to brownish-gray, medium- to coarse-grained, medium- to thick-bedded, slightly crossbedded; some dolomitic. Dnly upper few
I' 10+ feet exposed in extreme southeast corner of map area.

 

 

 

 

 

 

FIGURE 35.—Stratigraphic section and description showing lithologies of St. Louis and Ste. Genevieve Limestones and Lost River

Chert Bed in New Amsterdam quadrangle, west-central Kentucky (from Amos, 1972).

63

64 MISSISSIPPIAN ROCKS IN KENTUCKY

 

FIGURE 36.—Lost River Chert Bed (above hammer handle) about 2.5 ft thick. Roadcut on north side of Kentucky Highway 96 at Touristville,
Mill Springs quadrangle, Wayne County.

The distinctive colonial coral Schoenophyllum aggre-
gatum Simpson 1900 (Lithostrotion (Siphonodendron)
genevievensis and Lithostrotion harmodites of earlier
reports) has long been recognized as a useful guide fossil
for the Ste. Genevieve (for example, Ulrich and Smith,
1905; Ulrich, 1917; Butts, 1917, 1922). The coral has a
narrow stratigraphic but widespread geographic range
across the Ste. Genevieve outcrop belt west of the Cin-
cinnati arch (Dever and others, 1980). Its occurrence is
restricted to a zone 1—10 ft thick (more commonly, 1-2
ft thick) in the upper part of the formation. The zone
commonly is about 20 ft below the top of the Ste.
Genevieve in northern west-central, central, and south-
central Kentucky, and about 40—50 ft below the top in
southern west-central and western Kentucky. The coral
also occurs in the Ste. Genevieve in parts of south-
central Kentucky (Butts, 1922; Lewis, 1971b) where it
is locally abundant, but generally rare (for example,
Lewis and Taylor, 1976; Taylor, 1977). Some zonation
in this area was also noted by Lewis (1971b), but the

 

stratigraphic range of the coral within the unit has not
been well defined.

The top of the Ste. Genevieve in the outcrop west of
the Cincinnati arch is marked by a widespread zone of
altered limestone, 1—5 ft thick, which has been cor-
related with the Bryantsville Breccia Bed of Indiana
(Patton, 1949; Malott, 1952). Diagenetic features in the
limestone indicate that alteration resulted from
subaerial exposure and vadose diagenesis. The exposure
zone, mostly darker colored than the underlying lime-
stone, has a variety of lithologic expressions in the out-
crop belt, displaying various combinations of carbonate
lithotypes and diagenetic fabrics, ranging from rubbly
to massive, highly brecciated calcilutite and calcisiltite
to unbrecciated calcarenite containing micritic crusts
and stringers. Cherty silicification of the diagenetic
crusts and stringers is common.

The Ste. Genevieve crops out east of the Cincinnati
arch in the northeast-southwestward-trending belt of
Mississippian rocks along the western border of the

ROCKS 0F LATE OSAGEAN AND MERAMECIAN AGE 65

 

FIGURE 37.—Ste. Genevieve Limestone Member of Monteagle Limestone (Mmsg) and St. Louis Limestone (Msl). Arrows indicate contact
between members at top of l-ft cherty limestone bed. Kentucky Highway 80, Pulaski County.

Eastern Kentucky coal field; it is absent along the axis
of the arch. During the geologic mapping program, the
Ste. Genevieve in this eastern outcrop belt was reduced
in stratigraphic rank and mapped in south-central Ken-
tucky as the Ste. Genevieve Limestone Member of the
Monteagle Limestone (Lewis and Thaden, 1965; Lewis,
1971b); in northeastern, east-central, and northeast-
south-central Kentucky, it was mapped as the Ste.
Genevieve Limestone Member of the Newman Lime-
stone (Hatch, 1964; Cohee and West, 1965). In 1984, the
Ste. Genevieve in east-central and northeastern Ken-
tucky was assigned to the Slade Formation (Ettensohn
and others, 1984) after a more comprehensive under-
standing of the stratigraphic relations of Upper Missis-
sippian strata was reached.

Lithologies in the Ste. Genevieve are virtually the
same on both sides of the Cincinnati arch; some included
stratigraphic markers occurring on the west side have
been recognized east of the arch. The Lost River Chert
Bed is present 15—20 ft above the base on the east in
parts of south-central Kentucky, but it commonly is un-
fossiliferous (McGrain, 1969). Northeast of Somerset,
a bed of cherty limestone is at the top of the mapped
St. Louis; it weathers to a chert similar in appearance
to the Lost River Chert.

 

A zone of altered limestone, with features indicative
of subaerial exposure and vadose diagenesis, caps the
Ste. Genevieve east of the Cincinnati arch and has been
correlated with the Bryantsville Breccia Bed at the top
of the unit west of the arch (McFarlan and Walker,
1956). In south-central Kentucky as in central Ken-
tucky, the breccia bed is bracketed by diagnostic crinoid
fauna, Merarnecian Platycrinites penicillus in the Ste.
Genevieve, and Chesterian genus Talarocn‘nus. In
south-central Kentucky, the unit containing Talaro-
cn‘nus is named the Kidder Limestone Member of the
Monteagle Limestone (Lewis, 1971b). Northeastward
along the outcrop belt, identiﬁcation of the Bryantsville
Breccia Bed and top of the Ste. Genevieve is com-
plicated by (1) the presence of multiple exposure zones,
lithologically similar to the Bryantsville, in both the Ste.
Genevieve and overlying limestones, (2) absence or
sparse presence of diagnostic crinoid fauna, and (3)
northeastward thinning of the carbonate units (Dever,
Hester, and others, 1979).

Rocks in east-central and northeastern Kentucky
assigned to the Ste. Genevieve Limestone by Butts
(1922) and McFarlan and Walker (1956), and designated
as the Ste. Genevieve Limestone Member of the
Newman Limestone during the mapping program, for

66 MISSISSIPPIAN ROCKS IN KENTUCKY

example in the Tygarts Valley (Sheppard, 1964) and
Bangor (Hylbert and Philley, 1971) quadrangles, con-
sist of two distinct limestone units, separated by an ero-
sional unconformity (Dever, 1973, 1980b). The older
Platycn'nites-bearing Ste. Genevieve, composed main-
ly of bioclastic and oolitic calcarenite and capped by the
Bryantsville Breccia Bed, extends along the outcrop
belt from south-central into northern east-central Ken-
tucky (fig. 59) where the unit was erosionally truncated
following Late Mississippian movement along the Ken-
tucky River fault system. In addition to being truncated
on the north, the Ste. Genevieve in the east-central area
pinches out eastward toward the axis of the Waverly
arch, which remained a positive feature following the
earlier uplift that interrupted St. Louis deposition.

The younger limestone unit included in the Ste.
Genevieve was deposited after the period of uplift and
erosion associated with recurrent movement along the
Kentucky River fault system. Extensive erosion, mainly
on the upthrown (north) side of the fault system,
resulted in erosional thinning and truncation of the
Platycn'nites-bearing Ste. Genevieve and erosional thin-
ning and removal of the St. Louis and upper Borden
(Renfro, Nada, and upper Cowbell) in parts of northeast-
ern and east-central Kentucky. The younger limestone,
named the Warix Run Limestone Member of the N ew-
man (Dever, 1977) and later of the Slade Formation
(Ettensohn and others, 1984) was deposited on the ero-
sional surface developed on the uplifted northern block,
where it rests unconformably upon the Borden (mainly
on the Cowbell Member) and only locally upon the older
Ste. Genevieve and St. Louis. South of the fault system,
it rests unconformably upon the Ste. Genevieve and
St. Louis.

The Warix Run Member, as much as 100 ft thick, con-
sists mainly of crossbedded quartzose calcarenite, with
lesser amounts of calcilutite. The calcarenite is com-
posed of peloids, with sparse to locally abundant
micrite-enveloped grains, ooliths, and bioclastic grains.
The unit contains abundant quartz silt and sand, and,
locally, grades into calcareous sandstone (Klekamp,
1971). Material derived from eroded members of the
Borden and Newman is present in the basal part; grains
and granules of St. Louis chert and limestone are com-
mon constituents in much of the unit.

Rocks assigned to the Warix Run Member of the
Slade Formation accumulated in erosional lows on the
uplifted northern block of the Kentucky River fault
system, partly filling them. The member commonly
reaches its maximum thickness, as much as 100 ft
(Englund, 1976), near the middle of these lows, and it
thins and pinches out along the margins of each area.
Thus, the member forms a series of isolated deposits
in the present outcrop of northeastern and northern

 

east-central Kentucky (Dever, 1977; Ettensohn and
others, 1984). South of the Kentucky River fault system
where the unit is more widespread, it is thinner; here
it forms a blanketer deposit of calcarenite which has
been traced as far south as central Rockcastle County
(fig. 59).

In northeastern and northern east-central Kentucky,
Warix Run calcarenite is overlain by calcilutite of the
upper member of the Newman or Slade. The lithologies
of these two units form a transgressive-regressive se-
quence; the contact between them appears conformable,
in part sharp and in part intertonguing. Diagenetic
features indicative of exposure and vadose diagenesis
occur at the top of the Warix Run in parts of the area,
but their meager development and the absence of ex-
tensive alteration in the limestone suggest formation
during relatively brief periods of exposure.

On the Pine Mountain overthrust block in southeast-
ern Kentucky, an interval in the lower Newman Lime-
stone, composed mainly of oolitic and bioclastic
calcarenite, was identified as Ste. Genevieve by Butts
(1922) and Hauser and others (1957). The calcarenitic
interval, as much as 100 ft thick, overlies the zone of
cherty limestone and dolomite in the basal Newman
which was identified as St. Louis. Above the calcarenitic
interval are beds correlated by de Witt and McGrew
(1979) with the Taggard Red Member of the Greenbrier
Limestone of Wilpolt and Marden (1959), a West
Virginia unit characterized by the presence of
variegated limestone, mudstone, and shale. Butts (1922)
reported the occurrence of Lithostron'on (Siphonoden-
dron) genevievensis, now Schoenophyllum aggregatum
Simpson 1900, in the Ste. Genevieve at Pineville.
Platycrinites penicillus has been found in the interval
below the Taggard at localities in Letcher and Whitley
Counties, Ky. In West Virginia, the Taggard has been
placed at the Meramecian-Chesterian boundary (Wells,
1950) and in the basal Chesterian (de Witt and McGrew,
1979).

STRATIGRAPHIC RELATIONSHIPS

ST. LOUIS AND OLDER UNIT RELATIONSHIPS

West of the Cincinnati arch in Kentucky, field rela-
tionships between the St. Louis Limestone and the
underlying Salem Limestone or the Salem and Warsaw
Formations unit appear conformable, but lithologic
units at the mapped boundary vary (Withington and
Sable, 1969). On a 71/2-minute quadrangle scale, bound-
ary units commonly interlayer strata characteristic of
each unit, or else the lithologic contact can be sharp and
unequivocal. In larger areas, the relationships are
conformable and probably gradational or intertonguing,

ROCKS OF LATE OSAGEAN AND MERAMECIAN AGE 67

so that inconsistencies arise when regional correlations
are attempted.

Lineback (1972), in an attempt to reconcile differences
in the Salem and St. Louis boundary, traced contacts
in the subsurface eastward from the St. Louis, Mo., type
area of the St. Louis and westward from the Salem, Ind.,
type area of the Salem into areas of Illinois, and con-
cluded that the two formations partly intergrade and
reciprocally thicken and thin in Madison and Bond
Counties, Ill. Lineback virtually restricted the Salem
to strata containing biocalcarenitic and oolitic limestone
(“biocalcarenite facies”). His criteria for this usage are
similar to those used in western Kentucky. Both the
older (Baxter, 1960; Baxter and others, 1963, 1967) and
newer (Lineback, 1972) usages are shown on the cross
section (fig. 9).

Contact criteria between the Salem and St. Louis
Limestones in central Kentucky generally follow those
used in adjacent Indiana, but according to Pohl (1970,
p. 3), the criteria are inconsistent. The contact in central
Kentucky is placed between underlying biocalcarenite,
argillaceous shaly limestone, or saccharoidal dolomitic
limestone and overlying even-bedded cherty, dense,
micrograined or micritic, in part laminated limestone
and dolomitic limestone. The contact is based on gross
lithologic differences, but it is generally sharp to grada-
tional over a few feet (or at the most, 10—20 ft) in cen-
tral, west-central, and south-central Kentucky, where
the combined thickness of the St. Louis and Salem
ranges from about 100 to 400 ft. In western Kentucky,
however, the combined thickness increases from about
400 to more than 700 ft; here a unit of dark, partly
argillaceous limestone as much as 170 ft thick inter-
venes between biocalcarenite beds typical of the Salem,
and the dense micritic limestones considered typical of
the St. Louis (Trace, 1974). In some quadrangles the
Salem and St. Louis have been mapped as an undivided
unit (Fox, 1965; Rogers, 1963); in more recent maps the
units were differentiated (Trace, 1974; Rogers and
Hays, 1967 ). In western Kentucky, the uppermost
boundary of the Salem is placed at the top of the highest
biocalcarenite or oolitic limestone bed. This corresponds
to the boundary of Lineback (1972) in Illinois. Elsewhere
in Kentucky, the current lithologic boundaries of the
Salem also appear to generally correspond to those in
Illinois.

To early workers concerned with the relative age of
the St. Louis, faunal and lithologic differences indicated
that a hiatus or disconformity separated the St. Louis
from older rocks in the Mississippi River Valley
(S. Weller, 1909; Van Tuyl, 1925). More recent studies
indicate that conodont faunas are transitional upwards
into basal St. Louis beds, but that a sharp widespread
break in the conodont succession exists within the lower

 

part of the St. Louis and is associated with limestone
breccias in some places (Collinson and others, 1971,
p. 382). This break, between the Taphrognathus
van'ans-Apatognathus zone and the Apatognathus
scalenus—Cavusgnathus zone, may mark a widespread
hiatus. Similar positions of this zonal boundary are
reported by N icoll and Rexroad (1975) in Indiana adj a-
cent to north-central Kentucky, and in general the break
may correspond with occurrence of evaporite beds and
breccia beds in Illinois, which have been considered to
represent remnants of fractured carbonate beds asso-
ciated with dissolved evaporite strata (Collinson, 1964,
p. 9).

The Renfro Member of the Borden and Slade Forma-
tions apparently spans the Osagean-Meramecian bound-
ary; its conformable stratigraphic relationships with
underlying Borden clastic units and with the overlying
St. Louis have previously been discussed. In some
areas, however, glauconite in and underlying the basal
Renfro suggests a depositional hiatus there.

Reciprocal thicknesses of the Renfro Member and the
overlying St. Louis Limestone Member of the Slade For-
mation in east-central Kentucky are indicated by Weir
and others (1966, p. F20), although the Renfro includes
equivalents of Muldraugh, Salem and Warsaw Forma-
tions unit, and about the lower two-thirds of the St.
Louis. Thickness reciprocity is inferred by probable in-
tertonguing of the Salem and St. Louis in areas west
of the Cincinnati arch. Thus, the thicknesses of the com-
bined Warsaw-Salem-Renfro and St. Louis thicknesses
(fig. 57) are considered to represent a more generally
correct gross chronostratigraphic unit than do isopach
maps of the separate units.

ST. LOUIS—STE. GENEVIEVE RELATIONSHIPS

The contact between the St. Louis and Ste. Genevieve
Limestones and their equivalents in Kentucky east of
the Cincinnati arch is based on general lithologic and
faunal differences. It has been considered to be regional-
ly inconsistent in Illinois (Swarm, 1963, p. 27). Likewise,
in western Kentucky, some abrupt vertical offsets of
this contact (fig. 34) have resulted from differences in
contact criteria used by geologists who mapped adj a-
cent quadrangles (Pohl, 1970, p. 6), considerations of dif-
ferent criteria used by adjoining States, the lack of
detailed faunal studies, and the characteristically poor
exposures of rocks associated with this contact.

The Lost River Chert Bed of Elrod, discussed pre-
viously, has been used in west-central, central, and
south-central Kentucky (McGrain, 1969) to approximate
the St. Louis—Ste. Genevieve contact, which occurs in
areas of characteristically poor exposures. This zone or
its identical lithologies have been reported in many

68 MISSISSIPPIAN ROCKS IN KENTUCKY

quadrangle section descriptions from south-central to
western Kentucky. Identical lithologies have been
observed at least as far west as near Anna, southern
Illinois, by Sable, and may extend into Missouri and to
northwestern Illinois (F.J. Woodson, written commun.,
1983). In western and west-central Kentucky these
rocks have been mapped as part of the St. Louis
Limestone (Rogers and Hays, 1967). The Lost River
Chert Bed is considered by Sable to mark a widespread
virtually time-equivalent unit which lent itself by its
porosity and permeability to relative ease of
silicification.

Criteria for discerning the base of the Ste. Genevieve
Limestone vary considerably throughout Kentucky, as
they have in adjoining States. In Kentucky, the basal
contact has been mapped at the following successively
lower horizons and approximations of horizons:

1. The base of abundant oolitic limestone beds; this
contact corresponds to the base of the Fredonia Lime-
stone Member, and approximates the contact as
mapped in Illinois. It has been used commonly in west-
ern Kentucky, particularly in the Fluorspar district
(Trace and Amos, 1984). There it coincides with the base
of a laminated dolomitic limestone and the top of a zone
of nodular light-gray chert, the “upper” or “first”
cherty zone of mapping geologists in western Kentucky
(Trace, 1962). This is the contact currently recognized
in that area.

2. The base of all oolitic or biocalcarenitic limestone
beds in the St. Louis—Ste. Genevieve interval, at the top
of a “lower” or “second” cherty zone. This contact was
used in earlier mapping of the western Kentucky and
Illinois ﬂuorspar district (Weller and Sutton, 1951), and
has generally been considered to be the boundary be-
tween the two formations in Indiana (Malott, 1932;
McGrain, 1943) and in central to west-central Kentucky.

3. The uppermost occurrence of “lithostrotionoid”
corals which are commonly silicified and are con-
spicuous in deeply weathered areas. This involves only
minor lithologic differences but has been used in both
west-central and western south-central Kentucky
(Haynes, 1966).

4. At or as much as 20—50 ft below the Lost River
Chert Bed (Kepferle, 1963b; Taylor, 1976). The upper-
most “lithostrotionoid” corals occur within 10—25 ft or
less below the Lost River Chert Bed in south-central
and southern east—central Kentucky, but the interval
between the corals and the chert bed increases west-
ward and may be as much as 110 ft below the top of
the Lost River Chert lithologic equivalent in western
Kentucky. As previously mentioned, the lithologic
equivalents of the Lost River Chert Bed have been
observed in several western Kentucky quadrangles (for
example, Rogers and Hays, 1967). A possible candidate

 

for the Lost River Chert Bed equivalent in western and
west-central Kentucky is a chalky-weathering chert con-
taining abundant fenestellid bryozoans and brachiopods
below an interval of abundant nodular chert. This would
correspond to the lower, or second, chert of the upper
member of the St. Louis Limestone (Klemic, 1966a,
1966b; Klemic and Ulrich, 1967; Ulrich, 1966; Ulrich and
Klemic, 1966). Geologists mapping in that area, how-
ever, interpreted this lithology either to occur
sporadically throughout a thick stratigraphic interval
of as much as 250 ft in the upper member of the St.
Louis, or to represent secondary vadose or other ex-
posure silicification after deposition in keeping with a
new tectonic regime, possibly post-Paleozoic or post-
Mesozoic. They therefore did not relate this lithology
to a specific stratigraphic position. Nevertheless, the
unique abundance and association of fossil fauna are
very distinctive. The chert may be a diagenetic feature
or a replacement product of a widespread subtidal
deposit which preceded minor uplift or sea-level with-
drawal enough to expose the beds to supratidal condi-
tions, initiating silicification by vadose circulation. It
is quite certain that only one such zone containing one
or more cherty beds exists in Indiana and west-central,
central, and south-central Kentucky. It is difficult to
explain more than one such zone in western Kentucky,
which lies basinward of the above area, and where
repeated exposure seems more unlikely than in Shelf
areas.

In addition to the mapped contacts described above,
other St. Louis—Ste. Genevieve boundary positions have
been proposed. Pohl (1970) placed the boundary several
feet above the Lost River Chert Bed and proposed a new
subdivision, the Horse Cave Member of the St. Louis
Limestone, to include the transitional lithologies be-
tween the lowest prominent oolitic limestone in the St.
Genevieve and the highest occurrence of spheroidal
(“ball”) chert and amoeboid to lensatic chert in the St.
Louis dolomitic micritic lithologies. The approximate
upper limit of the Horse Cave, according to Pohl, marks
a “viable extinction” of a foraminifer genus and a
dasyclad algal genus. Macrofossils historically con-
sidered to be the guide fossils of the Ste. Genevieve
(Platycn'nites penicillus [Meek and Worthen] and
Pugnoides ottumwa [White]), however, extend as much
as 65 ft below the Lost River Chert Bed in western
south-central Kentucky according to Pohl. An even
higher St. Louis—Ste. Genevieve boundary in southern
Illinois, the base of the Spar Mountain Sandstone
Member of the Ste. Genevieve, is based on conodont
zonation (Rexroad and Collinson, 1963; Burger, Rex-
road, and others, 1966). In south-central Kentucky, con-
odont studies suggest to Rexroad that the upper limit
of the St. Louis lies roughly 2 to 10 ft or more above

ROCKS OF LATE OSAGEAN AND MERAMECIAN AGE 69

the Lost River Chert Bed (F.J. Woodson, written
commun., January 1983). These contacts, although they
combine both lithologic and faunal criteria, do not lend
themselves to mapping definition in that the lithologic
criteria are subtle, rarely observable, and extremely dif-
ficult to map in the poorly exposed areas in which they
occur.

As mapped, the relationships between the St. Louis
and Ste. Genevieve west of the Cincinnati arch are ap-
parently conformable and gradational. In south-central
Kentucky, east of the arch, the contact between the
cherty limestone and dolomite of the upper St. Louis
and the basal Ste. Genevieve calcarenite also appears
conformable. However, in the northeastern part of
south-central Kentucky, a conglomerate as much as 10
ft thick, composed of abundant clasts of St. Louis-type
chert in an oolitic matrix, has been reported at the base
of the Ste. Genevieve (Butts, 1922; Lewis, 1971b). In
parts of the area around Somerset, roadcut exposures
show that from 4 to 18 ft of Ste. Genevieve-type cal-
carenite is present between the base of the chert-bearing
conglomerate and the top of the cherty calcilutite and
calcisiltite of the St. Louis. Across most of east-central
Kentucky, the Ste. Genevieve rests unconformably
upon an exposure zone at the top of the St. Louis and
commonly contains a thin, basal conglomerate of St.
Louis chert and limestone clasts. The St. Louis was
removed from parts of northeastern Kentucky by intra-
Mississippian erosion; where present, it is unconform-
ably overlain by limestone and shale of the upper part
of the Slade Formation (formerly Newman Limestone),
which is considered to be a Chesterian equivalent.

SOURCES OF SEDIMENTS AND
DEPOSITIONAL ENVIRONMENTS

Although all preserved rocks of Meramecian age were
deposited in marine or marginal marine environments,
the distribution of terrigenous detrital material in-
dicates transport from source areas mostly east and
northeast of the Appalachian basin. Lesser sources of
sediment may have been east of the present-day south-
ern Appalachians, and small amounts of sediment may
have been shed from areas of low relief along the Cin-
cinnati arch and Ozark uplift. Williams (1957,
p. 315—316) suggested three source areas for detritals
in Warsaw and Salem rocks for the midcontinent
region—Wisconsin, Ozarkia, and Appalachia. Rubey
(1952, p. 50) reported sand grains derived from igneous
and metamorphic rocks in the St. Louis of eastern Mis-
souri, and suggested the presence of newly exposed land
areas in the Ozark region.

Directional data for the Garret Mill Sandstone
Member are reported by Lewis and Taylor (1979). They

 

interpreted the Science Hill Member of the Warsaw to
be a shallow deltaic sand body derived from an eastern
source area. Thus the Science Hill, at least, appears to
represent a late southwestward spillover along the
southern part of the Borden delta front after clastic
deposition had generally ceased in more northern areas
such as in central Indiana.

Source areas and transport directions are not known
for the conspicuous mudstones of the Salem and
Warsaw Formations unit (“Somerset Shale”) in the
Cumberland saddle area of south-central Kentucky and
adjacent Tennessee and those of the Salem in central
Kentucky. These mudstones are interpreted to be
genetically related to the sandstones in those units.
Their general distribution and thickness suggest west-
ward and northward transport. They, along with the
sandstones of the Garret Mill and Science Hill
Members, may have been swept westward from sources
east of the southern Appalachians possibly through
the Cumberland saddle, with the fine .clastic muds
carried farther westward and northward along the west
ﬂank of the Cincinnati arch. These latter two possi-
bilities seem likely because eroded pre-Mississippian
units on the Jessamine or Nashville domes probably did
not contain sufficient sand to be the source for the
sandstones.

A northeastward-increasing sand content and thick-
ening of the Warix Run Member of the Slade (Newman),
in northeastern Kentucky and correlatives in southern
Ohio, northern West Virginia, and western Pennsylvania
suggest a northeastern source area for that member (de
Witt and McGrew, 1979). Southwestward- and westward-
dipping regional paleoslopes are indicated during
Borden and, presumably, ensuing Mississippian deposi-
tion (Potter and Pryor, 1961). In contrast, sediment
transport to the northeast is indicated by paleocurrent
measurements in the calcarenites of the Warix Run in
western Carter County (Klekamp, 1971). The calcar-
enites, interpreted as tidal-channel deposits, were trans-
ported in migrating large-scale sand waves. Klekamp
suggested that northeastward transport in the area did
not necessarily reﬂect regional paleocurrent directions,
but may have resulted from coastline configuration,
wind direction, or ﬂood- and ebb-dominated tidal cur-
rents. The Warix Run of northeastern Kentucky was
deposited on the irregular topography of an intra-
Mississippian erosional surface cut down as deeply as
into the Borden. In western Carter County, the Warix
Run also was studied by Ferm and others (1971) and
Home and others (1974), who considered it to represent
tidal-bar belt deposits and, around topographic highs,
a complex of beach, bar, and tidal-channel deposits.

The Spar Mountain Sandstone Member (Tippie, 1945)
of the Ste. Genevieve Limestone in Illinois and the

70 MISSISSIPPIAN ROCKS IN KENTUCKY

younger Rosiclare (Aux Vases Sandstone) in Illinois and
western Kentucky are two detrital tongues in the up-
permost part of the otherwise carbonate-dominated
Meramecian sequence. Sand in both units was con-
sidered by Swann (1963) to have been transported from
northeastern source areas by a southwest-ﬂowing river
system, the Michigan river, which was active also dur-
ing Chesterian time. N 0 studies of current directional
criteria in these units are known to have been made in
Kentucky.

Depositional environments for Meramecian-age
sediments were largely very shallow marine except for
the darker and finer grained rocks of the Salem and St.
Louis in westernmost Kentucky and southernmost Il-
linois, where the water may have been moderately deep
(more than 500 ft?) during Salem and early St. Louis
time. The textures of the shallow-marine carbonate
rocks reﬂect a wide variety of depositional conditions
including oolite shoals, lagoons, reefoid or bank detritus
areas, and bars. Waters ranged from clear to turbid, and
from agitated to quiet. High-energy environments pro-
duced by wind-driven and (or) tidal currents combined
with shallow water (less than 100 ft) over most of Ken-
tucky are indicated by fossil fragmental, crossbedded,
and interlensing beds in the Harrodsburg, Salem and
Warsaw Formations unit, Ste. Genevieve, and Rosiclare
(Aux Vases) strata, whereas quiet water of probable bay
and lagoon intertidal-supratidal environments is
inferred for the St. Louis evaporites and other fine-
grained carbonate rocks.

The Harrodsburg’s relatively pure carbonate com-
position, abundant disarticulated fossil remains, and
crossbedding indicate a striking change from turbid
water, in which the underlying Borden was deposited,
to widespread clear water. Fossil-debris beds,
biorudites, that fill depressions marginal to Borden
delta platforms (fig. 17) probably are the disarticulated
hard parts of fauna indigenous to the platforms, swept
by currents into adjacent depressions on the sea ﬂoor
(Peterson and Kepferle, 1970).

Pinsak (1957) concluded that near-shore, shallow-
water offshore (shelf), and deep-water offshore (basin)
environments are represented in limestones of the
Salem Limestone of northern, central, and southwestern
Indiana. The shallow-water offshore environment of
Pinsak has been compared to the environment on the
present-day Bahama Banks (Sedimentation Seminar,
1966). Clear-water conditions are shown by the Salem
in Indiana, but markedly turbid water is indicated by
the argillaceous content of the Salem and the Salem and
Warsaw Formations unit in Kentucky and by associ-
ated sandstone strata in the Cumberland saddle area.

A distinct change to very low energy conditions is re-
ﬂected by rocks of the St. Louis Limestone. Evaporites

 

in the St. Louis Limestone in Illinois, Indiana, and
western Kentucky are indicators of a probable sabkha
shelf environment along the Cincinnati arch trend
(J orgensen and Carr, 1973). A seaway opening to the
south and intermittently connected with the Michigan
basin, where evaporite deposition also occurred, is in-
ferred to have been the source of waters for the

_ evaporite deposits in the St. Louis (Sable, 1979b). The

widespread limestone and dolomite breccias in the St.
Louis Limestone west, south, and east of known
evaporites are interpreted to result from dissolution of
contiguous evaporite beds which occurred over a much
wider area than present evaporite occurrences (Collin-
son, 1964, p. 7; Dever and others, 1978). Carbonaceous
mudstones and limestones in lower St. Louis beds of
west-central Kentucky (Kepferle, 1966b; Withington
and Sable, 1969), probable time equivalents of the
evaporites, may indicate that land along the Cincinnati
arch acted as a barrier to free circulation.

Rocks in the upper St. Louis largely represent a
general transition to clear, freely circulating but quiet
water without terrigenous clastic inﬂuence, perhaps
restricted by conditions similar to those of the Florida
Bay, in which organically precipitated carbonates from
organisms such as calcareous algae could accumulate.
A slow but major transgression during this time prob-
ably inundated all or much of the Cincinnati and Waver-
ly arches. Fine grain size, tabular beds, and general
scarcity of terrigenous elastic debris indicate that wave
and current activity was weak over a large area and that
source areas were low or remote. Dolomitic beds, the
relative scarcity of fossil remains, and scattered occur-
rences of gypsum also suggest that a hypersaline en-
vironment existed in some areas throughout much of
later St. Louis time.

During deposition of the Ste. Genevieve, clear, very
shallow oxygenated seas with normal open circulation
prevailed. A general vertical transition to shallow agi‘
tated water occurred from St. Louis to Ste. Genevieve
time throughout the basin, as indicated by an upward
increase of oolitic limestones, current structures, and
prolific shallow-water faunas and sedimentary limestone
breccias of probable supratidal and subaerial exposure
origin. Low-energy protected lagoon or bay conditions
in which lime mud was dominant were supplanted by
subtidal shelf and oolitic shoal and bank environments
as general marine transgression occurred. Terrigenous
clastics of the Spar Mountain, Aux Vases, and less
widespread units in the central and southern part of the
Eastern Interior basin are generally sheetlike deposits
in lobes that accumulated in this shallow marine en-
vironment. These clastics and more subtle examples of
cyclicity in carbonate depositional environments, as in
the Ste. Genevieve of south-central Kentucky (Sandberg

MERAMECIAN-CHESTERIAN SERIES BOUNDARY 71

and Bowles, 1965), represent the onset of cyclical deposi-
tion which characterizes the marine and continental
environments displayed by Chesterian deposits still to
come.

Paleogeographic interpretation indicates that during
most of Meramecian time, Kentucky was inundated by
a ﬂuctuating shallow sea. Seaway connections existed
mainly to the south, and connections with the Ap-
palachian basin may have been through the Cumberland
saddle. The Cincinnati and Waverly arches were inter-
mittently emergent. There, in northern east-central and
northeastern Kentucky, after uplift resulted in cessa-
tion of St. Louis deposition, the Waverly arch remained
partly emergent during the remainder of Meramecian
time and early Chesterian time.

Paleogeographic reconstruction for Meramecian time
would show topographically low peninsulas along the
Cincinnati and Waverly arches southward into Ken-
tucky. During an early St. Louis marine regression
(evaporite interval), low-lying land may have extended
northwestward along the Kankakee arch (Sable, 1979b).
In Ste. Genevieve time, source areas south or east of
Kentucky may have shed a small amount of sandy sedi-
ment northward.

Evidence for a Michigan river system (p. 85) lies in
the youngest Meramecian units in western Kentucky.
Prior to these, the river may have had little sediment-
carrying capacity due to widespread aridity, or because
it headed in lowlands. Most likely, the point where the
river debouched into the sea was too far distant for
much sediment to reach the Eastern Interior basin
region. In eastern Kentucky, however, either in late
Meramecian or early Chesterian time, northeast-
trending bars composed of carbonate and silica-sand ad-
mixture were prominent along the Waverly arch.

PALEOTECTONIC IMPLICATIONS

Trends and positions of major tectonic elements
inherited from Osagean time stabilized considerably
during ensuing Meramecian time, and infilling of ir-
regular depositional topography continued. The Eastern
Interior basin in western Kentucky was a major
negative feature, as shown by accumulation of more
than 1,100 ft of sediments in western Kentucky, and
thinning towards such positive features as the Cincin-
nati and Kankakee arches. Maximum subsidence of the
basin in western Kentucky occurred during Salem and
early St. Louis time. This major differential downwarp-
ing ceased in late St. Louis time, and wide areas sub-
sided gently and evenly, the exceptions being eastern,
northeastern, and east-central Kentucky, where the
Waverly arch and Kentucky River fault system were
rejuvenated.

 

The Cincinnati arch and Ozark uplift restricted
seaways during St. Louis evaporite deposition, although
they contributed little detrital sediment.

In northeastern, east-central, and eastern Kentucky,
discontinuous deposits of sediments imply a broad,
unstable platform, the Waverly arch (Woodward, 1961),
between the Cincinnati arch and the Appalachian basin
trough. The platform was modified by a shallow basin
in southeastern Kentucky and emergent areas in north-
eastern Kentucky during part of Meramecian time.

MERAMECIAN-CHESTERIAN SERIES BOUNDARY

The Ste. Genevieve Limestone is the youngest forma-
tion of the Meramecian Series in Kentucky. In adj oin-
ing Illinois, the youngest Meramecian (Valmeyeran)
rocks are assigned to different formations (Swarm,
1963). The top of the Ste. Genevieve in eastern Illinois
is correlative with the top of the Fredonia Limestone
Member of the Ste. Genevieve in western Kentucky.
The Rosiclare Sandstone and Levias Limestone
Members of the Ste. Genevieve in Kentucky correlate
with the Aux Vases Sandstone and Levias Limestone
Member of the Renault Limestone, respectively, in
Illinois.

The Ste. Genevieve is overlain by the Renault Lime-
stone in western Kentucky, the Girkin Formation in
southern west-central and south-central Kentucky, and
the Paoli Limestone in central and northern west-central
Kentucky. In the northern west-central area, a sandy
limestone or calcareous sandstone, mapped in the basal
Paoli, represents the Popcorn Sandstone Bed of Swann
(1963). It was identified as the Aux Vases Sandstone
of Malott (1952) on geologic quadrangle maps of the
area, for example in the New Amsterdam quadrangle
(Amos, 1972). East of the Cincinnati arch, the Ste.
Genevieve Limestone Member of the Monteagle Lime-
stone, as mapped, is overlain by the Kidder Limestone
Member of the Monteagle Limestone in south-central
Kentucky and by the upper member of the Newman
Limestone (Mill Knob through Poppin Rock Members
of the Slade Formation of Ettensohn and others, 1984)
in east-central and northeastern Kentucky.

The boundary between the Chesterian and Merame-
cian Series in Kentucky is marked by both a change in
crinoid fauna and a break in deposition. J.M. Weller and
Sutton (1940) considered the unconformity at the
Chesterian-Meramecian boundary to be the most impor-
tant stratigraphic break within the Mississippian
System in the Eastern Interior basin, citing variations
in thickness and absence of upper members or all of the
Ste. Genevieve within comparatively short distances,
overlap of Chesterian units on older beds, and thickness

72 MISSISSIPPIAN ROCKS IN KENTUCKY

variations and conglomerates in basal Chesterian for-
mations. Swarm (1963), however, considered that no im-
portant “time break” occurred between Ste. Genevieve
and Chesterian deposition.

Late Meramecian rocks are characterized by the com-
mon presence of Platycn'nites penicillus and early
Chesterian rocks by species of Talarocrinus (other than
T. simplex) (Swann, 1963). The abundance of crinoid-
bearing limestones in this part of the Mississippian sec-
tion in much of Kentucky makes the presence of these
forms a useful field criterion. In conjunction with the
crinoid “break,” the Bryantsville Breccia Bed exposure
zone at the top of the Ste. Genevieve has been traced
across western, west-central, central, south-central, and
east-central Kentucky, indicating widespread interrup-
tion of deposition at or near the end of Meramecian
time.

The change in crinoid fauna and the Bryantsville
exposure zone are the most reliable field criteria for iden-
tifying the top of the Ste. Genevieve. However, addi-
tional exposure zones with identical or similar
diagenetic features are present below the Bryantsville
in the Ste. Genevieve and above the Ste. Genevieve in
the Girkin Formation, Kidder Limestone Member of the
Monteagle, and upper member of the Slade (Newman)
Formation (Dever, Hester, and others, 1979; Dever,
McGrain, and Ellsworth, 1979; Dever, McGrain, and
others, 1979). In the outcrop west of the Cincinnati arch
and in south-central Kentucky east of the arch, vertical
distribution of the diagnostic crinoid fauna can be used
to distinguish the Bryantsville from older and younger
exposure zones. Criteria used for identifying the
Bryantsville in east-central Kentucky where diagnostic
fauna are absent or sparse are outlined by Dever,
Hester, and others (1979).

In the subsurface of eastern Kentucky, the top of the
Ste. Genevieve is even more difficult to determine.
There, it is largely delineated on the basis of projected
regional trends and, in a few wells, by the presence of
red shale and limestone. These latter lithologies are con-
sidered correlative with the Taggard Formation (Reger,
1926) of the Appalachian basin, which was placed in the
basal Chesterian by de Witt and McGrew (1979).

The Warix Run Member of the Slade Formation was
formerly identified as Ste. Genevieve (Butts, 1922;
McFarlan and Walker, 1956; Sheppard, 1964; McGrain
and Dever, 1967; Hylbert and Philley, 1971). The Warix
Run, separated from the Platycrinites-bearing Ste.
Genevieve by an erosional unconformity (Dever, 1973,
1980b), was deposited on a post-Ste. Genevieve ero-
sional surface. In northeastern Kentucky and northern
east-central Kentucky, the Warix Run rests unconform-
ably upon the Borden and, locally, St. Louis; and in east-
central Kentucky, from Menifee County southwestward

 

into Rockcastle County, it rests on the Bryantsville
Breccia Bed at the top of the Platycrinites-bearing Ste.
Genevieve.

The question of a Meramecian or Chesterian age for
the Warix Run has not been resolved. Microfaunal
studies have been inconclusive (Pohl and Philley, 1971;
Horowitz and Rexroad, 1972). Warix Run calcarenite
in northeastern and northern east-central Kentucky is
overlain by a calcilutitic unit, forming a transgressive-
regressive, fining-upward sequence. The calcilutite,
which locally intertongues with Warix Run calcarenite,
contains Endothyranella, indicating a late Genevievian
or later age (Pohl and Philley, 1971); it has been cor-
related with Talarocn'nus-bearing limestones to the
southwest and Chesterian equivalents in east-central
and south-central Kentucky (McFarlan and Walker,
1956). Determination of a Chesterian age for the Warix
Run would suggest possible correlation with the Pop-
corn Sandstone Bed (Swann, 1963), the basal Chesterian
unit in Lawrence County, Ind., equivalents of which
may be present in northern west-central Kentucky.

In northeastern and northern east-central Kentucky,
erosional thinning and, in parts of the area, complete
removal of Meramecian rocks followed post-Ste.
Genevieve movement along the Kentucky River fault
system (Dever, 1977, 1980b). The Ste. Genevieve is
preserved on the uplifted side of the fault system only
in limited parts of Bath and Menifee Counties.
Elsewhere, the youngest preserved Meramecian unit is
the St. Louis, which generally is unconformably overlain
by limestone and shale of the upper member of the
Newman (Slade), units considered by McFarlan and
Walker (1956) to be Chesterian equivalents.

ROCKS OF CHESTERIAN AGE

Rocks of Chesterian age (figs. 7, 8) are characterized
by distinct cyclical alternation of detrital and carbonate
strata in western Kentucky, and less obvious, but
recognizable cyclicity in eastern Kentucky. Whether
cyclicity is truly synchronous between western and
eastern Kentucky has not been determined. The strata
are preserved in three separate areas: west of the Cin-
cinnati arch from south-central Kentucky into Indiana,
Illinois and southeastern Missouri; east of the arch in
northeastern, east-central, and south-central Kentucky
and in the eastern Kentucky subsurface; and in the Pine
Mountain and Cumberland Mountain belts of south-
eastern Kentucky. Detailed correlations among the
somewhat different sequences in the three areas are
uncertain. The succession is the Chesterian Series in the
Eastern Interior basin and its age equivalent in the
other areas.

ROCKS OF CHESTERIAN AGE 73

A voluminous literature covers many aspects of
Chesterian Series rocks in the Eastern Interior basin.
The synthesis by Swann (1963) discusses Illinois, west-
ern Kentucky, and Indiana; it reviews nomenclatural
history of units; discusses time and rock stratigraphy,
biostratigraphy, and depositional framework; and it
includes an exhaustive, pertinent bibliography. Petro-
graphic and environmental studies of Chesterian Series
sandstone-dominated units include papers by Potter
and others (1958), and Potter (1962, 1963). Conodont
studies include those from the type Chesterian area in
southwestern Illinois (Rexroad, 1957). Rocks of
Chesterian age in Indiana which are directly pertinent
to Kentucky stratigraphy were discussed by Malott
(1952) and those in western Kentucky by Ulrich (1917),
Butts (1917), and McFarlan and others (1955).
Literature on Chesterian equivalents in eastern Ken-
tucky includes papers by Butts (1922), McFarlan and
Walker (1956), Vail (1959), Weir (1970), Dever and others
(1977), Ettensohn (1980), Ettensohn and Dever (1979),
and Ettensohn and Peppers (1979).

Because of the many units in the Chesterian Series,
the formations are not described individually here and
the reader is referred to the geologic quadrangle series
maps of 1961—1979 and reports such as Sedimentation
Seminar ( 1969); Calvert (1968); Englund and Windolph
(1971); Vincent (1975); McFarlan and others (1955);
McFarlan and Walker (1956); Ferm and others (1971);
Stouder (1938, 1941). Trace and Amos (1984) gave ex-
cellent descriptions and discussions of relationships of
Chesterian units in the Fluorspar district of western
Kentucky.

ROCKS WEST OF THE CINCINNATI ARCH

Rhythmically deposited units of limestone and minor
dolomite, and clay- to sand-size terrigenous clastics
characterize the Chesterian Series in the Eastern In-
terior basin. Subdivision into more than 20 formations
is shown on the correlation chart (fig. 8). Formations
have been defined in outcrops in the Chester district in
southwestern Illinois; in southern Illinois; in western,
west-central, and south-central Kentucky; and in south-
central Indiana. Correlations of nomenclatural units are
generally consistent across State boundaries, although
some units which are not divisible in some areas are
readily divided in others. Physical criteria for some for-
mational boundaries are different in adjoining States,
and variations in terminology and groupings of units
occur from State to State.

Chesterian strata in the Eastern Interior basin com-
prise five formal groups in Illinois and three in Indiana.
Groupings are based on lithologic similarity. Facies
changes were produced by shifting loci of terrigenous

 

clastic accumulations (Swarm, 1963, p. 21—22). Time-
stratigraphic divisions include the post-Genevievan,
successively younger Gasperian, Hombergian, and
Elviran Stages in Illinois (Swarm, 1963, p. 21—23); the
stage boundaries are considered to closely correspond
to rock-stratigraphic boundaries.

CARBONATE-DOMINATED UNITS

Widespread units that are mostly limestone west of
the Cincinnati arch in Kentucky include parts or all of
the Renault Formation, the Paint Creek, Paoli, Beaver
Bend, and Reelsville Limestones. the Girkin Formation,
the Beech Creek and Haney Limestone Members of the
Golconda Formation, and the Glen Dean, Vienna,
Menard, Clore, and Kinkaid Limestones. They range
from a few feet to 200 ft thick and generally consist of
relatively pure limestones with micritic microcrystalline
to sparry calcite matrix. Oolitic and pelletal limestones
occur locally in nearly all Chesterian carbonate-
dominated units. Argillaceous and limonitic material is
locally common. Framework grains identified in the ﬁeld
and under the microscope consist of foraminifers, corals,
bryozoans, brachiopods, echinoderms, mollusks (gastro-
pods and pelecypods), and arthropods (trilobites and
ostracodes), as well as ooids, pellets, detrital quartz, and
carbonate rock fragments. Chert and silicified limestone
are relatively rare, although they locally characterize
some units such as the Haney Member of the Golconda
Formation and the Vienna Limestone—and thus are
valuable mapping criteria. Dolomitic limestone beds are
not abundant, but some form marker “zones” such as
those with calcite concretions in the Beaver Bend
Limestone. Interbedded mudstone is generally more
abundant in the middle and upper Chesterian units such
as the Glen Dean, Menard, and Clore than in the lower
ones. Thin and discontinuous shale beds in the Girkin
Formation of west-central and south-central Kentucky
are distal facies of thicker, dominantly sandstone and
shale formations—Bethel (Mooretown), Sample, Elwren
(Cypress) to the north and west (see Sandberg and
Bowles, 1965). The lower surfaces of limestone units are
commonly more planar than those of detrital units (fig.
38), but some gradational and interfingering relation-
ships between limestone and underlying mudstone and
sandstone, such as the Sample and Reelsville in west-
central Kentucky (Sable, 1964), and local thickening of
limestone units at the expense of underlying clastic
units have been recognized. Several limestone units
within the Leitchfield and Buffalo Wallow Formations
in west-central Kentucky are correlated with the
Vienna, Menard, and Kinkaid of western Kentucky
areas (Stouder, 1938; Clark and Crittenden, 1965;
Gildersleeve, 1971; Johnson, 1977).

74 MISSISSIPPIAN ROCKS IN KENTUCKY

 

FIGURE 38.-Upper part of Mooretown Formation (Mm) underlying 18.5 ft of Beaver Bend Limestone (Mbb) overlain in turn by shale and
sandstone of the Sample Sandstone (Ms). Kentucky Stone Company Upton Quarry, Upton quadrangle, Hardin County.

Fairly persistent detrital units within carbonate rock
units are reported in the Paoli Limestone in central and
west-central Kentucky (Kepferle, 1963a), where prom-
inently crossbedded quartzose calcarenite and oolitic
limestone occur, and in the Renault Formation of west-
ern Kentucky, where green shale occurs (Amos, 1965;
Sample, 1965). The Paoli elastic beds may correspond
to the middle shale break of the Paoli in Indiana (Perry
and Smith, 1958). The Renault clastic beds are equiva-
lent to the Yankeetown Shale in southern Illinois
(Baxter and others, 1963, 1967).

Chesterian limestone units west of the Cincinnati arch
which are most widespread and therefore of great value
for long distance correlation are the Beech Creek (sub-
surface “Barlow Lime”), the Glen Dean (“Little Lime”),
the Vienna (“Brown Lime”), the. Menard, and the
Kinkaid.

Chert is particularly useful in correlation of the Haney
Limestone in the outcrop belt of central, west-central,
and south-central Kentucky. This is characteristically
a whitish-weathering replacement of oolitic and fossil
fragmental limestone and is conspicuous in weathering

 

residuum. Similar chert has been observed above the
Hartselle Formation in the Albany quadrangle, in
south-central Kentucky (Preston McGrain, written com-
mun., 1964, verified by Sable); but chert is also common
below the Hartselle equivalent in east-central and north-
eastern Kentucky. Chert also is a valuable widespread
correlation tool in the Vienna Limestone of west-central,
south-central and western Kentucky. This latter is also
a replacement chert but weathers to brown-tinged,
darker hues than does chert in the Haney.

TERRIGENOUS CLASTIC UNITS

Dominantly detrital units in western, west-central
and western central Kentucky include, in general
ascending order, the Bethel Sandstone or Mooretown
Formation, Sample Sandstone, Paint Creek Shale,
Cypress Sandstone or Elwren Formation equivalent
(Malott, 1919), Big Clifty Sandstone Member of the
Golconda Formation, Hardinsburg, Tar Springs,
Waltersburg, Palestine, and Degonia Sandstones. The
Grove Church Shale (Swann, 1963), the youngest known

ROCKS OF CHESTERIAN AGE 75

Chesterian formation of Illinois, is documented only in
the subsurface of Webster County, western Kentucky
(Hansen, 1975). The Leitchfield Formation of west-
central Kentucky includes largely mudstone equivalents
of the Tar Springs and younger elastic and minor
limestone units; the adjacent Buffalo Wallow Forma-
tion is similar lithologically but excludes Tar Springs
beds (fig. 8).

Mudstone and sandstone regionally constitute
roughly 50 and 25 percent respectively of Chesterian
rocks in Kentucky. The proportions and thicknesses of
mudstone and sandstone show great lateral variation.
In many places thick sandstones such as the Big Clifty
(fig. 39) make up most or all of a formation. Their lower
boundaries are commonly erosional, and truncation of
underlying beds can be intraformational or can involve
removal of one or more underlying formations. The
morphology and internal features of the sandstones
established some of them as bar-fingers and channel-
fill or tidal current deposits (Calvert, 1968); some, such

 

 

as the Big Clifty, may be regressive shoreline sheet
sands. Where the basal parts of elastic units are fine-
grained sedimentary rocks, perhaps bay-fill or overbank
deposits, their basal contacts are generally planar and
apparently conformable with underlying carbonate
units (fig. 40). However, splay(?) channel fills of mud-
stone or of mudstone and sandstone admixtures are also
present (Sable and Peterson, 1966, p. 25) although rarely
seen. Marine fossils occur in the lower parts of many
sandstones. Thin coaly beds and unfossiliferous red
shales are locally present in the upper parts of some
units such as the Mooretown and Big Clifty (Sable,
1964; Amos, 1972). Suggesting alternating shoreline
regression and transgression, the depositional regime
was that of lower delta plain and delta front
environments.

Most Chesterian sandstones are dominantly fine- to
coarse-grained orthoquartzites as compared to the
subgraywackes of Kinderhookian and Osagean ages.
Some sandstones in the upper Chesterian strata, such

FIGURE 39.—Big Clifty Sandstone Member of the Golconda Formation showing tabular beds, gently cross-laminated beds. and minor
channel fills. Note asphaltic exudation as drip and ﬂow features. Milepost 120.7. Western Kentucky Parkway, Big Clifty quadrangle,

Hardin County.

76 MISSISSIPPIAN ROCKS IN KENTUCKY

 

FIGURE 40.—Hardinsburg Sandstone (M h) conformably overlying Haney Limestone Member of the Golconda Formation (Mgh). Milepost
114.8, Western Kentucky Parkway, Summit quadrangle, Grayson County.

as in the Pennington Formation of eastern Kentucky,
are subgraywacke types, however. Framework is pre-
dominantly quartz and minor alkali and sodic feldspars,
chert, fossil fragments, rock fragments, and heavy
minerals. Locally, pyrite grains are common. Sideritic
nodules are also locally common. Matrix consists of
common quartz overgrowths, sparry calcite, micrite,
clay minerals (reported as illite and kaolinite in several
reports), and iron oxides. Other constituents are clay
pellets, carbonized plant impressions, Sideritic and
limonitic nodules, carbonized wood fragments, and tree
trunks, with a few in growth positions. Conglomeratic
sandstone is rare; the Mooretown (Bethel) Formation
contains scattered granules and pebbles of white quartz
in a thick channel fill in western central Kentucky, and
the Cypress Sandstone contains sandstone clasts in
western Kentucky.

A striking example of a thick channel-fill deposit is
the Bethel (Mooretown) Sandstone in central, west-
central and western Kentucky. This deposit extends for
more than 150 mi, and the channel has cut through more
than 250 ft of previously beveled pre-Bethel strata as

 

old as the St. Louis (Sable and Peterson, 1966; Reynolds
and Vincent, 1967) (fig. 41). This deposit is interpreted
to have been a submarine shelf channel fill (Sedimenta-
tion Seminar, 1969). Other elongate channel sandstone
bodies have integrated distribution patterns suggesting
delta plain stream channels, some of which are repeated
as stacked channel deposits (Potter, 1962, 1963) in the
north-central part of the western Kentucky ﬂuorspar
district (Trace and Amos, 1984).

Another thick, southerly-trending unit is the Tar
Springs Sandstone in northern west-central Kentucky
(Clark and Crittenden, 1965; fig. 42, this report), con-
sidered by Calvert (1968) to have been a tidal current
sand.

Siltstone, mudstone, and claystone in strata of
Chesterian age are commonly medium dark gray to
greenish gray. Reddish-gray varieties occur in several
clastic units, including the upper part of the Big Clifty
Sandstone Member of the Golconda Formation and
units within the Leitchfield and Buffalo Wallow Forma-
tions; they are common in the Hardinsburg and
Palestine Sandstones of western Kentucky. Dark-gray

ROCKS OF CHESTERIAN AGE 77

Map area , /
l
/
,‘7/

Louisville
r
\ _\ //\ ' I.
a ‘ -' ~, \ xv
; “ -L \\ i "
/" Henderson \"
f Owensboro o
Elizabethtown
HOPKINS
./{ COUNTY

   
 

   

0 10 20 30 “MILES
;|_._l_a

 

 

NW
FEET

SE

 

r—ZOO

Sample Sandstone

 

 

 

 

Paoli Limestone

~100 datum\

 

Beaver Bend Limestone

 
 

Mooretown Formation ' m

 

 

 

 

Ste. Genevieve Limestone

 

 

 

0 1 2
l l l

3 4 5 MILES
| | l

FIGURE 41.—Generalized map showing distribution of Mooretown Formation—Bethe] Sandstone channel fill in central to western Kentucky
and cross section showing relationships of channel fill to other rock units. (Map modified from Reynolds and Vincent. 1967; cross section

northwest of Elizabethtown from Peterson, 1964.)

to black carbonaceous shale with siderite nodules is
associated with thin and discontinuous coal beds in the
upper part of the Big Clifty and the Mooretown (Bethel)
in west-central Kentucky and in the Cypress, Hardins-
burg, and Tar Springs Sandstones of western Kentucky.
Perhaps the thickest is a l-ft bed reported in the Cub
Run quadrangle (Sandberg and Bowles, 1965) of south-
central Kentucky.

ROCKS EAST OF THE CINCINNATI ARCH

East of the Cincinnati arch, Chesterian age equiva-
lents crop out along the western border of the Eastern
Kentucky coal field in a belt extending across south-
central, east-central, and northeastern Kentucky; they
also are exposed along Pine and Cumberland Mountains
in southeastern Kentucky. Carbonate rocks are domi-
nant in the lower part of the succession of Chesterian

 

correlatives; the upper part mainly consists of ter-
rigenous clastic rocks.

During the 1960—1978 geologic mapping in south-
central Kentucky, the interval dominated by carbonate
rocks was assigned, in ascending order, to the Kidder
Limestone Member (Lewis, 1971b) of the Monteagle
Limestone (Vail, 1959; Stearns, 1963), Hartselle Forma-
tion (Smith, 1894) (a relatively thin unit of sandstone
and shale), and Bangor Limestone (Smith, 1890). In
east-central, eastern, and northeastern Kentucky, cor-
relatives of the Kidder, Hartselle, and Bangor general-
ly were mapped together as the upper member of the
Newman Limestone (Campbell, 1893). The overlying
interval dominated by terrigenous clastics was desig-
nated as the Pennington Formation (Campbell, 1893),
but in northeastern Kentucky, following the usage of
Englund and Windolph (1971), these rocks generally
were mapped in the Newman Limestone. The Carter

78 MISSISSIPPIAN ROCKS IN KENTUCKY

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

Buffalo
Buffalo _ Wallow
WaIIovv Formation
FormatIon
Tar Tar
Springs _ Springs
Formation '. '_ Formation
u ‘5 a
Q) .0 I— a.)
a; s2
2) E 3 E
'1’ a:
C C
0 o
‘0‘: :7;
Q’ a)
E E
—‘ _I
C :
‘5 I m
G) I I (D
0 o
I: I I C
2 .“3
0 _ | l | I I I I | I I I I I I I _ ~ 0
E" 8 I I l | I I | I I I | I | I I a) 3
3 E | I I I I | I I I I I I I I I I 3 E
o a, I I 1 I I I I I I I I I I I 3 a;
-' E I I I I I l I I I I I I I I l I I I I ﬂI I_ E
__ ‘ I I I I I I I I I DII—I _
—‘ —‘ | I I I I
_ I | I I I I I I I | I
I l I I l I I l I l I l | l l

 

 

 

 

 

 

FIGURE 42.—Diagram showing relationships of the Glen Dean Limestone. Tar Springs Formation, and lower part of Buffalo
Wallow Formation in Mattingly quadrangle, west-central Kentucky. A, local section at Bull Creek; B, typical of remainder

of quadrangle (from Clark and Crittenden, 1965).

Caves Sandstone (Englund and Windolph, 1971) is a
linear body of sandstone occurring within the
terrigenous-clastic sequence of northeastern and north-
west eastern Kentucky. It commonly rests disconform—
ably upon limestones of the Newman.

In east-central and northeastern Kentucky, as previ-
ously mentioned, the names Slade and Paragon Forma-
tions have replaced the Newman Limestone and
Pennington Formation used during the mapping pro-
gram (Ettensohn and others, 1984). The Slade Forma-
tion (fig. 43) includes 9, possibly 10 (Warix Run)
members of Chesterian age, including the unit former-
ly identiﬁed with the Glen Dean and Bangor Limestones
to the west and south, and the Cave Branch Bed. The
overlying Paragon Formation, mostly fine grained

 

clastic rocks (fig. 44), is divided into four informal
members of contrasting shale and carbonate rock.
Because the gross units, Slade and Paragon, closely ap-
proximate the stratigraphic limits of the previously
used Newman and Pennington terminology, which is
well entrenched in the literature, these terms are used
interchangeably in the discussion here, with Newman
and Pennington generally ﬁrst mentioned because they
are the terms used on the geologic maps. Changes from
the older names are based on a rationale which considers
the older Appalachian basin terms inappropriate
because of correlation difficulties. The various members
of the Slade and Paragon are significant subdivisions
of these units in that they allow study of detailed
depositional environments and authigenic processes.

ROCKS OF CHESTERIAN AGE 79

 

FIGURE 43.—Roadcut mostly in Slade (Newman) Formation. Bench above truck is top of Mill Knob Member. Lower tongue of Breathitt
Formation (arrow) (Pennsylvanian) above topmost vertical face of Slade limestone. Milepost 60.7, Interstate Highway 75, Mount Vernon

quadrangle, Rockcastle County.

The limestone, dolomite, shale, and sandstone of the
Kidder, Hartselle, Bangor, and the equivalent upper
member of the Newman (Slade) form a relatively
uniform succession of lithologic subunits which can be
traced along the outcrop belt bordering the Eastern
Kentucky coal field (Butts, 1922; McFarlan and Walker,
1956). Carbonate rocks in the lower part of both the
Kidder and upper member of the Newman (Slade) con-
sist of bioclastic and oolitic calcarenite alternating with
calcilutite and dolomite. To the northeast, the lower in-
terval is dominantly calcilutite. The calcarenites and
calcilutites form multiple fining-upward sequences, com-
monly capped by exposure zones which, in part, are
altered by secondary silicification and dolomitization.
A prominent exposure zone which can be traced the
length of the outcrop belt caps the interval of fining-
upward cycles. It is overlain by the Cave Branch Bed
(Dever, 1980a), a thin but widespread shale that forms
a useful marker within the carbonate section.

Overlying the Cave Branch Bed in the middle part
of the Kidder and upper Newman (Slade) are oolitic and
bioclastic calcarenite and thick-bedded, crinoidal
calcirudite, characterized by an association of Agassizo-
crinus, Pentremites, and a distinctive, large, uniden-
tified crinoid columnal (McFarlan and Walker, 1956). In
northeaStern Kentucky, deposits of calcilutite and
dolomite underlie this subunit. The crinoidal calcirudite

 

is overlain by a complex sequence of bioclastic and
oolitic calcarenite, calcilutite, dolomite, and shale. The
presence of nodular chert is a distinctive feature of these
upper calcarenites, particularly in east-central and
northeastern Kentucky.

The Hartselle Sandstone in south-central Kentucky
and along the Kentucky-Tennessee State line consists
of very ﬁne to medium-grained quartzose sandstone and
shale. N ortheastward along the outcrop belt, the sand-
stone pinches out and the Hartselle and its equivalent
in the upper Newman (Slade) are composed of shale
characterized by thin, discontinuous beds and lenses of
calcilutite. The Bangor and correlative limestone of the
upper Newman (Slade) are mainly bioclastic calcarenite
containing zones of dolomitic limestone and locally,
oolitic calcarenite. The limestones are commonly argil-
laceous, in part silty and sandy, and locally cherty; they
generally are darker colored than limestones of the
Kidder and underlying Newman (Slade). The unit is
thick bedded in the lower part and grades upward into
thinner beds with shale partings, forming a gradational
contact with the Pennington (Paragon) Formation. The
upper, thinner bedded interval is very fossiliferous.

The Pennington (Paragon) and correlative rocks of
northeastern Kentucky are composed of shale, sand-
stone, siltstone, limestone, and dolomite. The sequence
of lithologies, in ascending order, consists of (1) gray

80

MISSISSIPPIAN ROCKS IN KENTUCKY

 

FIGURE 44.—Section at Strunk Crushed Stone Company quarry, Tateville, Burnside quadrangle, Pulaski County.
Monteagle Limestone (Mm); Hartselle Formation (Mh); Bangor Limestone (Mb), and Paragon (formerly Penn-
ington) Formation (Mp).

ROCKS OF CHESTERIAN AGE 81

shale, (2) interbedded dolomite and limestone (south) or
sandstone (north), and (3) gray shale grading upward
into red and green shales with interbeds of dolomite and
siltstone. In south-central Kentucky, carbonate rocks
are important constituents of this dominantly
terrigenous-clastic sequence. Fossiliferous limestone is
interbedded with the basal gray shale, which is overlain
by an interval of interbedded dolomite and limestone;
a thin but widespread limestone occurs near the middle
of the formation, and dolomites are associated with the
upper red and green shales (Ettensohn and Chesnut,
1979a, 1979b). Northeastward along the outcrop belt,
carbonate-rock deposits are restricted mainly to the
middle limestone unit and upper dolomites. Sandstones
occur in the middle and upper Pennington (Paragon) of
south-central Kentucky and in the middle and lower
parts of the unit in east-central and northeastern Ken-
tucky. They commonly are slightly argillaceous and
micaceous; but orthoquartzitic sandstones, similar to
those of the Mississippian and Pennsylvanian Lee For-
mation, are present within the sequence and include the
Carter Caves Sandstone of northeastern Kentucky and
a sandstone in the upper Pennington (Paragon) of south-
central Kentucky (Smith, 1978). Thin coals have been
found locally in the Pennington (Paragon) in east-central
Kentucky (Ettensohn, 1977; Ettensohn and Peppers,
1979; Rice, 1972).

In southeastern Kentucky, rocks of Chesterian age
are present along Pine Mountain in the Newman Lime-
stone (upper member and part of lower member), Penn-
ington Formation, and tongues of Lee Formation
(Campbell, 1893) occurring within the Pennington.
Chesterian correlatives exposed on Cumberland Moun-
tain along the Kentucky-Virginia and Kentucky-
Tennessee State lines are the upper member of the
Newman Limestone, Pennington Formation, Pinnacle
Overlook Member (Englund, 1964a), and Chadwell
Member (Englund, 1964a; Englund and others, 1963)
of the Lee Formation. The White Rocks Sandstone
Member of the Lee Formation, formerly listed as
Mississippian in age, is now assigned to the Pennsyl-
vanian (Englund and others, 1985).

The Newman on the Pine Mountain overthrust block
has been divided into two informal members. The lower
member is composed of limestone and minor dolomite;
the upper member is mainly shale, with lesser amounts
of limestone, dolomite, siltstone, and sandstone. The ap-
proximate boundary between Meramecian and
Chesterian equivalents is within the lower member,
about 250 ft below the top, based on the position of the
Taggard Red Member of the Greenbrier Formation
(Wilpolt and Marden, 1959). The Taggard is a distinc-
tive unit, commonly from 5 to 20 ft thick, consisting
of grayish-red to grayish-red-purple and greenish-gray

 

calcarenite, calcilutite, mudstone, and shale. It is a
useful marker in the central and northeastern parts of
the Pine Mountain belt; in the northeast, however,
distinct units with similar coloration and lithology also
occur at several positions in the lower half of the lower
member of the Newman.

Limestones of the Newman lower member above the
Taggard are principally bioclastic calcarenites. Im-
mediately above the Taggard, they are partly oolitic and
interbedded with calcilutite and dolomite. The calcar-
enites are increasingly argillaceous and darker colored
upward within the member and, in the upper part, inter-
bedded with argillaceous limestone and thin shales.
Fossil content also increases upward within the
member. Chesterian crinoid genera, Talarocn'nus and
A gassizocrinus, have been found in the lower member
in Harlan and Letcher Counties (Greenfield, 1957).

The calcarenite, argillaceous limestone, and shale of
the lower member grade upward into the upper member
of the Newman, which is mainly medium gray to dark-
gray and greenish-gray shale, with lesser amounts of
siltstone. Limestone is interbedded with calcareous
shale in the lower part; sandstone, in discontinuous
lenses, occurs in the middle and upper parts. At the
northeastern end of Pine Mountain, a zone of grayish-
red and greenish-gray shale with lenses of dolomite is
present at the top of the Newman (Alvord and Miller,
1972).

Along Pine Mountain the Newman is overlain by the
Pennington Formation, which consists mainly of
siltstone, sandstone, and shale. Rocks assigned to the
Pennington principally are of Late Mississippian age,
but the presence of Pennsylvanian rocks locally in the
upper part of the formation is indicated by the occur-
rence of the spore, Laevigatosporites (mostly L. ovalis)
(Maughan, 1976). Average thickness of the formation
is about 450 ft in the southwestern part of Pine Moun-
tain, 900 ft in the central part, and 750 ft in the north-
east. At Pineville, the Pennington thickens abruptly,
from 500 ft on the west side to 1,100 ft on the east side
of the Rocky Face fault, a north-south-trending fault
with strike-slip displacement (Froelich and Tazelaar,
1974).

The N ewman-Pennington contact in the northeastern
part of Pine Mountain is placed at the base of the Stony
Gap Sandstone Member (Reger, 1926) of the Penn-
ington, which was mapped along the outcrop belt in
Pike County and part of Letcher County (Elkhorn City
through Whitesburg quadrangles). To the southwest,
correlative rocks were mapped as part of the lower
member of the Pennington. The Stony Gap is dominant-
ly crossbedded, quartzose sandstone, with tongues and
interbeds of variegated shale and siltstone. Southwest-
ward from central Letcher County, sandstone in the

82 MISSISSIPPIAN ROCKS IN KENTUCKY

lower Pennington commonly is less quartzose, ripple-
bedded, and interbedded and interlaminated with shale
and siltstone. In the southwestern part of the outcrop
belt, siltstone and shale are dominant lithologies in the
lower member (Newell, 1975; Rice and Newell, 1975).

The upper Pennington above the Stony Gap in the
northeastern part of Pine Mountain consists mainly of
siltstone and shale, with lenticular bodies of sandstone.
The siltstone and shale range from variegated grayish
red and greenish gray to dark gray and carbonaceous.
Lithologic descriptions on the geologic quadrangle
maps indicate that the sandstone content of the upper
Pennington is greater to the southwest. Tonguelike
bodies of conglomeratic, quartzose sandstone, similar
to that of the overlying Lee Formation, occur in at least
two positions within the Pennington (Maughan, 197 6).
These have been mapped as the “lower tongue of Lee
Formation” (Englund and others, 1964), “sandstone
tongue(?) of Lee Formation” (Csejtey, 1970; Froelich,
1973; Froelich and Tazelaar, 1974), and “sandstone
tongue of Lee(?) Formation” (Rice and Maughan, 1978).
The sandstone or lower tongue of Lee occurring in the
upper member of the Pennington in the Pineville area
(Englund and others, 1964; Froelich and Tazelaar, 1974;
Rice and Maughan, 1978) is considered to be correlative
with the Pinnacle Overlook Member of the Lee on
Cumberland Mountain (Rice, 1984).

The Little Stone Gap Member (Miller, 1964) of the
Pennington, composed of marine fossiliferous shale, silt-
stone, and limestone, was mapped in Pike County and
part of Letcher County. There, it occurs near the middle
of the Pennington section above the Stony Gap
Member. To the southwest, discontinuous deposits of
fossiliferous, calcareous shale and siltstone, and lime-
stone near the middle of the formation may be correla-
tive with the Little Stone Gap. Marine fossil-bearing
beds also occur locally at various positions in the upper
and lower parts of the Pennington.

Coal beds are fairly common in the upper Pennington
of Letcher and Harlan Counties, and one bed has been
reported at the base (Maughan, 1976). The coals general-
ly are thin, but locally are as much as 3 ft thick
(Froelich, 1973). Plant remains occur in sandstones of
the Pennington, including the Stony Gap Member.

The youngest Mississippian unit in southeastern
Kentucky is the Chadwell Member of the Lee Forma-
tion, which crops out along Cumberland Mountain. Dur-
ing the mapping program, the Chadwell (previously
designated as “sandstone member A”) and the generally
overlying White Rocks Member were considered to be
of probable Pennsylvanian age, but they were assigned
a Late Mississippian age by Englund (1979). Englund
and others (1985) later reassigned the White Rocks back
to the Pennsylvanian designation. The older Pinnacle

 

Overlook Member of the Lee (previously designated as
“lower tongue of Lee Formation”), which also crops out
on Cumberland Mountain, occurs mainly as a tongue
of Lee sandstone within the Pennington. All three
members consist of very light gray to white, fine- to
coarse-grained, quartzose sandstone which is partly con-
glomeratic, with well-rounded quartz pebbles 1%; to 1 in.
in diameter. They are thick bedded to massive, and
crossbedded. The presence of an unusual abundance of
quartz pebbles is the main distinguishing characteristic
of the White Rocks Member (Englund, 1964a).

The Chadwell Member overlies the Pennington. Basal
sandstone of the Chadwell intertongues with upper
Pennington shale and laterally grades into argillaceous
sandstone of the Pennington (Englund, 1964a). South
of Cumberland Gap, the upper member of the Penn-
ington pinches out and the Chadwell rests on the Pin-
nacle Overlook Member (Englund, 1964b). The White
Rocks Member (Pennsylvanian) locally overlies the
Chadwell. The White Rocks thins southwestward from
its type locality on the crest of Cumberland Mountain,
near Ewing, Va., and pinches out southwestward about
5 mi northeast of Cumberland Gap. Thicknesses of the
Lee members are varied; they reach a maximum of
about 350 ft for each member. In southwestern Virginia,
northeast of the Kentucky outcrop, the Chadwell and
Pinnacle Overlook intertongue with rocks of the Blue-
stone Formation (equivalent to upper Pennington) and
pinch out to the northeast (Englund, 1979).

A Chadwell correlative may be present on Pine Moun-
tain in southwestern Harlan County, where the Penn-
ington is overlain, in ascending order, by conglomeratic
sandstone and marine shale (C.L. Rice, US. Geological
Survey, oral commun., 1981). The sandstone and shale
were mapped as the basal part of the Middlesboro(?)
Member of the Lee (Froelich, 1972).

LITHOLOGIC TRENDS AND INTERBASIN CORRELATIONS

A lithofacies depiction of the total Chesterian succes-
sion in the Eastern Interior basin (Sable, 1979a) results
from combining the many rhythmic alternations of
detrital and carbonate units, which are modified by ir-
regular amounts of erosion in the upper part of the in-
terval. In the most complete Chesterian rock sections
in the southern part of the basin, sandstones are most
abundant in southern and southeastern Illinois, and are
common in western Kentucky. Lobate sandstone
tongues converge towards the center of the basin in
southern Illinois as shown by Swann and Bell (1958).
Increase of carbonate rocks along the southernmost
limits of Chesterian strata in west-central and south-
central Kentucky results mostly from the southward
thinning and disappearance of pre—Big Clifty detrital

ROCKS OF CHESTERIAN AGE 83

units into the Girkin Formation (fig. 45). A westward
increase in mudstone in Chesterian units such as in the
Big Clifty is reported by Swann (1963, p. 15; 1964,
p. 649) and is shown by comparison of Kentucky GQ
maps (for example, Big Clifty quadrangle versus Olney
quadrangle); a southeastward increase in the mudstone:
sandstone ratio also occurs in post-Glen Dean rocks of
west-central Kentucky.

Relationships between Chesterian units of the East-
ern Interior basin and presumed equivalents in the Ap-
palachian basin are critical for understanding the
Mississippian System in Kentucky. These rocks have
been removed by erosion along the axis of the Cincin-
nati arch; outcrops in the two basins are about 50 mi
apart at their nearest point in the Cumberland saddle
area of south-central Kentucky. Two major lithologic
trends indicate a continuity of deposition between the
two basins. The southeastward decrease in detrital
input in the Eastern Interior basin during early
Chesterian time extended into the western part of the
Appalachian basin, as reﬂected by the dominance of car-
bonate rocks in both the Girkin Limestone, west of the
Cincinnati arch, and the Kidder Member of the Mont-
eagle Limestone, east of the arch. Chesterian rocks in
most of the Eastern Interior basin, as noted previously,

Big Clifty Sandstone
Member of
Golconda Formation
Beech Creek
Limestone
Member of
Golconda
Formation
ypress
Formation

Paint Creek
Limestone

Bethel
Sandstone

Renault
Limestone

 

consist of a cyclic sequence of limestones alternating
with sandstones and shales. In contrast, the upper part
of the Chesterian succession in both the eastern part
of the Eastern Interior basin (Buffalo Wallow and
Leitchfield Formations) and the western part of the Ap-
palachian basin (Paragon and Pennington Formations)
mainly consists of shale with lesser amounts of sand-
stone and relatively minor amounts of limestone and
dolomite.

Specific correlations between Chesterian-age units of
the Eastern Interior and Appalachian basins, based on
megafaunal elements and the similarity and sequence
of lithologic units, were proposed by Butts (1922),
Stokley and McFarlan (1952), and McFarlan and Walker
(1956). The stratigraphic units and nomenclature of
McFarlan and Walker (1956) have been utilized in
reports of the Kentucky Geological Survey and in
papers and theses of other Kentucky workers, but were
not considered applicable by the US. Geological Survey
during the geologic mapping project because of lack of
mappable continuity and specific paleontologic and
lithologic correlations of many units east and west of
the Cincinnati arch.

The Hartselle Formation, as mapped in eastern south-
central Kentucky (Lewis and Thaden, 1965), the only

i:
.9
,_.
(U
E
l.
O
u.
.E
x
.L'
(3

 

FIGURE 45.—Relationships of Girkin Formation with coeval units of Chesterian age as mapped in the Homer quadrangle. west-central Ken-
tucky. A, typical of the western part of the quadrangle; B, typical of central and eastern parts (modified from Gildersleeve, 1966).

84 MISSISSIPPIAN ROCKS IN KENTUCKY

sandstone occurring in the carbonatedominated section
between the Ste. Genevieve and Pennington (Paragon),
may be a key to understanding relationships between
Chesterian—age units in the two basins. Correlation with
two different sandstone-bearing units west of the Cin-
cinnati arch has been suggested: Hardinsburg Sand-
stone (Stokley and McFarlan, 1952; McFarlan and
Walker, 1956; Lewis and Thaden, 1965; Ettensohn,
1980) and Big Clifty Sandstone Member of the Golcon-
da Formation (Butts, 1922; Peterson, 1956; Vail, 1959;
Horowitz and Strimple, 1974; Sable, 1979a; Horowitz
and others, 1979). It should be noted that the wide-
spread shale facies in the Hartselle was correlated with
the Golconda by Butts (1922) because in northern Ten-
nessee, the shale lies above the sandstone facies which
was correlated with the “Cypress” (Big Clifty).

Correlation of the Hartselle with the Hardinsburg is
based in part on characteristics of the enclosing lime-
stones. The Hartselle is overlain by limestone similar
to the Glen Dean of west-central Kentucky, and is
underlain by limestone similar to the Haney Limestone
Member of the Golconda Formation, and at some places
containing chert like that in the Haney. However,
Haney-type chert in limestone overlying the Hartselle
has been reported in the vicinity of Albany, Clinton
County (see section, “Carbonate-dominated units,”
p. 73).

Microfaunal elements from a core taken in northern
Tennessee, just south of the Kentucky-Tennessee State
line, show that the Hartselle is overlain and underlain
by beds of Golconda age, the basal Bangor and upper-
most Monteagle respectively. This indicates that the
Hartselle is correlative with the Big Clifty, the middle
member of the Golconda in the Eastern Interior basin
(Horowitz and others, 1979). Massive sandstones are
well developed in the Hardinsburg Sandstone in the
northern part of west-central Kentucky, but the
sandzshale ratio of the Hardinsburg decreases
southward (compare Mattingly and Homer geologic
quadrangle maps). The Hartselle in south-central Ken-
tucky and north-central Tennessee contains consider-
able sandstone. It thus appears that the Hardinsburg
clastic ratios do not correspond with those of the
Hartselle. The Big Clifty, on the other hand, is largely
massive thick-bedded sandstone to its erosional limits
in western south-central Kentucky; and its lithofacies
pattern can, with a fair degree of assurance, be com-
pared favorably to that of the Hartselle. Directional
depositional components (crossbedding and lobes of
sandstone lithofacies (fig. 46)) of the Big Clifty in west-
and south-central Kentucky (Potter and others, 1958),
if they represent primary sand input, suggest westward
and northwestward transport through the Cumberland
saddle (unlike most Michigan river southwestward

 

transport features). This may further suggest that the
thick Big Clifty sandstones may be related to the Hart-
selle in Kentucky.

In addition to possible correlation of the Hartselle
with units west of the Cincinnati arch, a southwestern
source for the type Hartselle Sandstone of Alabama was
suggested by Thomas and Mack (1982). The body of
sandstone in northern Alabama pinches out to the
northeast (Thomas, 1972, 1974). Its exact relationship
with the sandstone in the Hartselle of northern Ten-
nessee and south-central Kentucky, about 100 mi north,
is uncertain because detailed studies along the outcrop
belt in Tennessee have not yet been done. Thus, the
Hartselle Sandstone of Alabama may or may not equate
with that in Kentucky.

During the mapping program, rocks in the unit that
Butts (1922) had identified as Glen Dean in south-
central and east-central Kentucky were assigned to two
formations. The thick- to medium-bedded bioclastic
limestone forming the main part of the Glen Dean of
Butts (1922) was designated as the Bangor Limestone
in south-central Kentucky; it was included in the up-
per Newman Limestone (Slade Formation) in east-
central and northeastern Kentucky. The shale and in-
terbedded limestone at the top of the Glen Dean of
Butts (1922) were included in the Pennington Forma-
tion; these uppermost beds are the principal source of
Glen Dean—age megafauna (Butts, 1922; Horowitz,
1965; Weir and others, 1971; Ettensohn and Chesnut,
1979b; Chesnut, 1980). The Bangor in the northern Ten-
nessee core studied by Horowitz and others (1979) con-
tains foraminifera and conodonts which indicate a Glen
Dean age for at least part of the formation; the older
Golconda age, as noted above, was determined for the
basal Bangor.

On Pine Mountain in southeastern Kentucky, lime-
stone similar to the Bangor is present in the upper part
of the lower member of the Newman. The overlying
upper member of the Newman is correlated with the
Bluefield Formation of southern West Virginia and
southwestern Virginia (Wilpolt and Marden, 1959),
which generally is considered to be of Glen Dean age
(Butts, 1940; Cooper, 1944).

Age determinations for the unit designated as
Paragon Formation of Ettensohn and others (1984)
(formerly Pennington Formation), along the western
border of the Appalachian basin, and reported correla-
tions with Upper Mississippian units of the eastern Ap-
palachian basin and Eastern Interior basin are not
entirely compatible. Paleontological data indicate the
presence of younger rocks in the Paragon than are sug-
gested by lithostratigraphic correlations. Recent studies
of the palynology and microfauna indicate a Glen Dean
to Grove Church age for the Paragon in and near the

ROCKS OF CHESTERIAN AGE

88° 87° 86° 85°
I J ( ( i ‘ s I I I

 

 

39" — EXPLANATION ,7 ' 1.,“ § .7
»- Sandstone <— General direction of sediment ,jﬁ) 1 Q. n ,2 j V
' transport (from Potter and a?“ ,1 .. ,me \v A \
o 30 70 100 PERCENT others. 1958) {w ~- , g» , _ a ,7 j
"\ ;

 

   

it MNQKS

37° ——

  

 

 

 

 

)UU MILES

FIGURE 46.—Distribution of sandstone and shale in Big Clifty Sandstone Member of Golconda Formation in west-central, central, and western
Kentucky. and Hartselle Formation in south-central Kentucky, showing sediment transport directions. Kentucky fault patterns shown

on base.

outcrop belt along the western border of the Eastern
Kentucky coal field (Ettensohn and Peppers, 1979;
Horowitz and others, 1979). Rocks designated as Para-
gon in this outcrop belt have been correlated with the
upper member of the Newman, or Bluefield, on the Pine
Mountain overthrust block (Englund and Windolph,
1971; Englund and Randall, 1981). The Bluefield, as
noted previously, is considered to be Glen Dean in age.

Englund and Henry (1979) and Englund and Randall
(1981) indicated that as a result of progressive west-
ward, post-Mississippian truncation of Upper Mississip-
pian units, the equivalents of the Pennington Formation
on the Pine Mountain overthrust block are not present
in the Paragon outcrop belt along the western border
of the Eastern Kentucky coal field. The Pennington on
Pine Mountain is approximately correlative with the
Hinton Formation, Princeton Sandstone, and Bluestone
Formation of West Virginia and Virginia (Wilpolt and
Marden, 1959), which range in age from Menard to
Grove Church or younger (Gordon and Henry, 1981).

SOURCES OF SEDIMENTS AND
DEPOSITIONAL ENVIRONMENTS

An ancient southward-ﬂowing river system, first en-
visioned by Stuart Weller (1927, p. 26) and later named
the Michigan river (Swarm, 1963, p. 12), is interpreted

 

to have transported detrital Chesterian sediments into
and across the Eastern Interior basin. Evidence for the
Michigan river consists of regional thickness and facies
distribution of Chesterian units (Swann and Bell, 1958),
southward-dipping crossbeds in generally southward
trending sandstone bodies (Potter and others, 1958;
Potter, 1963), and channel-fill morphology of sandstone
bodies (Potter, 1962, p. 28—29; 1963; Reynolds and
Vincent, 1967; Sedimentation Seminar, 1969). The main
sediment source was in eastern Canada, either in the
Canadian shield or east of the northern Appalachian
fold belt. The possibility that the uplifted Franklinian
geosyncline in northern Canada may have been a major
source was also suggested by Swarm (1964). According
to Swann, the Michigan river ﬂowed southwestward
across Michigan and northern Indiana into a shallow
sea, where detrital sediments accumulated as a birds-
foot delta projecting southward beyond a N. 65° W.-
trending shoreline. Lateral northwest or southeast
shifts in the course of the Michigan river of as much
as 200 mi produced belts of sand and mud in different
parts of the region at different times. Major northeast
and southwest oscillations of the shoreline, perhaps as
much as 600 to 1,000 mi, produced numerous marine
transgressions during which carbonates were deposited
in the basin and surrounding areas and the amount of
clastic detritus was relatively low. The resulting highly

86 MISSISSIPPIAN ROCKS IN KENTUCKY

variable but rhythmic depositional complex of
sediments includes shallow-marine detrital and car-
bonate facies, littoral detrital facies, and continental
detrital and coaly facies. Superimposed on the system
of shifting shorelines and positions of the river system,
Swarm (1963, p. 14—15) postulated a northwest-ﬂowing
sea current or drift which carried muds northwestward,
leaving relatively clear water in the southeastern (south-
central Kentucky) part of the basin where limestone ac-
cumulated. Sea depths in the Eastern Interior basin
during Chesterian time have been estimated on the
basis of water depths near modern deltas, increased by
the estimated amount of compaction of the sediments
deposited. These estimates place the depth of water at
50—75 ft, with a range of from 30 to 100 ft (Swarm, 1964,
p. 652—653).

A low dip of the paleoslope, low relief of the sea ﬂoor,
and a widespread shallow-marine environment are in-
dicated by (1) the distribution of tabular carbonate units
which individually cover thousands of square miles and
indicate similar depositional environments across these
large areas, (2) the presence of oolitic limestones and
current-bedding features, and (3) numerous shallow-
water fossils. Likewise, the widespread sheets of ter-
rigenous clastics also indicate generally uniform
shallow-marine, shoreline and lower delta plain environ-
ments, in contrast to environments of the earlier Borden
deltaic beds, which prograded into deep water and were
onlapped by deep-water carbonates deposited along the
delta front.

The major sources for terrigenous elastic sediment of
the Pennington Formation were highlands of sedimen-
tary and metamorphic rocks lying generally east of the
Appalachian basin (de Witt and McGrew, 1979).
Whether detrital material from Appalachian basin
source areas contributed to Chesterian rocks west of the
Cincinnati arch in the Eastern Interior basin as may
be suggested by Big Clifty—Hartselle distribution (p. 84)
is uncertain. Some contribution in late Chesterian time
may also be suggested by lithologic similarity of the
Pennington and Paragon Formations of eastern Ken-
tucky and the Leitchfield and Buffalo Wallow Forma-
tions of west-central Kentucky. Carbonate-dominated
units of earlier Chesterian and Meramecian age were
deposited across the Cumberland saddle area as shown
by their lithic and faunal similarities, and their thick-
ness trends indicate that the saddle was then a negative
element athwart the Cincinnati arch. In southeastern
and eastern Kentucky, sandstones in the Pennington
and Paragon, which thin to extinction westward and
southward, have been interpreted to have been trans-
ported southwestward along the Appalachian basin
from a northeastern source area (Vail, 1959, p. 59), but
in view of the proposed Paragon-Bluefield correlations

 

discussed previously (p. 84—85), this conclusion may
be invalid. The northwest sea-current drift proposed
by Swann, which moved muds discharged by the
Michigan river westward, explains the mudstone in
the western part of the Eastern Interior basin. A
connection across the Oimberland saddle area (Cumber-
land strait) also could provide a passage for late
Chesterian (Pennington-Paragon) fine-grained sedi-
ments from Appalachian sources carried by westward-
ﬂowing marine currents to form Leitchfield-Buffalo
Wallow clastic strata in the Eastern Interior basin.
Mixing of Appalachian sediments with those from the
Michigan river would thus have taken place in west-
central Kentucky.

Depositional environments of the rocks of Chesterian
age in the upper member of the Newman (Slade), the
Kidder-Hartselle-Bangor sequence, and Pennington
(Paragon) along the western border of the Eastern Ken-
tucky coal field have been studied extensively in recent
years. Carbonate rocks in the lower part of both the
Kidder and the upper member of the Newman (Slade),
below the Cave Branch Bed, form multiple fining-
upward sequences. These sequences indicate
transgressive-regressive cycles, with shallow, subtidal
carbonate sands grading upward into prograding tidal-
ﬂat and supratidal deposits of lime mud and silt. (Dever,
1973, 1977). Individual sets of supratidal and tidal-ﬂat
deposits commonly are capped by exposure zones. The
entire interval of fining-upward sequences is capped by
a prominent exposure zone which can be traced the
length of the outcrop belt of these units.

Limestone, shale, and sandstone of the upper
Newman (Slade) overlying the widespread exposure
zone, and sediments of the younger Pennington
(Paragon) Formation were deposited in a major
transgressive-regressive sequence (Ettensohn, 1975,
1977, 1980; Ettensohn and Chesnut, 1979b). The trans-
gressive sequence consists of intertidal mud ﬂats (Cave
Branch Bed), lagoonal lime mud (calcilutite), carbonate
sand belt (calcarenite and calcirudite), shallow open-
marine deposits (fossiliferous calcarenite, calcilutite,
and shale), and deeper open-marine deposits (shale and
calcilutite). The change from transgression to regres-
sion is considered to have occurred during deposition
of this open-marine shale, an equivalent of the Hartselle.
Succeeding regressive deposits consist of carbonate
sand belt (calcarenite; Bangor equivalent), lagoonal
deposits (fossiliferous shale and limestone of basal Pen-
nington (Paragon)), and carbonate and detrital tidal
ﬂats with local coal-forming swamps (dolomite and
sandstone; see Ettensohn and Peppers, 1979). The
regressive sequence was temporarily interrupted by a
marine advance with deposition of a thin but wide-
spread limestone. Succeeding the limestone are prodelta

MISSISSIPPIAN-PENNSYLVANIAN-YOUNGER ROCK RELATIONSHIPS 87

deposits of prograding shoal-water deltas or tidal ﬂat
sediments (shale, siltstone, and dolomite of uppermost
Pennington-Paragon). As each of the depositional en-
vironments migrated with transgression or regression
across the region, its lithologic expression assumed a
sheetlike geometry forming a stack of widespread,
tabular lithologic units (Ettensohn, 1977).

The linear body of the Carter Caves Sandstone in
northeastern Kentucky has been described variously as
an offshore bar (Englund and Windolph, 1971), a
beach-barrier island system (Home and others, 1974),
a tidal-channel deposit paralleling the Waverly arch
(Ettensohn, 1977), and a distributary-channel deposit
of a Pennington (Paragon) delta system (Short, 1978).
Deposition of Chesterian-age rocks in northeastern and
east-central Kentucky was inﬂuenced by the Waverly
arch, which was a positive feature, and by recurrent
movement along the Kentucky River fault system, as
described in Dever and others (1977), Ettensohn and
Dever (1979), and Ettensohn and Peppers (1979).

Depositional environments of the Newman and Pen-
nington in southeastern Kentucky on the Pine Moun-
tain overthrust block have not yet been studied
extensively, but the environments of deposition for cor-
relative units in southwestern Virginia and southern
West Virginia have been discussed by Englund (1979),
Englund and others (1981), and Miller (1974). Rocks of
Chesterian age show a general upward transition from
shallow-marine shelf and nearshore-marine deposits to
nearshore tidal-ﬂat, beach, lagoon, and marsh deposits.
The Pennington, in particular, is characterized by
deposition in alternating near-coastal terrestrial and
nearshore marine environments.

PALEOGEOGRAPHY AND PALEOTECTONIC IMPLICATIONS

In general, stream and shoreline positions lay in
western Kentucky through much of Chesterian time.
The Michigan river system was the dominant transport-
ing agent. A delta-complex at its mouth included both
sheet sands and linear sand bodies such as channel fills,
point bars, and distributary mouth bars. Sediment
discharged by the Michigan river may have been
deposited as far south as northwestern Alabama and
northeastern Mississippi. Only minor coaly beds formed
in interdistributary swamps on the delta plain, suggest-
ing that rates of clastic deposition were high, with rapid
shifts of distributaries and splay channels. Highlands
east of the Appalachian basin contributed detrital
sediments, some of which in late Chesterian time were
probably carried through the Cumberland saddle and
intermingled with Michigan river sediments in western
Kentucky. The northern part of the Cincinnati arch in
Ohio and possibly Kentucky was a low-lying peninsula

 

or shoal area which contributed little sediment but may
have acted as a partial barrier between the Eastern In-
terior and Appalachian basins.

Abundant marine faunas and oolitic limestones in-
dicate subtropical or tropical conditions during
Chesterian time. Scattered coaly beds and fairly abun-
dant fossil plant remains in sandstones and siltstones
suggest humid terrestrial conditions. Red shales are
common in some detrital units such as the Big Clifty,
Pennington, and Leitchfield Formations; they may in-
dicate the presence of deep residual soils in source areas
(Vail, 1959, p. 52).

Mild but persistent subsidence characterized broad
areas of western Kentucky, and sediment thicknesses
indicate that the greatest subsidence was in the Fair-
field basin in Illinois and Moorman syncline in Ken-
tucky. The lateral persistence of individual formations
and their small thickness variations show that sub-
sidence was relatively even. Minor differential move-
ments probably occurred along the La Salle anticlinal
belt, folds in Illinois and Indiana (Siever, 1951, p. 569),
and the Rough Creek and Pennyrile fault systems. A
regional southwestward-dipping paleoslope and a mildly
negative trough across Michigan, Indiana, Illinois, and
western Kentucky controlled the trend of the Michigan
river system. The area of the Cumberland saddle is
presumed to have been a mildly downwarped element
between the Jessamine and Nashville domes that con-
nected eastern and western basins in Kentucky.

In eastern Kentucky, the area between the present
Cincinnati arch and the Appalachian basin, the Waverly
arch, was a relatively high and unstable platform during
Chesterian age deposition. Minor differential movement
occurred along the Kentucky River fault system.

MISSISSIPPIAN-PENNSYLVANIAN-
YOUNGER ROCK RELATIONSHIPS

Detrital rocks of terrigenous origin of Pennsylvanian
and Cretaceous age unconformably overlie rocks of all
Mississippian series in Kentucky. Only in the southeast-
ern part of Kentucky has it been described that Missis-
sippian and Pennsylvanian strata are conformable.
Recent literature on systemic relationships includes a
comprehensive synthesis by Rice and others (197 9).

Pennsylvanian rocks overlying Mississippian strata
are dominantly lower delta plain deposits of sandstone,
mudstone, conglomerate, clay, and coal with very minor
marine beds including limestone and mudstone. These
basal strata of Morrowan and Atokan ages are the lower
part of the Lee and Breathitt Formations in the eastern
part of Kentucky, and the Caseyville Formation in the
west. About 100 mi separates the two main belts of
exposures.

88 MISSISSIPPIAN ROCKS IN KENTUCKY

A major erosional hiatus between Mississippian and
Pennsylvanian rocks is well documented in much of
Kentucky; it is marked by progressively deeper bevel-
ing of pre-Pennsylvanian strata north of Kentucky in
northern Illinois and Indiana, where as much as 1,500
ft of Mississippian section may have been removed
(Sable, 1979a). Channels superimposed on a gently roll-
ing surface are as much as 450 ft deep in southeastern
Illinois (Siever, 1951). In western Kentucky, as much
as 900 ft of Mississippian strata is believed to have been
removed, with channels incised as much as 250 ft into
the Mississippian surface (Bristol and Howard, 1971).
In northeastern and east-central Kentucky pre-
Pennsylvanian erosion is evinced by a southward- or
south-southeastward-sloping beveled surface locally
marked by paleokarst topography developed on Missis-
sippian limestones (Dever, 1971; Weir, 1974b; Hoge,
1977). The beveling appears to have been controlled by
growth of the Cincinnati and Waverly arches (Englund,
1972). In west-central Kentucky, pre-Caseyville For-
mation beveling has truncated progressively older

 

Chesterian units in general eastward or southeastward
directions, toward the Cincinnati arch, as shown in
figures 47, 48, and 49; the illustrated localities lie along
an 18-mi line of section and show the Caseyville For-
mation lying on post-Kinkaid Limestone beds, on post-
Menard Limestone beds, and on the Vienna Limestone,
respectively, from northwest to southeast. Channeling
into this surface is locally pronounced, and relief along
the borders of the deeper channels is commonly about
60—80 ft. Reconstruction of Mississippian strata prior
to Pennsylvanian erosion indicates that 160 ft of Borden
and Newman and an undetermined amount of Penning-
ton may have been removed in the vicinity of the Bell
Branch channel in Pomeroyton quadrangle (Weir and
Richards, 1974). On the west side of Indian Fort Moun-
tain in the Bighill quadrangle (Weir and others, 1971),
as much as 240 ft of upper Newman (Slade) and
Pennington (Paragon) strata may have been present
prior to erosion.

Several southwest-trending sub-Pennsylvanian chan-
nels in central, west-central, and western south-central

 

FIGURE 47.—Sandstone of Caseyville Formation (PC) with basal thin coal bed (c) overlying limestone equivalent of Kinkaid Limestone in
Buffalo Wallow Formation (Mbwk). Intervening shale probably of Mississippian age. Caneyville Crushed Stone Company quarry, Caneyville
quadrangle. Grayson County.

MISSISSIPPIAN -PEN N SYLVAN IAN-YOUN GER ROCK RELATIONSHIPS 89

 

FIGURE 48.—Caseyville Formation (PC) disconformably overlying Buffalo Wallow Formation (Mbw). Exposed Buffalo Wallow units are
(descending) unnamed gray shale; 6-ft thick Menard Limestone equivalent (M bwm), unnamed greenish and reddish shale and dolomite
(Mbwd). Milepost 101.0, Western Kentucky Parkway, Caneyville quadrangle, Grayson County.

Kentucky are incised as much as 250 ft in the Mississip-
pian surface, following truncation which may have
removed as much as 1,000 ft of Mississippian strata
(Rice and others, 1979). East of and colinear with the
southernmost channel, isolated hilltop outliers of sand-
stone, shale, and conglomerate, interpreted to be
remnants of basal Pennsylvanian channel fill, extend
along a west-southwest linear trend across parts of Ed-
monson, Hart, and Larue Counties, west-central and
south-central Kentucky (Burroughs, 1923; McFarlan,
1943, p. 96). These outliers mapped in Hibernia (RE.
Moore, 1976) and other quadrangles overlie strata of the
Ste. Genevieve and upper part of the St. Louis Lime-
stones. Growth of the Cincinnati arch prior to channel
cutting is suggested by these relationships. The main
channel continues southwestward across the southern
part of the western Kentucky coal field as part of the
channel complex overlying Chesterian units, described
by Shawe and Gildersleeve (1969) and Sedimentation
Seminar (1978), and termed the Brownsville channel by
Bristol and Howard (1971, p. 9). Rice and Weir (1984)
interpreted the Brownsville paleochannel to be the
distal part of the much larger southeast-gradient

 

Sharon-Brownsville paleovalley system which headed
north of Pennsylvania.

In general, Pennsylvanian sandstones are composi-
tionally and texturally similar to Mississippian sand-
stones, and in many places unfossiliferous mudstone,
shale, and siltstone constitute the strata between un-
mistakable Mississippian and unmistakable Pennsylva-
nian rocks so that recognition of the systemic boundary
is difficult. Two-group discriminant functional analysis
of geochemical data of Pennsylvanian and Chesterian
sandstones was reported by Connor (1969) and Connor
and Trace (1970) to be of potential use in future
discrimination of similar rocks of the two systems.

In southeastern Kentucky, the Mississippian-
Pennsylvanian systemic boundary occurs within rock
sequences that generally were considered to be con-
formable. Along Pine Mountain, the boundary locally
is in shale and siltstone sequences of the Pennington
Formation. The Pennington rocks are mostly of Late
Mississippian age, but the occurrence of the spore,
Laevigatosporites (mostly L. ovalis), in the upper part
of the formation indicates the presence of rocks of
Pennsylvanian age (Maughan, 1976). On Cumberland

90 MISSISSIPPIAN ROCKS IN KENTUCKY

 

FIGURE 49.—Sa.ndstone of Caseyville Formation (Pc) overlying 8-ft thick Vienna Limestone (Mv). Basal Caseyville consists of 6 ft of quartz-
pebble conglomerate. Section below Vienna is 26 ft of Tar Springs Formation (Mts) shale. sandstone, and limestone overlying Glen Dean
Limestone (Mgd). Cardinal Stone Company Quarry, Bee Spring quadrangle, Edmonson County.

Mountain, Englund and Smith (1960), Englund (1964a),
and Englund and Delaney (1966) reported intertongu-
ing and lateral gradation between rocks of the Lee and
Pennington Formations which, at that time, generally
were considered to be of Pennsylvanian and Mississip-
pian age, respectively. Lower members of the Lee, the
Pinnacle Overlook, Chadwell (previously designated as
“sandstone member A”), and White Rocks Members,
which intertongue with the Pennington, later were
assigned a Late Mississippian age by Englund (1979).
The oldest Pennsylvanian unit on Cumberland Moun-
tain, according to Englund (1979), was the Dark Ridge
Member (previously designated as “sandstone and shale
member B”) of the Lee Formation, which apparently
was conformable with Mississippian rocks of the Lee
and Bluestone Formations. The Dark Ridge was de-
scribed as gradational and intertonguing with the
Chadwell and White Rocks (Englund, 1964a, 1964b;
Englund, Landis, and Smith, 1963). Rice (1984),

 

however, suggested that the systemic boundary may
be an unrecognized paraconformity within the Dark
Ridge. More recently, Englund and others (1985)
reassigned the White Rocks to the Pennsylvanian
System and showed the White Rocks and Dark Ridge
Members of the Lee Formation to unconformably
overlie Mississippian rocks of the Bluestone Formation
and Chadwell Member of the Lee Formation.

The conclusion that an unconformity between the
Mississippian and Pennsylvanian Systems extends
throughout eastern Kentucky is supported by a sub-
surface study that, in southeastern Kentucky and
southwestern Virginia, identifies a 350-ft-deep,
southwest-trending valley fill at the base of the Penn-
sylvanian called the “Middlesboro paleovalley” (Rice,
1985). The reassignment of the White Rocks Sandstone
Member of the Lee Formation from Mississippian to
Pennsylvanian age by Englund and others (1985) has
suggested to Rice (in Shepherd and others, 1986) that

MISSI SSI PPIAN—PENN SYLVANIAN -YOUN GER ROCK RELATIONSHIPS 91

sediments of the White Rocks filled the Middlesboro
paleovalley where that valley intersects Cumberland
Mountain.

An intra-Pennsylvanian unconformity originating
within the lower part of the New River Formation of
West Virginia and Virginia was considered by Englund
(1974, 1979) and Englund and Henry (1979) to be coex-
tensive with an unconformity at the base of the Middles-
boro Member of the Lee in southeastern Kentucky and
with an unconformity occurring between Mississippian
and Pennsylvanian rocks along the western border of
the Eastern Kentucky coal field. More recently the un-
conformity beneath the Pineville Member of the New
River Formation in West Virginia and Virginia was ex-
tended southwestward beneath the White Rocks, Dark
Ridge, and Middlesboro Members of the Lee Formation
along the Kentucky-Virginia State line where it is coex-
tensive with the Mississippian-Pennsylvanian systemic
boundary (Englund and others, 1985). Where the Mid-
dlesboro Member of the Lee overlies the Pride Shale,
its base marks the unconformity at the systemic bound-
ary according to CL. Rice (written commun., 1987).

The presence of an unconformity at the Pennsylvanian-
Mississippian systemic boundary in northeastern
Kentucky has been challenged by several workers.
Deposition of the Carboniferous rocks of northeastern
Kentucky as part of a gradational sequence of westward-
prograding ﬂuvial, deltaic, lagoonal, shoreline, and
marine sediments was proposed by Home and Ferm
( 1970), Swinchatt (1970), Ferm and others (1971), Home
and others (1971), Ferm (1974), and Home and others
(1974). Their depositional model involved rocks of the
Borden, Newman (Slade), Pennington (Paragon), Carter
Caves, Lee, and Breathitt; intra- and post-Mississippian
unconformities, identified by earlier workers, were con-
sidered to be facies or depositional boundaries. In the
model, orthoquartzites (Carter Caves, Lee) were inter-
preted as beach-barrier deposits which grade landward
into lagoonal-bay and deltaic-ﬂuvial shales, subgray-
wackes, and coals (Breathitt). The beach-barrier sand-
stones grade seaward into marine shales (Borden and
Pennington) which surround limestones (Newman)
deposited as carbonate barriers, bars, and tidal-ﬂat-
island complexes. The model was developed incor-
porating the exposures of Carboniferous rocks along
Interstate Highway 64 in Carter and Rowan Counties.

Study of the rock units along the interstate highway
and in adjacent areas has shown that the field relation-
ships do not support the proposed model (Dever, 1973;
Ettensohn, 1975, 1979a, 1980, 1981; Rice and others,
1979). For example, a large body of marine shale which
is shown in the model as occurring between and grada-
tional with isolated limestone bodies and with ortho-
quartzites does not exist at that stratigraphic position

 

(fig. 50, upper cross section). The area of northeastern
Kentucky used for the depositional model is on the up-
thrown block of the Kentucky River fault system and
astride the axis of the Waverly arch, but these tectonic
elements were not recognized during the model’s devel-
opment. Uplift and movement along these features
during the Carboniferous resulted in extensive intra-
and post-Mississippian erosion in the area and partly
controlled deposition of Mississippian and Pennsylva-
nian units (Dever, 1973; Dever and others, 1977;
Englund, 1972; Ettensohn, 1975, 1980, 1981; Ettensohn
and Dever, 1979; Ettensohn and Peppers, 1979; Haney,
1976; Haney and others, 1975; Home and Ferm, 1978;
Sergeant and Haney, 1980; Short, 1978).

Short (1978) also proposed that the Mississippian-
Pennsylvanian systemic boundary in northeastern Ken-
tucky is within a gradational sequence, with continuous
deposition from deltaic and marine sediments of the
Pennington into deltaic and ﬂuvial sediments of the Lee
and Breathitt. He recognized that local disconformities
occur in the sequence at the base of channel-fill sand-
stones, but suggested that the extensive intra-
Mississippian unconformity in the area commonly is
misidentified as a post-Mississippian unconformity.
Both Short (1978) and Ettensohn and Peppers (1979)
suggested a Late Mississippian age for the Olive Hill
Clay Bed of Crider (1913) which traditionally has been
assigned to the Pennsylvanian. Their suggested age is
based on the occurrence of palynomorphs that may in-
clude pollen from an eroded Mississippian upland, ac-
cording to CL. Rice (written commun., 1987).

On the other hand, Haney (1979, 1980) reported that
the systemic boundary is coincident with an unconform-
ity in areas that were tectonically active during the Car-
boniferous, such as northeastern Kentucky, but that in
areas distant from active tectonic features, such as
south-central Kentucky, the systemic boundary com-
monly is within a gradational sequence. In the latter
areas, occurrences of an erosional surface between Penn-
sylvanian and Mississippian rocks are considered to be
related to downcutting during intra-Pennsylvanian ero-
sion. Rice (1980), in rebuttal, contended that regional
relationships indicate the presence of a widespread un-
conformity coincident with the systemic boundary, and
he reported the development of an extensive consequent
drainage system after emergence of the region in Late
Mississippian time. Subsurface analysis of oil and gas
logs by C.L. Rice and others (written commun., 1987)
indicates that Lower Pennsylvanian strata thin and
onlap onto older Mississippian rocks northward in east-
ern Kentucky. This suggests that some localities in
northeastern Kentucky were emergent during Early
Pennsylvanian time—no strata were deposited there.
In parts of south-central Kentucky where no obvious

Lee Formation

 

 

    
 

MISSISSIPPIAN ROCKS IN KENTUCKY

EAST

Breathitt Formation

" FEET

 

4 MILES Modiﬁed from Farm and others |197i, fig, 13]
| and Home and others (1974, fig. 2] 10”

200

 

Slade and
Paragon Fms

 

 

 

 

 

 

   

BORDEN
FORMAT I0N

Orthoquartzite sandstone
Limestone

Dolomite

 

 
   

EXPLANATION

 

 

Red and green shale

 

 

_ Siltstone

 

 

 

Sandstone, siltstone, and

 

shale undifferentiated

FIGURE 50,—Generalized cross sections showing two interpretations of Pennsylvanian and Mississippian unit relationships along Interstate
Highway 64, northeastern Kentucky (from Rice and others, 1979, modified from Dever, 1973).

erosional surface is recognized between Mississippian
and Pennsylvanian rocks, the systemic boundary is con-
sidered to be a paraconformity which may be misinter-
preted as a normal bedding plane or facies-related
boundary (Rice, 1980).

Detrital rocks of the Tuscaloosa and McNairy Forma-
tions of Late Cretaceous (Gulfian) age lie on an erosional
surface of Mississippian and older rocks in the Missis-
sippi Embayment area of southern Illinois, westernmost
Kentucky, Tennessee, and southeastern Missouri. More
than 3,000 ft of Mississippian strata were removed by
pro-Cretaceous erosion in westernmost Kentucky and
southern Illinois, the result of the growth of the Pascola
arch and its denudation (Schwalb, 1969). There is little
indication of growth of the Pascola arch in western Ken-
tucky during Mississippian time; only the isopach pat-
tern of latest Chesterian sediments (fig. 64), may suggest
inﬂuence of a positive feature such as the Pascola arch
to the south and west of western Kentucky. Through
much of Mississippian time, however, a south-southwest-
trending depositional trough or basin in which more
than 3,300 ft of Mississippian rocks accumulated is
evidenced by the record of Mississippian strata. Thicken-
ing of most Mississippian stratigraphic intervals shown
in figures 54-63 is towards the Pascola arch.

 

SUMMARY OF THICKNESSES AND
LITHOLOGIC TRENDS

The generalized isopach maps (ﬁgs. 51—64) portray
thicknesses of single or multiple units of Mississippian
strata and are based largely on data from geologic quad-
rangle maps of Kentucky and individual measured sur-
face sections; they also rely upon subsurface data from
published reports and from files of the Kentucky Geo-
logical Survey. Isopach intervals are variable between
maps and also within some individual maps. Thick-
nesses shown reﬂect in-place depositional thicknesses
modified by subsequent burial and induration. The iso-
pachs give only general unit thicknesses for two reasons:
(1) Considerable range in unit thicknesses is given in col-
umnar sections of mapped quadrangles so that an
accurate representative average thickness figure for a
unit in any given quadrangle is difficult to obtain. The
average thicknesses used in thickness calculations were
therefore based on average thicknesses shown in the col-
umnar sections and augmented by map inspection. (2)
Some unit thicknesses given in well descriptions are
highly subjective because of varying interpretations by
different geologists who described the subsurface rocks

SUMMARY OF THICKNESSES AND LITHOLOGIC TRENDS 93

 

89° 88° 87° 86° 85° 34° 83° 82"
I I K I l l E « \. ,, l l ‘
39° _ l WT
,2 > » may. 3
2 f. {x
3 \V .
/ mammal ,5,

38° —

  
 
     
  

El “Nikita

  

00’» ~ “

,i» A m

[800 |

  

 

 

 

D 50 100 MILES

 

FIGURE 51.—Total thickness (in feet) of Mississippian rocks in Kentucky. Solid isopachs show thicknesses of strata underlying Pennsyl-
vanian rocks; dotted lines where overlain by Cretaceous and Tertiary rocks or in exposure areas of Mississippian rocks. See ﬁgure 2
for distribution of units. Selected faults and fault zones shown on Kentucky base.

 

  

89° 88° 87° 86° 82°
39° I I a! I I
— / T
k 4.: a
i \
/ iNE)lL‘aN.-¥ , ...... X
. I {W}
WEST

"I VIRGINIA

llLENOIS

37° —

   

 

 

 

glﬁfwu: WT. ,n. ,_ n . , . 'I l l l
l] 50 100 MILES
|_;_;l_l_J%—l

FIGURE 52.—Cumulative thickness (in feet) of Bedford Shale, Berea Sandstone. and Sunbury Shale in eastern part of Kentucky, and thickness

of approximate age-equivalent strata (Maury Formation, Rocldord Limestone. and Hannibal Shale equivalents) in other parts of Kentucky.
+, thickness uncertain, less than 10 ft.

94 MISSISSIPPIAN ROCKS IN KENTUCKY

89° 88° 87° 86° 85° 84° 83° 82°

 

39° _ l, \ T‘

 
  
  
   
   
 
 
 
 
  
    

1/ YNIREANA

38" ~

Elilfmﬂié

 

 

 

[1 El] 100 MILES

 

FIGURE 53.—Thickness (in feet) of terrigenous elastic units of the Borden Formation (New Providence, Nancy. Cowbell, Halls Gap. Wildie,
N ada, Holtsclaw, Farmers, Kenwood), New Providence and Grainger Formations, and basal shale beds (New Providence Memberi in
the Fort Payne Formation.

89° 88° 87° 86° 85° 84° 83° 82°

 

 

l
39°- EXPLANATION < :_’ N ,

Dominant to predominant E Cane Valley Limestone

dark-gray to brown shale Member of Fort Payne

and siltstone (from sub- Formation
surface records: limits
uncertain)‘

Outcrop and surface area of
crinoidal biocalcirudite and
green shale 10 to 165 ft
33° — thick mmsm

lLLlNOlb

 

 

 

 

[] 50 100 MILES

 

FIGURE 54.—Thickness (in feet) of Muldraugh Member of Borden Formation and of Fort Payne Formation exclusive of basal shale units,
and showing distribution of unusual lithic components.

SUMMARY OF THICKNESSES AND LITHOLOGIC TRENDS 95

 

     
   

 

59° 88° 37° 86° 85° 84° 83° 82°
I I \ | | T
39" _ a) .1 r“:
l f ,1
3
j ENDMEEA
38° — \? ——*

{LUNOXS

v: :43: MA

    

 
 

I ’HEE‘NESSEE‘ L l l

“j C300 <

 

 

 

U 50 100 MILES
\_l_;.1__x__J—|

FIGURE 55.—Combined thickness (in feet) of Muldraugh—Fort Payne-Salem-Warsaw—Harrodsburg units in central to western Kentucky,
and probable equivalent Renfro Member of Slade Formation (Msr) (previously of Borden Formation) in east-central and southeastern
Kentucky. Dashed line is arbitrary cut-off of Renfro Member isopachs.

39° 88° 87° 86° 85° 84° 83° 82°

 

Willi}

   

1/ IM’NANA

    

 

iUJNOlS:

 

 

 

[J 50 100 MILES

 

FIGURE 56.—Thickness (in feet) of Harrodsburg Limestone in central. south-central, west-central, and western Kentucky. Dashed isopachs
in areas of little control.

96

MISSISSIPPIAN ROCKS IN KENTUCKY

 

  
 
 
    

89° 88° 87° 86° 85° 84° 83° 32°
l | l | I in _ 1 l 1
39°_ m M M
: > , ‘3 OHM": {I
l j E z
/) INDIANA
I' /_.
f 1
/ x /'
x a,
“0/ wee?
» magma
aa°— f _
{LLINOIS

 

    

 
    

-\,’._ ‘-

J TENNEEaJEEMI ‘ “

 

 

 

 

  
   
   

[] 50 100 MILES
|_L__l_|_l__l—¥
FIGURE 57.——Cumulative thickness (in feet) of Salem, Salem-Warsaw, St. Louis, and Renfro (excluding Muldraugh equivalents) units
in Kentucky.
89° 88° 67° 86° 85° 84° 83° 82°
I | | I 7 A l l 1
33° _. E ("l \ ’ (KN (“\f—
g 3 , V \\ OHIO f T
i “o f .4 ,/ ~’
\_ _ , i
,3 1NMANA J “ z
/ ‘ pi
/ , ,
/ , ”a"
x L,”
f ‘
J "7’ WEST
5 Q magma
38° — ,5“ ~ k) _
,7 , (\
{LUNOIS 1

    

 

 

 

100 MILES

FIGURE 58.—Thickness (in feet) of St. Louis Limestone and St. Louis Member of Newman Limestone and Slade Formation. Top of unit
placed at top of Lost River Chert Bed of Elrod (1899) in south-central and west-central Kentucky and inferred equivalent in western
Kentucky.

SUMMARY OF THICKNESSES AND LITHOLOGIC TRENDS 97

 

 
  
  
 
 
  
  
 
   
 
 

   

88° 88° 87° 83° 85° 34° 83° 82°
I I I | I { ,. l l I
39” —— EXPLANATION ' ’ , / \—
Distribution of Warix Run Area of reported distribution / "
Member of Platycrinites penicillus and \
Distribution ofconglomeratic Lithostrotion harmodttes m .9
beds at base of Ste south-central Kentucky—p. :
Genevieve reported isolated occurrence ,2
§ A
rea of reported outcrop ‘, , 2:. ‘
‘\\\ occurrence of Lost River ﬂ ,\ r’ ..'> ” ~ . _ a 1.1:??? A
o Chert Bed in south-central A ’ * ' “ ‘ ° 3 U ‘ I“
39 ‘ Kentucky _

lLUNOIS

 

 

 

u ' so 100 MILES
l_i__l_l_l_L___—l

FIGURE 59.—Thickness (in feet) of Ste. Genevieve Limestone and probable and inferred equivalents (Warix Run and Ste. Genevieve Members
of Slade Formation and Newman Limestone, and Ste. Genevieve Limestone Member of Monteagle Limestone). Base of Ste. Genevieve
in west-central and western Kentucky placed at top of Lost River Chert Bed of Elrod (1899) or inferred equivalent.

89° 33° 37° 35° 85° 34° 33° 32°
l | | l l 1

 

 

39°- EXPLANATION .j‘ j ..... \ mm [/ij—
‘ 3 ' A .1 (

Shale Sandstone

 
 
 
  
  
 
 
  
   

 

Sandstone Limestone
and shale . , and dolomite
0 50 100 PERCENT
—0—— Approximate erosional edge
of Chester Series and
38° — equivalent rocks

{Llth'HS

 

 

 

0 50 100 MILES
l_1._\_|_L_l_—l

FIGURE (SQ—Thickness (in feet) and generalized lithofacies of Chesterian rocks and equivalents in Kentucky.

98 MISSISSIPPIAN ROCKS IN KENTUCKY

     
 
 
  

 

     
  
  

 

89° 88° 37° 95° 32.,
I I
390— ”W I l L—
I / \
I 1
/ INDIANA :3
/ x
“/1 XV;
r" j WEST
53 VIRGINIA
38°~
ILLINOIS
LN.
K
I} .350
/\
K» ,CP'u/Jgo‘x I ' .
37°— \ I ~~~~~ \
MO J .
5:!“ hf»~~m_mnnwh‘:_/I I TENNESSEE. I l I I I

 

 

 

0 50 100 MILES

 

FIGURE GIL—Thicknm (in feet) of Kidder Limestone Member of Monteagle Limestone and equivalents in eastern to south-central Kentucky,
and cumulative thickness of Paoli Limestone through Elwren Sandstone. and Renault Limestone through Cypress Sandstone in west-
eentral and western Kentucky.

89° 88° 37°

 

39°

   

i
I
/ INDIANA
/

ILLINOIS

 

 

 

[I 50 IUD MILES

I_I_I__I_I_L—__I

FIGURE (ta-Thickness (in feet) of combined Hartselle Formation and Bangor Limestone in eastern to sonar-central Kentucky and thickness
of Golconda Formation in west-central and western Kentucky.

SUMMARY OF THICKNESSES AND LITHOLOGIC TRENDS 99

89° 88° 87" 86° 84° 83° 82°

l |

   
  

 

39° _.

lNDlANA

38° —
lLLlNOiS

 
    

.._.z.-v «V -,

 

 

 

 

TENNESSéE
[] 50 100 MILES
I_I_I__I_.I__L______l

FIGURE 63.—Thickness (in feet) of Paragon and Pennington Formations in eastern to south-central Kentucky and combined thickness of
Glen Dean—Tar Springs—Vienna—Waltersburg formational units in west-central and western Kentucky. Dashed isopach lines indicate
approximate or inferred thickness. Area of anomalous thickness patterned.

 

  
 
 
      

  
   

   

 
      

 

 

 

89° 88° 87° 88° 34° 33° 82°
T | H I | l l l
39° _ \_
I OHIO \d
\ A /
I INDIANA KVN‘ ' kl
/ I’ N’TVT‘ H
/ ‘ J
Lz’J
K
, f WEST
\ VIRGINIA
38° —
ILLINOIS
5‘»
C
\
."j
7 \
\\/ (P/;::\O\.\
37° — \S\<
‘1 , £+
MO I .
k,_ -
(4
$:£[~J{:’MT,W;_ .I.. 1.4m I TENNESSEE I M I ‘ " ”I
u 50 mo MILES
L_I___L___L_.I___|_._____l

FIGURE 64.—Generalized thickness (in feet) of interval from base of Menard Limestone and equivalent units in Leitchfield and Buffalo Wellow
Formations to base of Caseyville Formation (Pennsylvanian).

100

at different times in their nomenclatural history. These
isopach maps were compiled by Sable, who has at-
tempted to use criteria consistent with surface interpre-
tation where possible. Because of the above factors,
general thickness trends shown on the maps are consid-
ered to be valid, but positions of isopachs should not
be considered to be exact. For instance, the known linear
trend in the thick, narrow Bethel (Mooretown) sandstone-
filled channel (see section, “'Ibrrigenous clastic units,”
p. 74), is not adequately shown in the corresponding
thickness map (fig. 60), which portrays the early
Chesterian interval within which the Moorean occurs
(recall fig. 8). In the following discussion relating to
figures 51—64, some reiteration of material previously
discussed is included for the sake of emphasis.

Factors affecting the total thickness of Mississippian
rocks (fig. 51) in Kentucky are post-Mississippian ero-
sion, variations in Mississippian sediment volume
distribution, and Mississippian tectonic movements.
Erosional effects include regional beveling after deposi-
tion of Chesterian rocks but prior to deposition of Penn-
sylvanian strata, and Mesozoic and Cenozoic erosion
that has stripped Mississippian rocks from the Pascola
arch in westernmost Kentucky and the Cincinnati arch
in central Kentucky. Thickness variations of Mississip-
pian deltaic deposits in parts of the region, shifts in loci
of detrital deposition, and filling of troughs peripheral
to deltaic platforms by carbonate and detrital rock
sediments appear to be less important factors than late
or post-Mississippian erosion in portrayal of present
total thicknesses.

In western Kentucky, Mississippian strata thicken
towards the Fairfield basin in Illinois to more than 2,800
ft in Crittenden County, western Kentucky. Deviations
in general trend directions suggest control by warping
along the Rough Creek—Shawneetown and Pennyrﬂe
fault zones. In eastern Kentucky, a broad, irregular
northeast-trending platform area is represented by
700—900 ft of Mississippian strata, east of which the
rocks thicken rapidly into the Appalachian basin. The
thickest strata east of the Cincinnati arch trend are on
the east side of the Cumberland saddle area in south-
central and southeastern Kentucky, and locally in
northeastern Kentucky east of the Waverly arch.

The elastic wedge of latest Devonian and earliest
Mississippian rocks in Kentucky, comprising the
Bedford-Berea-Sunbury units, is less than 200 ft thick.
Isopachs of these units (fig. 52) indicate the distal parts
of major deltaic deposits (Pepper and others, 1954), with
minor depocenters in northeastern Kentucky, and
southwest-trending lobes which are interpreted to
reﬂect deltaic lobes originating east of Kentucky. The
Pine Mountain fault does not appear to affect isopach
trends. '

 

MISSISSIPPIAN ROCKS IN KENTUCKY

In south-central and western Kentucky, elastic rocks
of Kinderhookian age (Maury Formation and Hannibal
Shale equivalents) are for the most part extremely thin
(ﬁg. 52), and their constitutents appear to represent lag
deposits resulting from winnowing of distal deltaic
deposits by bottom currents. Some areas in which Fort
Payne or New Providence strata rest directly on the
Chattanooga Shale, mostly in western Kentucky (see
section, “Kinderhookian-Osagean unit relationships,”
p. 35), may indicate areas of mild uplift scoured by
strong current action and swept clean of distal deltaic
fines.

Isopachs of the terrigenous clastic units in the
Borden, basal Fort Payne, New Providence and
Grainger Formations (fig. 53) indicate an irregular
depositional platform area in eastern Kentucky in which
axes of the thicker clastic lobes trend southwestward,
normal to the conspicuous northwest-trending Borden
delta front slope deposits that extend in an almost
straight-line trend from west-central through south-
central Kentucky. The Borden delta front and the mild-
ly irregular delta platform topography in eastern
Kentucky represent the last phase and ﬁnally cessation
of active Early Mississippian deltaic deposition in Ken-
tucky. Southwestward paleoslopes were dominant, and
eastern source areas were probably the important
contributors to the sediment wedge as shown by distri-
bution and transport indicators of some sandstones
(fig. 16).

West of the Borden delta front, thin distal green
shales are generally less than 20 ft thick, and locally
absent in western Kentucky (fig. 53). A northwest-
trending belt in which the carbonate and siliceous rocks
of the Fort Payne Formation strata directly overlie the
Chattanooga Shale occurs in the southern part of
western Kentucky subparallel to the Borden front, and
may indicate that a low submarine platform occupied
this area during deltaic deposition. Another area in
which basal clastics are absent is part of the ﬂuorspar
district of western Kentucky, which may have been a
dome before collapse and mineralization. Southwest of
this, however, a thickening of greenish shales inter-
preted here to be Osagean in age suggests the upper
end of a westward-deepening mud-filled trough.
Possibly the area in the ﬂuorspar district was a low plat-
form that acted as a barrier to sediments from the
Borden delta of Illinois (Willman and others, 1975,
p. 137) and the Borden delta deposits in Kentucky.

Carbonatedominated units, which form the Mul-
draugh Member of the Borden Formation and the bulk
of the Fort Payne Formation (fig. 54), thicken south-
westward parallel to the Borden delta front across west-
central and south-central Kentucky, and, farther west,
thicken westward and southwestward to more than

SUMMARY OF THICKNESSES AND LITHOLOGIC TRENDS

400 ft towards the Fairﬁeld basin of Illinois and to more
than 600 ft into the Mississippi Embayment. Litholo-
gies are mostly dark siliceous silty dolostones and silty
limestones except for two northwest-trending belts of
strikingly different strata recorded in subsurface sec-
tions: (1) a belt of crinoidal biocalcirudite with associ-
ated green shale which extends northwestward from the
outcrop area of Cane Valley Limestone Member of
south-central Kentucky, and (2) a dominantly dark gray
to brown shale and siltstone, in places more than 200 ft
thick, occupying most of the interval which is commonly
siliceous carbonate rocks. The shale belt is parallel to
the biocalcirudite—green shale belt, and on strike with
the Kniﬂey Sandstone Member of the Fort Payne For-
mation of south-central Kentucky. (See section, “Fort
Payne Formation,” p. 48.)

The biocalcirudite—green shale interval in the Fort
Payne Formation largely coincides with a northwest-
trending belt of thicker Fort Payne strata. The greater
thickness is inferred to be the result of uneven deposi-
tional bottom topography rather than filling of a tec-
tonically controlled depression. Spatial relationships
and directional trends of the crinoidal biocalcirudite and
green shale and the related Cane Valley Limestone
Member, the dark shale unit, the Knifley Sandstone
Member, and the Borden delta front indicate genetic
relationships of these features. The biocalcirudites and
green shale formed southwest of and adjoining the
Kniﬂey Sandstone Member and the dark shale unit. The

dark shale unit may be equivalent to the Knifley, a.

distal fine sediment winnowed from the Kniﬂey sands
by longshore currents and deposited to the northwest
by them. The crinoidal limestones and the green mud-
stones, rather than being a slope edge deposit derived
from an area of delta platform origin as interpreted by
Sedimentation Seminar (1972), may have been sub-
marine bank deposits growing in a peridelta belt prox-‘
imal to the deltaic bottom slope deposits and swept by
currents that maintained a nutrient supply, but from
which muds were trapped by the in-place crinoid organ-
isms. Green coloration of the muds may indicate
accumulation in a mildly reducing, open circulation envi-
ronment. Basinwards, darker sediments accumulated
more slowly perhaps in a deeper anoxic environment,
as indicated by the tabular bedding and fine-grained to
aphanitic character of much of the Fort Payne strata.

In figure 55, cumulative thicknesses are shown for all

pre-St. Louis carbonate-dominated units of Osagean and .

Early Meramecian age, including the Harrodsburg.
Salem, and Warsaw Limestones. In this and in the
Harrodsburg Limestone thickness map (fig. 56),
westward-thickening strata culminate in a west- to
northwest-trending trough trending toward Fairfield
basin of Illinois. The 300-ft isopach on figure 55

 

101

encompasses a broad area in west-central and south-
central Kentucky, roughly reciprocal with the thickness
trends shown for the Muldraugh—Fort Payne carbonate
rock thicknesses in figure 54.

Thicknesses of the Renfro Member in southeastern
Kentucky (fig. 55) are derived mostly from subsurface
information, and correspond reciprocally with the thick-
nesses of the underlying Borden Formation clastic
strata of figure 53. This type of relationship is con-
sidered to result from filling of underlying irregular
depositional topography by the overlying carbonates.
However, the degree of thickening into the trough in
western Kentucky is not related to such causes, but in-
dicates filling of a structural trough which either
predated carbonate deposition, existing under starved
basin conditions, or. more likely, which developed in late
Osagean and early Meramecian time during carbonate
sediment deposition.

Rocks of mostly Meramecian age (ﬁgs. 56—59) are best
developed in parts of eastern and western Kentucky.
Pre-Pennsylvanian and intra-Mississippian erosion has
removed much strata from former areas of deposition
in parts of northeastern Kentucky. PreCretaceous ero-
sion has removed very thick sections of rocks in
westernmost Kentucky and southernmost Illinois.
Thicknesses of Meramecian rocks (Sable, 1979a) range
from 400 to more than 1,100 ft. The thickest accumula-
tion of sediment was in western Kentucky, southern
Illinois, and southwestern Indiana. Marginal to the Cin-
cinnati arch in eastern Kentucky, thicknesses are gener-
ally less than 200 ft.

Strata of Meramecian age are predominantly car-
bonate rocks throughout the State. The bordering areas
in which terrigenous detrital components are concen-
trated are those in which the upper, dominantly lime-
stone units, such as the St. Louis and Ste. Genevieve,
have been eroded; in south-central Kentucky (Cumber-
land saddle area) and western central Kentucky, fine
detritals in the lower units such as the Salem and
Warsaw make up a significant proportion of the strata
that remain. Mudstone is also present in these units on
the west side of the Cincinnati arch from southern In-
diana to the Nashville dome in Tennessee. Quartz sand
grains are present in the Ste. Genevieve Member of the
Newman (Slade) and Warix Run Member of the Slade
in northeastern Kentucky and their equivalents in
southern Ohio and West Virginia, and farther south in
eastern Kentucky.

The thicknesses of the Harrodsburg—Warsaw Lime-
stone interval in western Kentucky, more than 300—500
ft in Trigg and Crittenden Counties (fig. 56), are inter-
preted to reﬂect filling of troughs peripheral to banks
formed on earlier Fort Payne deposits. (See Dever and
McGrain, 1969, fig. 6.) Westerly transport directions of

102

carbonate debris and westward thickening into the Fair-
field basin in Illinois are indicated. Of interest is the con-
siderable volume of clean, winnowed and sorted
bryozoan- and crinoid-bearing debris interpreted to
originate from shoals on the surface of the older Borden
delta and Fort Payne accumulations. The regional strike
of the Harrodsburg Limestone in Kentucky is nearly
at right angles to strike of the earlier Borden delta and
to the strike of the Borden and Fort Payne carbonate
units of figure 54. This seems to reﬂect general filling
of peridelta areas and the basin in southern Illinois.

In Illinois, the St. Louis and the Salem Limestones
have generally been considered to be intertonguing
units so that their boundary cannot be considered to
approximate a regional time rock boundary (Lineback,
1972). The ﬁeld relationships and reciprocal thicknesses
of the St. Louis and Renfro indicated by comparisons
of quadrangle maps do not reﬂect the intergrading and
interlensing of the two units as much as they do incon-
sistent mapping of the contact. In figure 57, the upper
limit of the St. Louis is arbitrarily placed at the top of
the Lost River Chert Bed in south-central, central, and
west-central Kentucky and at the top of inferred equiva-
lents in western Kentucky. The resulting thickness map
shows a gradual thickness increase towards the Fair-
field basin of Illinois; the regional strike of the isopachs
swings from north-northwest east of the Cincinnati arch
to northeast and east-northeast in western Kentucky,
reﬂecting rather even basin filling and indicating that
the irregular bottom topography which characterized
the region during Borden deposition had disappeared
probably by Salem or early St. Louis time.

Thickness of the St. Louis Limestone alone in Ken-
tucky (fig. 58), shows a rather even basinward thicken-
ing similar to that discussed above. Difficulties of
tracing both the base and top of this unit in western
Kentucky, however, make thickness interpretations
there uncertain.

The Ste. Genevieve Limestone and its equivalents
represent largely shoal water deposition over a wide
area. Isopachs of this unit and its presumed equivalents
in eastern and southeastern Kentucky, the Ste.
Genevieve Member of the Monteagle and Newman
Limestones and Slade Formation (fig. 59), show an ap-
parent westward—thickening trend across the Cincinnati
arch and development of a distinct basin in western
Kentucky south of the Moorman syncline. There thick-
nesses exceed 300 ft. Thicknesses in southeastern Ken-
tucky are uncertain but appear to be generally about
50—7 5 ft, except for a few estimates along Pine Moun-
tain, where strata in the Newman Limestone which are
lithologically like the Ste. Genvevieve and contain
Platycn'nites penicillus are as much as about 100 ft
thick.

 

MISSISSIPPIAN ROCKS IN KENTUCKY

All originally deposited thicknesses of Chesterian age
rocks (fig. 60) in Kentucky have been modified by post-
Mississippian erosion. The youngest known Mississip-
pian units, the Kinkaid Limestone and Grove Church
Shale, occur in the southern part of the basin in western
Kentucky. There, nearly 1,200 ft of Chesterian Series
rocks are recorded, but the original upper surface of the
Mississippian deposits cannot be confidently restored.
Within Chesterian units that are most widespread, how-
ever, the Glen Dean Limestone and older Chesterian
strata thin towards the margins of the Eastern Interior
basin; this thinning is considered to be depositional thin-
ning towards uplifted belts.

Strata of the Chesterian Series and equivalents are
more than 1,100 ft thick in both eastern and western
Kentucky. The strata are thin or absent, however, in
northeastern Kentucky in the Waverly arch area and
along the margins of the Cincinnati arch. Irregular
thicknesses along the east side of the Cincinnati arch
are largely the result of irregular original thickness dif-
ferences as affected by positive tectonic and irregular
topographic features. The latter feature resulted from
earlier intra-Mississippian erosion and to a lesser extent,
from Late Mississippian or pre-Pennsylvanian erosion
prior to deposition of the Lee and Breathitt Formations.
In western Kentucky, basinal outlines are shown along
the Moorman syncline and subparallel to the faults of
the Pennyrile fault system.

Thicknesses of intra-Chesterian intervals (ﬁgs. 61-64)
show the continuing basinal character of western Ken-
tucky, but do not definitively show the Cincinnati arch
trend. A depocenter in or near the Cumberland saddle
area is indicated in ﬁgures 61 and 62, as is the inﬂuence
of the Waverly arch either during or following deposi-
tion of the Paragon Formation (fig. 63). Thickening of
sediments towards the present Pascola arch in early and
middle Chesterian time (ﬁgs. 61—63) indicates that the
arch area was still negative, but isopachs of latest
Chesterian strata (fig. 64) show an east-west-trending
basin along the Pennyrile fault system, suggesting that
the arch began to grow before the Mississippian-
Pennsylvanian erosional interval began. In the eastern
part of Kentucky, isopachs of rocks show a northeast-
trending belt of variable sediment thickness ranging
from 150 to more than 200 ft which, taken with the
dominant carbonate lithofacies, seems to reﬂect a trend
of carbonate banks fringing the Appalachian basin
(fig. 60).

In summary, the selected thickness maps of the
Mississippian and maps showing lithologic trends
(Craig and Connor, 1979) reﬂect the onset of deltaic
deposition in eastern Kentucky during Late Devonian
and Early Mississippian time followed by the main
southwestward progradation of the Borden delta. A

CYCLIC DEPOSITION

trough or depocenter was maintained in extreme south-
western Kentucky during Kinderhookianl?) and early
Osagean time. This shifted northward in late Osagean
and early Meramecian time, and concurrently a
southwest-dipping paleoslope developed. Depositional
infilling west of the present Cincinnati arch appears to
have slowed by middle Meramecian (middle and upper
St. Louis) time. In western Kentucky a shallow deposi-
tional basin developed, the depocenters of which appear
to have shifted throughout Chesterian time. This shift-
ing may have been in part due to tectonic causes such
as sags along the Pennyrile fault system, and in part
to depositional deltaic depocenters, depending on
whether tectonism or sediment load was a more impor-
tant factor.

In eastern Kentucky, history of downsinking of the
northeast-trending Appalachian basin is not well docu-
mented in this report because reliable subsurface infor-
mation is scarce. Early Mississippian strata of the
Borden show little indication of strong northeast linear
trough development. Rocks considered equivalent to the
St. Louis, Ste. Genevieve, and Chesterian strata below
the Paragon and Pennington likewise show no appreci-
able basinal thickening in Kentucky, but Pennington
strata do show a decided southeastward thickening,
which may reﬂect both tectonic downsinking and
loading response to Pennington sediments.

The role of the Cincinnati arch in Kentucky during
the Mississippian is not clear. Judging from the isopach
thicknesses and trends, the arch seems to have been
rather quiescent or only mildly active until late
Chesterian time. The Waverly arch also does not seem
to have exerted control of sediment thicknesses until
post-St. Louis time. Between Osagean and Chesterian
time, the area of the Cincinnati arch in Kentucky may
have been a broad, irregular platform downwarped to
the west, with vague northeast- to north-trending
hingelines west of the present arch axis. Subsequent
growth of the arch is shown by the apparent increased
rates of thinning of St. Louis, Ste. Genevieve, and
Chesterian age strata towards the arch in south-central
Kentucky. The average thicknesses of St. Louis beds
west and east of the present Cincinnati arch are about
229 and 166 ft respectively, with a thickness ratio
west:east of 1.4. Beds of the Ste. Genevieve average
about 166 and 68 ft thick respectively, with a ratio
west:east of 2.4. Depositional rates thus appear to have
increased from St. Louis through Ste. Genevieve time
within Kentucky, with areas to the west demonstrating
a greater relative amount of downsinking during this
interval. During Chesterian time and the post-
Mississippian—pre-Pennsylvanian interval, the arch
seems to have been well developed, as indicated by
the Bethel (Mooretown) channel remnants which cut

 

103

progressively older Mississippian strata northeastward,
and by the basal Pennsylvanian channel remnants in
central Kentucky, which also cut progressively older
Mississippian strata eastward.

CYCLIC DEPOSITION

Two major pulses of terrigenous detrital sediments
during Mississippian time followed deposition of the
Sunbury Shale, which overlies the shallow- to deep-
water deltaic Bedford-Berea units. These terrigenous
pulses resulted in the Borden-Grainger offshore delta
units and the rhythmic Chesterian lower delta plain
deposits. The two pulses were separated by a wide-
spread shallow-water marine transgression in which ex-
tensive carbonate rocks as well as minor evaporites were
deposited. Transitional units between the earlier deltaic
deposits and the carbonate-dominated succession are
the Muldraugh and Renfro Members of the Borden For-
mation and siliceous carbonate rocks of the Fort Payne
Formation.

The Devonian-Mississippian Bedford-Berea-Sunbury
units become finer grained upwards, but in the eastern
United States, including most of Kentucky, conditions
favorable to carbonate sedimentation did not occur until
after the pulse of clastics of the Borden-Grainger suc-
cession. However, the organic, low-energy Sunbury
Shale is probably the eastern time equivalent of exten-
sive carbonate rocks in Indiana, Illinois, Missouri, and
Iowa, and thus may represent the late stage of a
regional broad cycle of deposition. Cessation of Borden
deltaic deposition and subsequent infilling by the
upward-fining carbonate succession culminating in the
St. Louis carbonate and evaporite strata may represent
a second broad cycle. The Ste. Genevieve does not in
itself seem to fit neatly into a cyclic framework, but
thinner rhythmic upward-fining units are reported
within the Ste. Genevieve (Sandberg and Bowles, 1965)
and are discussed herein.

St. Louis lithologies reﬂect very low energy environ-
ments and probably represent protected lagoonal and
shallow bay conditions over a large area. Internal
cyclicity is reﬂected by upward-fining successions gen-
erally beginning by a pulse of fine terrigenous debris,
green shales such as those reported in quadrangles in
west-central Kentucky (Kepferle, 1963b; Withington
and Sable, 1969), and associated fine-grained calcar-
enites and calcsiltites with scattered fossils which grade
upward into calcilutite limestone. Because the St. Louis
is very poorly exposed, little evidence for cyclical repeti-
tions is recorded from surface exposures. However, in
exploratory well cores, Kepferle and Peterson (1964)
recorded upward alternations of oolitic crossbedded

104

limestone to fine-grained dolomitic limestone and
dolomite to fine-grained micritic limestone capped by
cherty micritic limestone. (Also see Fox and Seeland,
1964; Seeland, 1968.)

Examples of cyclicity in Ste. Genevieve strata prior
to the rhythmic repetition of Chesterian carbonate and
clastic strata are discrete units of widespread sandstone
and green shale in the upper part of this otherwise
carbonate-dominated unit. Shales which may be Spar
Mountain Sandstone equivalents and units referred to
the Aux Vases Sandstone (Amos, 1971, 1972) and
Rosiclare Sandstone Member (Amos, 1965) are ex-
amples of periodic interruption of the dominantly
carbonate depositional regime. These shales in them-
selves do not necessarily support a hypothesis of
cyclicity, because they may simply be the results of
lateral migration of deltaic channels and channel splays
in a general downsinking region. However, in south-
central and the southern part of west-central Kentucky,
the Ste. Genevieve Limestone and the overlying Girkin
Limestone display rhythmic alternations of carbonate
rocks. Units of crossbedded sandy or oolitic limestone
are overlain by fossil fragmental and micritic limestone
and capped by cherty limestone (Sandberg and Bowles,
1965). Upward-coarsening cycles are also recognized in
the Girkin (Dever and Moody, 1979b). Columnar sec-
tions in other quadrangles (Sable, 1964; Ulrich, 1966)
suggest similar cyclicity in which crossbedded oolitic
or sandy limestone and green shale represent repetitive
breaks in otherwise similar-appearing sections. In the
lower part of the Ste. Genevieve, a succession of alter-
nating dolomite and oolitic limestone 30 ft thick has
been recognized in Breckinridge, Hardin, Hart, and
Warren Counties. As more detailed information
becomes available, intraformational cyclic successions
such as the preceding may become valuable correlation
indicators.

In east-central and northeastern Kentucky, columnar
sections of considerable detail (Delaney and Englund,
1973; Gualtieri, 1967b; McFarlan and Walker, 1956, pl.
2) show lithologic repetition which has been ascribed to
relatively local tectonic movements of the Waverly arch
and along the Kentucky River fault system (Dever and
others, 1977). In these sections, thin units of subtidal
oolitic and sandy limestone, irregularly bedded micritic
limestone, dolomitic limestone, cherty limestone and
green shale, or a part of this succession are broken by
breccias with dark micritic matrix interpreted to repre-
sent subaerial exposure or vadose zones (Dever and
others, 1977). Associated structures include birdseye
texture and contorted piercement “tepee” structures
that attest to repeated submergence and emergence
during Chesterian and late Meramecian time. Inter-
preted from a tectonic viewpoint, the above evidence

 

MISSISSIPPIAN ROCKS IN KENTUCKY

might indicate repeated oscillations of the source area
as well as tectonic- or otherwise-induced sea-level
movements within the depositional area.

Causes for the apparent cyclical sedimentation in the
Chesterian are unresolved. J .M. Weller (1956) suggested
that tectonism was chieﬂy responsible for late Paleozoic
cyclicity. Swarm (1964, p. 656—657) considered climatic
ﬂuctuations, primarily changes in rainfall in the source
region of the Michigan river combined with even basin
subsidence, to be the chief factor. He (Swann, 1964,
p. 654) also cited eustatic sea-level changes as possible
factors in some cycles in the upper Chesterian series,
and suggested that alternation between pluvial and in-
terpluvial stages corresponded to ﬂuctuations of the
margins of continental glaciers in the southern hemi-
sphere. About 13 world-wide synchronous depositional
sequences of Kinderhookian through Chesterian ages
resulting from eustatic sea-level changes were postu-
lated by Ross and Ross (1985). Major regressions
include those in early Kinderhookian, early and middle
Meramecian (St. Louis evaporites), and latest
Chesterian times. They cited possible mechanisms for
the eustatic changes to include oceanic crust volume
changes, ocean trench activity, and orogenic activity
along continental margins. “Jostling” of continental
plates during a mid-Paleozoic collision of northeastern
North America and Europe (Craig and Varnes, 1979)
may have periodically uplifted source areas and may be
a realistic tectonic concept to account for apparent
cyclicity.

PALEONTOLOGY

During the geologic mapping program, fossils were
used as practical mapping aids during quadrangle map-
ping. Paleontologists of the US. Geological Survey
assisted in field identifications of genera and species of
macrofossils, and in some instances laboratory deter-
minations of fauna and ﬂora. An excellent summary of
the paleontological zonation of Mississippian rocks in
the United States (Dutro and others, 1979) included
discussions of foraminiferal, brachiopod, ammonoid,
and conodont zonation.

MACROFOSSILS

Taxonomy and zonation of macrofossils in Mississip-
pian rocks stem from extensive early studies in Ken-
tucky and adjacent States by S. Weller (1920, 1926),
J.M. Weller (1931), Butts (1915, 1917, 1922), and ED.
Ulrich (1917), and others, reviewed and updated by
Weller and Sutton (1940). Crinoids, brachiopods,

PALEONTOLOGY

blastoids, bryozoans, solitary and colonial corals, and
echinoids are dominant forms in the Mississippian
assemblage; pelecypods, gastropods, and trilobites are
locally abundant. Crinoid studies by Horowitz (1965),
and biofacies studies of a Chesterian rock unit (Vincent,
1975) are two examples of the many selective studies
done in recent years.

One significant break in the crinoid fauna marks the
Meramecian-Chesterian Series boundary in western,
west-central, central and south-central Kentucky—that
between strata containing Platycn'nites penicillus Meek
and Worthen, a Meramecian form, and strata contain-
ing Chesterian forms of Talarocrinus spp. The faunal
change corresponds to the time of formation of a wide-
spread zone of altered limestone and breccia (Bryants-
ville Breccia Bed) interpreted to have developed during
subaerial exposure and diagenesis. Other series bound-
aries, with the possible exception of the Kinderhookian-
Osagean boundary, appear to occur within intervals of
continuously deposited strata, and faunal criteria for
specific boundary demarcation are not conclusive.
However, many specific and generic forms have proved
valuable aids in practical recognition and mapping of
stratigraphic units. Figure 65 shows general strati-
graphic occurrences of selected fossil faunal elements
which have been helpful in discriminating Mississippian
rock units in Kentucky.

Macrofossils of Late Devdnian and Kinderhookian
ages are scarce in Kentucky. Sparse brachiopod and
molluscan faunas in the New Albany, Chattanooga,
Ohio, Bedford, and Sunbury Shales present difficulties
in determining the systemic boundary (Campbell, 1946).
Studies of macroﬂoras by Cross and Hoskins (1951,
1952) in regard to the Devonian-Mississippian systemic
boundary were also generally indeterminate and in some
cases contradictory to age determinations of macro-
fossil assemblages. Savage and Sutton (1931) concluded
that the lower part of the black shale (New Albany
(Chattanooga)) in Allen County, Ky., is Devonian and
the upper part Mississippian.

Macrofossils in units of mostly Osagean age, the
Borden, Fort Payne, Grainger and lower part of the
Harrodsburg, are generally uncommon and scattered
in terrigenously derived clastic strata. Fossil remains
consist largely of scattered crinoid columnal frag-
ments and brachiopod shells, with minor bryozoan
fragments. Locally, concentrations of a more varied,
well-preserved fauna occur in thin limestone lenses in
clay shale of the New Providence Member of the
Borden, such as the localities at Buttonmold Knob,
Bullitt County, and Kenwood Hill, Jefferson County
(Butts, 1915, 1917; Conkin, 1957). The New Providence
fauna has most recently been studied by Kammer

 

105

(1982). In northeastern Kentucky, in Stricklett and
Head of Grassy quadrangles, large spiriferoid
brachiopods, crinoids, and bryozoans are locally
abundant in the Cowbell Member of the Borden (Morris,
1965b, 1966a). Chaplin (1980) reported on fossil con-
centrations in the Nancy, Cowbell, and Nada Members
of the Borden. Mason (1979) and Mason and Chaplin
(197 9) also described cephalopod faunas in the Farmers,
Nancy, and Cowbell Members; and ammonoid faunas
indicating westward progradation of the Borden in
Kentucky are discussed by Gordon and Mason (1985).
In the Borden, Fort Payne, and lower part of the
Harrodsburg, biostromal accumulations of crinoid
and bryozoan debris occur at various horizons. They
include the Cane Valley Limestone Member and Beaver
Creek limestone member which are thick and areally
extensive in south-central Kentucky (Thaden and
others, 1961; Sedimentation Seminar, 1972; Kepferle
and Lewis, 1974). Biostromal lenses in which crinoid
and other echinoderm fragments are common are
also reported in the Fort Payne in western Kentucky
(Rogers, 1963; Hays, 1964; Fox and Olive, 1966).

Macrofossils of Meramecian age, compared to the
earlier Mississippian occurrences, are abundant and
varied, particularly in the Salem, the Warsaw, the Salem
and Warsaw map unit, and the Harrodsburg Limestone.
Brachythyrid, spiriferoid, and productid brachiopods,
solitary corals, echinoid spine and test fragments,
fenestrate bryozoans, and ubiquitous crinoid columnals
form calcarenitic, calciruditic, and coquinoid beds.
Fossils are best preserved in limy shales (Lewis, 1971a;
Kepferle, 1967) and are generally more fragmental in
the coquinoid and marly fine-grained limestones and
dolomitic limestones. Fewer forms are found in the St.
Louis Limestone, which is largely micritic, containing
scattered colonial corals, brachiopods, and echinoids;
fenestrate bryozoans are locally common, as are
crinoidal calcarenite and calcirudite. Faunal elements
in the Ste. Genevieve are more abundant and varied
than those in the St. Louis, and shelly coquinoid beds
occur particularly in the upper part of the unit. Many
of these are crinoidal biorudites and biocalcarenites in
which stem segments and calyx bases of the crinoid
Platycrinites are abundant.

Formations of Chesterian age in western to south-
central Kentucky, consisting of rhythmically alter-
nating units of limestones and detrital rocks, have
abundant macrofossils in the carbonate rocks, and scat-
tered remains in the clastic strata. Some carbonate units
can be identified with a good degree of assurance by
recognition of species of brachiopods, crinoids, and
bryozoans and to a lesser extent of corals and
pelecypods. Limy shales are locally very fossiliferous;

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

106 MISSISSIPPIAN ROCKS IN KENTUCKY
{’3 SOUTH-CENTRAL
E WESTERN WEST-CENTRAL AND
U) E AST- CE NTRAL
1% 3 5% m
5 g c Splrifer lncrebescens Hall ﬁg 5 Splrlfer lncrebescens Hall 5
2" 9 3 .9“ .. 5
,5: g Composlta subquadrata Hall 2E E Composlla subquadrula Hall 27%
- c m .. ._
3 “’3 E138 E E Pterotocrlnus spp.
% Sulcatoplnna mlssourlensls m 5 £8
E 2 c 2 5% A m d s *
a % Pterolocrinus spp.” g 9 Pterolocrlnus spp.‘ 2’73? Tc me es pp.
5 E c 3 Prismopora serrulala Ulrich 5.3
3 G: Archimedes Spa” 5 E Archimedes spp.“ m]
E
l3
3
.1 i
a 3 5 3 5.5 Archimedes spp.
g? C E g
-5 E g; 2 Large crinoid stem segments“
0.2 m a: Inflalla lnflala (McChesneyv
H“! H l l 2 Pentremltes spp.“
' 2 Agasslzocrlnus spp.
H .2 3 A . , s:
§§ g g, g .6 gasslzocr nus spp. .E
5 5 9 Talarocrlnus spp. £1, E Lithodromus veryl (Greene) '3 2
an: "’ . C . '6: 9,
gt 3 is: Talarocrmus spp. 3 ‘9 Talarocrinus spp.
0- 5: °_ :2 E E
or: o ._
___________ _9_>__. __ _._____.______ 2 -‘ _________________
a: Schoenophyllum aggregatum Eﬁnopmlﬁn agregalum :
> "thhostrotlon" (Slphonodendron) g "Lllhostrotlon"($lp oncdendmn) a Platycrlnltes penlclllus Meek and
o o
‘5 g geneuleuensls .5 2 geneuleuensls a Worthen
w ._. a) 2 Z
5 g Plalycrlnlles peniclllus Meek and E, 8 Platycrlnlles penlclllus Meek and
GE Worthen (a E
. ,_ Worthen
E! " 05 —‘
w G
Acrocyathus proliferum Acrocyathus prollferum — ”thhostrotlon" proliferum Hall
a "thhostrotlon"prollferum Hall” "Lllhoslrotlon"prolllerum Hall” 0
"a; : Acrocyathus florlfarmls floriformls g E .2 : Lithostrotlonella castelnaul Hayasaka
g .3 thhoslrotlonella caslelnaul a .9 Acrocyalhus Ilorllormis ﬂorlformls E
E j E HaYasaka. —l g thhostrotlonella caste/naul .1
—— ._ é _ Hayasaka’ u:-
3 ”’4 Melonechlnus sp. (n_’ Melonechlnus sp. SWMSOPOTU 3P-
: Syrlngopora sp. Archeocldarls sp.
é Syrlngopora Sp.
LU a) Endothyra balleyl Hall
2 E 8 Endothyra balleyl Hall
(is? Hapslphyllum sp. a E
V) c
- E .9 la "'
3 Brachythyrls subcardilformis (Hall) g g Brachythyris subcardllformlslHaII) ; g
(0.5 H . h ll 5 'g g Hapslphyllum sp.
Echlnocrlnus Sp.“ '1 apsq) y um p. g E
2 Splrlfer Iateralls %_I
In
E % Penlremlles conoldeus
'- a:
gag Talarocrlnus sp. g 5% Marglnlrugus magnus
_l '- 0-1
‘33 g —7 — Splrlfer Iateralis
—?‘— :I: 3
g C
3-2 Ortholetes keokuk (Hall)
a. g 0 : Very large crinoid stern segments“
2 t 5 S rln olh "'5 text slHaIl) £2
u H
:5 8m y g y é‘ig
c ..
8 g a) 8 g Very large crinoid stem segments' ”-
>3 u.
2.:
0.0)
g I
a:
Z

 

 

 

 

 

 

 

 

 

 

 

 

 

 

FIGURE 65.—Stratigraphic occurrences of Mississippian fossil fauna, selected on basis of abundance and ease of identification,
that are helpful in recognition of map units in Kentucky. Asterisk following name denotes very abundant forms. Lined areas
denote intervals of largely terrigenous detrital strata.

PALEONTOLOGY

shales in the Glen Dean Limestone and younger units
are notable examples. Brachiopods in sandstones are
locally common but mostly rare and scattered in such
units as the Big Clifty and Bethel (Mooretown).

In east-central and northeastern Kentucky and along
Pine Mountain, the lower Chesterian equivalents below
the Cave Branch Bed are generally less fossiliferous
than their western counterparts. Units in the upper
parts of the section such as the Bangor Limestone and
its equivalents and limy lenses in the Pennington For-
mation contain a more abundant macrofossil fauna. On
Pine Mountain, Chesterian equivalents are considerably
more fossiliferous than Meramecian units. In general,
probable supratidal conditions and hypersalinity con-
tributed to the dearth of fossil remains in the lower part
of the Chesterian equivalent section; rapid clastic
deposition and high turbidity probably inhibited the
growth and abundance of forms now found in most of
the Pennington Formation rocks.

MICROFOSSILS

The microfossils in Mississippian rocks of Kentucky
include conodonts, endothyrid and paleotextulariid
foraminifers, and ostracodes. Although no systematic
studies of the entire system have been done in Ken-
tucky, foraminiferal studies include those by Browne
and Pohl (1973), Browne and others (1977), Conkin
(1954, 1956, 1960, 1961), Pohl and others (1968), Pohl
(1970), and Pohl and Philley (1971). Conodont studies,
following zonation used in the Mississippi valley (Col-
linson and others, 1962, 1971), include those by Rex-
road (1958, 1969), Nicoll and Rexroad (1975), Rexroad
and Liebe (1962), Horowitz and Rexroad (1972), and
Chaplin and Mason (1979).

Conodonts currently seem to be the most reliable
faunal elements used in separating Devonian and
Mississippian marine rocks (see section “Devonian-
Mississippian systemic boundary,” p. 32) and for zona-
tion of Mississippian marine strata. Zones and ranges
of conodont assemblages have been established in the
upper Mississippi River valley (Collinson and others,
1962; Rexroad and Scott, 1964, reviewed in Collinson
and others, 1971) (fig. 66). Their use in delineating the
Devonian-Mississippian system boundary in the East-
ern Interior basin has also been discussed by Sable
(1979a) and in the Appalachian basin by de Witt and
McGrew (1979) and Sandberg (1981). Locations of
conodont collections from Upper Devonian and Lower
Mississippian rocks in Kentucky are shown in fig-
ure 67; the collections are discussed below and in
tables 3-6.

 

107

CONODONT IDENTIFICATION S

Samples of Late Devonian and Early Mississippian
ages were collected at scattered localities in east-central,
south-central, and central Kentucky during the geologic
mapping program. They were identiﬁed and commented
on by J .W. Huddle. His identifications and comments
are included in tables 3—6, arranged on the basis of
stratigraphic position and geographic area of the
samples collected.

Huddle (written commun.) commented in 1967:

The interpretation of these collections [Samples USGS 22787-PC
through 22794-PC] and those in KG-65—43 is not certain"*. Most
of the forms that were starred as possibly reworked are heavy forms
that could stand transportation or remain in lag concentrates. It seems
certain that the deposition in late New Albany and early New Prov-
idence time was very slow and gaps may be present, but the presence
of phosphate nodules ﬁsh remains and conodonts representing several
zones suggests that the area was under water and the gaps are due
to non-deposition and submarine erosion. The absence of the cono-
dont zones in the lower part of the New Providence in southern In-
diana and north-central Kentucky [central Kentucky of this report]
was determined by Rexroad and Scott (1964).*** The sequence [up-
permost New Albany Shale and lowermost Borden Formation] seems
to be most complete to the southeast, in a diagonal tier of quadrangles
from Burkesville to Panola. The Bactrognathus-Taphrognathus zone
is at or near the base of the New Providence in the Lebanon Junction,
Howardstown and New Haven quadrangles and the Bactrognathus-
Polygnathus communi zone is present in the Shepherdsville quad-
rangle. The older Mississippian zones in these quadrangles in north-
central Kentucky are thin or absent.

FLORA

Fossil plant remains occur as replacements by iron
and manganese oxides and silica, and as carbonized
woody fragments, tree trunks (some in growth at-
titudes) in sandstones such as the Big Clifty, and
macerated plant remains in carbonaceous shale and thin
coaly beds in the Bethel (Mooretown), Big Clifty, and
younger units. Spores occur in the coaly parts of these
Chesterian units and in the Pennington Formation (Et-
tensohn and Peppers, 1979). The zonation of plant
microfossils in western Kentucky and in adjacent States
Eastern Interior basin strata has established criteria
for distinguishing Mississippian terrigenously derived
clastics from lithologically similar Pennsylvanian strata
(Jennings, 1977).

Rare lycopod occurrences in the basal part of the St.
Louis Limestone of central Kentucky were reported by
Browne and Bryant (1970) and from the Salem and
Warsaw unit in south-central Kentucky by Dever and
Moody (1979b). Concentrations of plant debris are
reported in this part of the Mississippian succession in
the Elizabethtown quadrangle (R.C. Kepferle, oral
commun, 1965) and in the Rock Haven quadrangle
(Sable, oral commun., 1968).

MISSISSIPPIAN ROCKS IN KENTUCKY

108

.253 .323 6.8 "82500 Bod Ramayana 8in< Auto Z 5 83am ginmmmmmﬂﬁ 23 new mnoﬁuowmmmﬁu 3:8 235:8 3.38 we :Omﬁmmaooldw 55on

 

.I :uxuox .miszx manoéczmoaoi

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

.xﬁm :20. 3:2 02
um». 88:; Stucco—Em
M 38:56 SEE—Bani. 58:sz .w._m‘.33:um u=ubozocgm
w. 39:51
W. 2235.5 .mruESEPiug c=owocoﬁ5 5235.6 .9832 233:0...»5
w 88:53.33: .m 2238 tumoou .m.
32:25 2303 .mrcauamon u:uuo:ocsw I I analooﬂuaoolqwl I #3353 “3.0520
Ban: 1308 .m. I ”35955 3305320
conga—2 032.5535 maﬁuzméonouammtwAu‘EEum 33053.0 I I “$84th m I .33: 2380 .m.
2.5.59 2:35.39 n=£._u:u3oktwmﬂm.5ua m
:20 Son. ”EsEEoo maﬁuzmaﬁlﬁsﬁuzmohuum I 33:353.: .Q maﬁuzméoaovamﬂ
O was 35252.3 .maﬂucmoauum
s coauﬁtsm 2:331 .m u .50 .656 a: “a:
m mzﬁucmoinuhhaﬁuzmozuum nhELObuzsu .0 u Wantoumiumw‘ O
w 2.383 336530
A .
W 33—03. mzﬁuzuoinoh 3:83 330530 maﬁuzmoiuuhazcuxw. 2.30530
9
L1 M 3833
w «co—mm maﬁunmaun<ézutu= uncguzm2£auh 8mm 5 250:8 .25 53:55.
m :8: 83:5,:
m v.30.— .«w uaﬁutmgauoéazﬂaum maﬁa—5923‘
a
m. o>£>o=m0 .Sw
U
.322“. 3896.3 mzﬁuzmgaunvaaumsg m‘ioﬁunb
tam 9:62.00 I.| c
_ ..
g .
32:0
.320 :83 v.33 usﬁutmmzauoﬁaumsza 26053.0
m>___m.u_
>95:
92925.5: 8:3 02
W :80 :20 :xuﬁ maﬁucmowuiﬁauuﬁzn 33.230
% 35:5 .5...
m a: o_>
w mhznmhmzw; 3.5:» 352.30qu
3952
2:35;
9.20 2:326: maﬁutmgauszfuzmoqu
Ecomcn.
>§>>§§3£®Smﬁﬁ > ( £:;:g££
amtom ncoumczou. >m=~> Eaﬁﬂmﬂé Amom: 39.on can .zoom .com:___oo 8mm: .wumnmg

 

109

PALEONTOLOGY

.hxoaunovm Quanwoéuaom v5 €550-38 .1550 E 833% mﬁaoa.a:cvc:8 §Eo>oﬂ v.8 ﬁﬁﬁmmmmmg we mﬂmnwavaaw wﬁkoam 9:2le manure

 

<~Z.O~=>

  

SNNNémmNm

 

 

 

 

 

 

 

 

 

 

        
  

NmémnNN g

_ _ _
mmmmuzzmﬁ

 
 

  

E: 8:20....) :m
:oF8> «552 22

.2280 etc 8

3_Ewmo> =>

5.3.00 5

.529»: u.—

“m~_m:5_ua=c 9.55:8 m5
; E rhea: mvaoizv coca—Eon cache... 5 2:5 02:00:20

 

9:95th :2 //,,I\\

 

 

 

 

 

 

 

 

-/’\/’/

 

:wgwum :2 uamguohm a 29.3. E 9022.:
K. SEED 8 3:0 >20 8
mmémnuw a ., 22:5 :5 2.33.1.5 5
, x, , / aao m=n_._ 3: 55031530: 3:
NEWN a ,, / I 5.35 .5 5:: >3~z :z
T , . ommmwwbﬁ A mwhwm um cone—.3. .5533 3 :1
H \ if: x8.— «Emm ._n_ ~___>mv_w:no:w 5
V. x I. 2::— 2 9.09m om
umcocmrﬁznsa ﬂmnmhvmzc
«.235 \L
hwmg acozaEuou 2.55; to..— van
cannon .«o 23:52 851305 :32 van 3252 we 8an
«mo—p.532 an 3353 .0 .355: >256 393280 .m.D mmn NNKmA NN
Gaza—nuke“—
j awn—w “En 2.3355 can—3&2 *0 9:5sz £59. gm
I <Z<_DZ_ “Em otcvm Eat meEMm we 53:5: auaaw .mﬁmo—OaO .m.D FOB NN.Om©NN
\ «Bun—m PEA—5w 3 .>:n£< 252
vb \ we :2: «mo—Ewan: Ea: «29:: «o 23E:: 2an ES"— omor 9 U For
I _
xi ( l
L 3:2 2: JOEKZSQXM
_ b _ , _ _

 

owm

 

0mm 03 0mm

   

0mm

omm

110

TABLE 3.—Conodonts from top one inch of New Albany and
Sunbury Shales in east-central, eastern centraL and south-central
Kentucky

[Locations shown in fig. 67; identifications by J .W. Huddle. 1964; number of
specimens shown in parentheses]

 

Samples 1010-1090 collected from uppermost 1 in. of New
Albany or Sunbury Shales.

Samples 1010: State Hwy. 15, about 3 mi east of Clay City
and ‘Aa mi northeast of where Hwy. 15 crosses Hatton Creek,
Clay City quadrangle, Kentucky.

Hindeodella sp. (7)

Eupn’oniodina sp. (6)

Spathognathodus sp. (1)

Scolecodont (1)

Fish scales, conical fish teeth, plant trash, linguloid and or-

biculoid brachiopods

Sample 1020: about 1,000 ft south of southernmost part of
Ravenna, Ky., city limits, on the left bank of Cow Creek above
L&N Railroad, Irvine quadrangle, Kentucky.

Hindeodella sp. (6)

Polygnathids indeterminate (4)

Bryantodus sp. (1)

Siphonodella sp. (4)

Elictognathus sp. (1)

Euprioniodella sp. (1)

Ligonodina sp. (1)

Conical fish tooth (1)

Sample 1030: About 2 mi south of Panola, Ky. on dirt road
about 600 ft southwest of Knob Lick School and about 1,300
ft northeast of Knob Lick Cemetery, Panola quadrangle,
Kentucky.

Euprioniodina sp. (9)

Hindeodella sp. (51)

Spathognathodus sp. (3)

Ozarkodina sp. (8)

Gnathodus? sp. (2)

Lonchodina sp. (1)

Polygnathids indeterminate (16)

Linguloid and orbiculoid brachiopods

Sample 1040: US. Hwy. 25, about 21/2 mi south of L&N
Railroad tunnel in Berea, Ky., and about ‘% mi north of Boone
Gap, Berea quadrangle, Kentucky.

Siphonodella Sp. (17)

Siphonodella duplicata Branson and Mehl (2)

Hindeodella sp. (33)

Ligonodina sp. (2)

Bryantodus sp. (2)

Euprioniodina sp. (4)

Spathognathodus sp. (9)

Pn'oniodina sp. (2)

Ozarkodina sp. (2)

Dipiododella sp. (1)

Falcodus sp. (1)

Polygnathids indeterminate (6)

Linguloid brachiopod
Sample 1050: About 2% mi east of Crab Orchard on US. Hwy.
150, about 900 ft east of Slate Creek and about 800 ft west
of Lincoln-Rockcastle County line, Brodhead quadrangle,
Kentucky.

 

MISSISSIPPIAN ROCKS IN KENTUCKY

Hindeodella sp. (14)
Ozarkodina sp. (1)
Lonchodina? sp. (1)
Euprioniodina sp. (1)
Hibbardella sp. (1)
Polygnathid indeterminate (3)
Orbiculoid brachiopods

Sample 1060: On edge of reservoir near road from old US.
Hwy. 27 to Mayfield, Ky., about 1/2 mi northeast of Halls Gap,
Ky., Halls Gap quadrangle, Kentucky.

Hindeodella sp. (14)

Hibbardella sp. (2)

Spathognathodus sp. (3)

Pinacodus? sp. (1)

Euprioniodina sp. (1)

Lonchodina sp. (1)

Ligonodina? sp. (1)

Spores abundant

Sample 1070: State Hwy. 1248 about 3% mi southwest of

West Somerset on the east side of Lake Cumberland near
bottom of hill, Delmer quadrangle, Kentucky.

Eupn'oniodina sp. (1)

Hindeodella sp. (1)

Spathognathodus sp. (2)

Abundant spores and linguloid brachiopods

Sample 1080: About 31/2 mi southwest of Cartersville, Ky.
along road about 3/5 mi southwest of Pine Grove Church and
about 3/5 mi southeast of Bottom Lick Knob, Paint Lick
quadrangle, Kentucky.

Hindeodella sp. (1)

Spathognathodus sp. (1)

Shark’s tooth

Linguloid and orbiculoid brachiopods

Sample 1090: About 11/2 mi southeast of Bighill, Ky. along
road and about 800 ft east of point where the road crosses
Owsley Fork of Red Lick Creek, Bighill quadrangle, Kentucky.

Hindeodella sp. (5)

Siphonodella sp. (1)

Spathognathodus sp. (1)

Polygnathids indeterminate (2)

“Worm trails”

“All of the above samples were collected from the top inch
of the black ﬁssile shale (New Albany or Sunbury Shale). Near-
ly all of the samples are more or less weathered and most of
the conodonts are altered. The conodonts characteristic of the
Sunbury fauna are Siphonodella, Elictognathus, Pinacodus
and Gnathodus and these genera are also found in the Maury
Formation. Samples 102C, 1040, and 1090 are definitely
Mississippian, Kinderhookian in age. Samples 103C and 1060
are probably Mississippian in age. The age is based on the
questionably identified Gnathodus and Pinacodus. There are
no diagnostic conodonts in the other samples, and they
could be either Devonian or Mississippian. Hass (Prof. Paper
286, p. 24) reports a Siphonodella fauna in the basal New
Providence Shale at his locality 6 (1956, p. 27). near Sample
107C in Pulaski Co. *** this same Siphonodella fauna also
occurs in the top of New Albany Shale throughout southern
Indiana ***”

 

PALEONTOLOGY 1 1 1

TABLE 4.—Conodonts from basal part of Nancy (New Providence)
Member of the Borden Formation and from basal beds of the Fort
Payne Formation in east—central and south-central Kentucky

[Locations shown in fig. 67; identifications by J .W. Huddle, 1967; number of
specimens shown in parentheses; starred forms may be reworked]

 

Samples 227 95-PC—22799-PC from lowermost beds of Nancy
Member, Borden Formation.

USGS 22795-PC, field No. 103A, lowest bed in the Nancy
Member, Borden Formation (New Providence Shale), about
2 mi south of Panola, Ky., 600 ft southwest of Knob Lick
School and about 1,300 ft northeast of Knob Lick Cemetery,
Panola quadrangle, Kentucky.

Diplododella sp. (5)

Gnathodus antetexanus Rexroad and Scott (19)

Hindeodella sp. (137)

Ozarkodina roundyi Hass (18)

Polygnathus communis Branson and Mehl (81)

Polygnathus sp. (10)

*Pseudopolygnathus dentilineata Branson (29)

*Siphonodella quadruplicata? (Branson and Mehl) (3)

*Spathognathodus crassidentatus (Branson and Mehl) (11)

*S. stabilis (Branson and Mehl) (8)

Synpﬂoniodina sp. (3)
The presence of Gnathodus antetexanus suggests that this
may be as young as Cu II beta (G. semiglaber—Polygnathus
communis zone). Spathognathodus crassidentatus suggests
that Cu II alpha (Siphonodella cooperi-S. isosticha zone) may
be represented and Pseudopolygnathus dentilineata is charac-
teristic of Cu I strata. Very slow deposition or a lag concen-
trate is suggested. This seems more likely than reworking,
another possible explanation.

USGS 22796-PC, field No. 104A, lowermost Nancy Member,

Borden Formation (New Providence Shale Member), US.

Hwy. 25, about 21/2 mi south of Berea, Ky., and about 3/: mi

north of Boone Gap, Berea quadrangle, Kentucky.
Polygnathus communis Branson and Mehl (1)
Pseudopolygnathus prima Branson and Mehl (6)
Siphonodella duplicata (Branson and Mehl) (1)
Siphonodella quadruplicata (Branson and Mehl) (32

fragments)

Spathognathodus stabilis (Branson and Mehl) (1)
Spathognathodus sp. (3)

This collection probably belongs in the Siphonodella duplicata

zone of Collinson and others (1962).

USGS 2297-PC, field No. 105A, lowermost Nancy Member,
Borden Formation (New Providence Shale Member), about
23/4 mi east of Crab Orchard, Ky., on US. Hwy. 150, about

 

900 ft east of Slate Creek and about 800 ft west of the Lin-r
coln and Rockcastle County line, Brodhead quadrangle,
Kentucky.

Bactrognathus? sp. (1)

Hindeodella sp. (5)

Polygnathus communis Branson and Mehl (1)

Pn'oniodina sp. (1)

Pseudopolygnathus pn'ma Branson and Mehl (8)

Siphonodella duplicata (Branson and Mehl) (3)

Siphonodella sp. (29 fragments)

*Spathognathodus aculeatus (Branson and Mehl) (4)

Spathognathodus sp. (4)
The presence of Siphonodella duplicata and P. prima suggests
that this collection also represents the S. duplicata assemblage
zone. The specimens of Spathognathodus aculeatus are
interpreted as reworked. Bactrognathus has never been
reported below the “Sedalia” Limestone. Its presence in this
sample may be explained by contamination, misidentification,
stratigraphic leak, or as an extension of its range.

USGS 22798-PC, field No. 107A, lowermost Nancy Member,
Borden Formation (New Providence Shale Member), State
Hwy. 1248 about 3% mi southwest of West Somerset on the
east side of Lake Cumberland near bottom of hill, Delmer
quadrangle, Kentucky.

Gnathodus commutatus Branson and Mehl (2)

Hindeodella Sp. (14)

Ozarkodina roundyi Hass (2)

Polygnathus symmetrica Branson (8)

Polygnathus sp. (2)

*Spathognathodus aculeatus (Branson and Mehl) (10)

S. anteposicornis Scott (7)

S. linguliferus? (Branson) (3)

S. praelongus Cooper (45)

S. stabilis (Branson and Mehl) (4)
This collection is placed in the upper part of the
Spathognathodus costatus zone (equivalent to the Louisiana
Limestone) bcause of the presence of Spathognathodus
anteposicornis and Gnathodus commutatus.

USGS 22799-PC, ﬁeld N 0. 107B, same locality as 107A, from
rock less than 1 in. thick transitional between the New Albany
Shale and New Providence Shale Member.

Hindeodella sp. (2)

Spathognathodus praelongus Cooper (4)

S. stabilis (Branson and Mehl) (3)

Synprioniodina sp. (1)
Spathognathodus praelongus has previously been reported
from the S. costatus zone in Montana.

 

112

TABLE 5.—Conodonts from basal (New Providence) part of Nancy
Member of the Borden Formation in east-central and south-central
Kentucky

[Locations shown in fig. 67; identiﬁcations by J .W. Huddle, 1967; number of
specimens shown in parentheses; starred forms may be reworked]

 

Samples 22793-PC and 22794-PC from lowermost Fort
Payne Formation and basal beds of Nancy Member, Borden
Formation.

USGS 22793-PC, field No. 2 of 5/7/65 Huddle, Branson, and
Lewis. Limestone lens about 1 ft above the Chattanooga Shale
in the Fort Payne Formation, in a small west-ﬂowing tributary
of Bear Creek, 1.23 mi west of Seminary School, Burkesville
quadrangle, Kentucky.

Bryantodus sp. (1)

Gnathodus antetexanus Rexroad and Scott (79)

Hindeodella sp. (14)

Polygnathus communis Branson and Mehl (13)

Pn'oniodina pulcher Branson (2)

Pseudopolygnathus multistriata Mehl and Thomas (6)

*Siphonodella duplicata (Branson and Mehl) (4)

*S. quadruplicata (Branson and Mehl) (24)

*S. obsoleta? Hass (2)

*Spathognathodus aculeatus (Branson and Mehl) (9)

Synpn'oniodina sp. (2)
The presence of Gnathodus antetexanus and Pseudo-
polygnathus multistriata in this collection suggests that the
lower part of the Fort Payne Chert is equivalent to the upper
part of the Rockford Limestone of Indiana (Rexroad and Scott,
1964) and the “Sedalia” Limestone. The assemblage probably
represents the Gnathodus semiglaber-Pseudopolygnathus
multistriata zone of Collinson and others (1962).

USGS 22794-PC, field N o. 1 of 5/5/65 Huddle, Branson, and
Weir (same locality as USGS 22796-PC and USGS 7652-SD).
0.0-0.8 ft above base Nancy Member of the Borden Forma-
tion (New Providence Shale Member), US. Hwy. 25 about
21/2 mi south of Berea and 3/: mi north of Boones Gap, Berea
quadrangle, Kentucky.

Hindeodella sp. (2)

Polygnathus communis Branson and Mehl (2)

P. inornata Branson (1)

P. permarginata Branson (1)

P. scorbiformis Branson (1)

Pseudopolygnathus prima Branson and Mehl (1)

Siphonodella duplicata (Branson and Mehl) (10)

S. cooperi? Hass (7 fragments)
This collection is assigned to the Siphonodelhz duplicata zone
of Collinson and others, 1962.

 

TABLE 6.—Conodonts from basal part of New Providence Shale
Member of the Borden Formation in central Kentucky

[Locations shown in fig. 67 ; identifications by J .W. Huddle, 1967; number of
specimens shown in parentheses; starred forms may be reworked]

 

Samples 22787-PC—22792-PC from lowermost beds of New
Providence Shale member. Borden Formation.

USGS 22787-PC, ﬁeld No. 5 of 5/12/65 Huddle and Branson.
0-1.2 ft above base of New Providence Shale Member, on knob
south of road to Bemheim Forest opposite the County Farm
near Kentucky Turnpike, Ky. State coord. 1,584,600 E. by
158,100 N ., Shepherdsville quadrangle, Kentucky.

 

MISSISSIPPIAN ROCKS IN KENTUCKY

Bactrognathus hamata Branson and Mehl (7)

Hindeodella sp. (18)

Neopn'oniodus sp. (2)

Ozarkodina sp. (1)

Polygnathus communis Branson and Mehl (7)

*Polygnathus inomata Branson (6)

*P. longiposita Branson (6)

*P. symmetn’ca Branson (5)

*Pseudopolygnathus pn'ma Branson and Mehl (18)

*P. sp. (1)

*Siphonodella cooper-i Hass (13)

*Siphonodella duplicata (Branson and Mehl) (14)

*S. quadruplicata (Branson and Mehl) (4)

*Spathognathodus aculeatus (Branson and Mehl) (23)
The specimens in this collection do not look reworked and the
abundance of S. aculeatus suggests the presence of the Lower
or Middle Spathognathodus cos tatus zone (to V1). The young-
est zone present, the Bactrognathus—Polygnathus communis
zone, is indicated by Bact'rognathus hamata and P. communis.

USGS 22788-PC, field N o. 1 of 5/12/65 Huddle and Branson.
0.6-1.8 ft above the base of the New Providence Shale
Member, roadcut on north side of US. Hwy. 62, about 1%
mi east of Boston, Ky., Lebanon Junction quadrangle, Ken-
tucky State coords. 1,596,400 E. by 107,700 N.

Gnathodus texanus Roundy (2)

Hindeodella sp. (6)

Spathognathodus sp. (2)

Fish remains and white hollow spheres uneven in size
All of the conodonts are weathered white. The age is
presumably the same as USGS 22792—PC and represents the
time equivalent of the upper Burlington Limestone, according
to Rexroad and Scott (1964).

USGS 22789-PC, field No. 2 of 5/ 12/65 Huddle and Branson.
0.0-0.6 ft above base of New Providence Shale Member, same
locality as USGS 22788-PC.

Bactrognathus distorta? Branson and Mehl (1)

Bryantodus sp. (1)

Hindeodella sp. (17)

Polygnathus communis Branson and Mehl (2)

*Siphonodella sp. (two or more species) (17 fragments)

*Polygnathus inomata Branson (8)

*P. bngiposita? Branson (2)

*P. cf. P. scorbiformis Branson (10)

*Pn'oniodina sp. (1)

*Spathognathodus aculeatus (Branson and Mehl) (21)

*S. costatus? (Branson) (3)

*Spathognathodus aff. S. disparilis (Branson and Mehl) (5)

Spathognathodus sp. (4)
The conodonts in this collection are weathered white and the
forms starred are thought to be reworked. It is possible that
there are thin layers at the base representing the
Spathognathodus costatus zone and the Siphonodella zones,
but reworking of the older forms seems more probable. This
collection apparently belongs in the Bactrognathus—
Taphrognathus zone of Collinson and others (1962) and is
equivalent to the upper part of the Burlington Limestone.

USGS 22790-PC, ﬁeld No. 7 of 5/12/65 Huddle and Branson.
Basal 0—1.3 ft of the New Providence Shale in roadcut on north
side of Blue Gap, coord. 2,053,500 E. by 493,500 N., New
Haven quadrangle, Kentucky.

Gnathodus texanus Roundy (2)

Hindeodella sp. (1)

REFERENCES CITED

TABLE 6.—Conodonts from basal part of New Providence Shale
Member of the Borden Formation in central Kentucky—Continued

[Locations shown in fig. 67; identiﬁcations by J .W. Huddle. 1967; number of
specimens shown in parentheses; starred forms may be reworked]

 

Pn’oniodina sp. (1)

*Pseudopolygnathus pn’ma Branson and Mehl (3)

*Siphonodella sp. (4)

*Spathognathodus aculeatus Branson and Mehl (8)

S. sp. (4)
Most of the conodonts are fragmented, and only a few can be
identified. They are weathered white, and some are stained
with limonite.
The presence of Gnathodus texanus indicates that the upper
part of this sample, at least, belongs in the Bactrognathus—
Taphrognathus zone.

USGS 227 91-PC, ﬁeld N o. 2 of 5/11/65 Huddle, Branson, Sable,
Peterson, and Kepferle. Basal 0.8 ft of the New Providence
Shale Member, new roadcut in Kentucky. State Hwy. 247, Ky.
State coord. 2,053,200 E. by 470,300 N., Howardstown
quadrangle, Kentucky.

Bryantodus sp. (1)

Gnathodus antetexanus Rexroad and Scott (14)

G. distorta Branson and Mehl (3)

Hindeodella sp. (35)

Lonchodina sp. (1)

Ozarkodina sp. (1)

*Palmatolepis gracilis Branson and Mehl (1)

Polygnathus communis Branson and Mehl (27)

*Polygnathus inomata Branson (12)

*P. permarginata Branson (2)

*P. symmetrica Branson (12)

Prioniodina sp. (1)

Prioniodina sp. (1)

Pseudopolygnathus prima Branson and Mehl (12)

*Siphonodella duplicata (Branson and Mehl) (10)

*8. sp. (10)

*Spathognathodus aculeatus (Branson and Mehl) (60)

*S. anteposicornis Scott (3)

*S. dispan'lis (Branson and Mehl) (3)

S. aff. S. werneri Ziegler (54)

S. sp. (6)
Conodonts and fish remains are abundant in this sample, but
most of them are broken and cannot be identified. The relative
abundance of the species is not properly represented by the
counts above.
The interpretation of this collection is uncertain. A zone of
large phosphate nodules at the base suggests slow deposition
and lag concentration. The starred forms represent Late Devo-
nian and Early Mississippian conodonts. They may have been
reworked or several zones may be present in the basal 0.8 foot
interval of the New Providence Shale Member. Inch by inch
collecting would be necessary to prove or disprove this
possibility.
The youngest form present is Gnathodus antetexanus, which
suggests that at least the upper part of the interval is
equivalent to the upper part of the Rockford Limestone; and
the presence of Spathognathodus anteposicornis suggests that
the equivalent of the Louisiana Limestone is present.

USGS 22792-PC, field N o. 3 of 5/11/65 Huddle, Branson, Sable,
Peterson and Kepferle. 0.8—1.9 ft above the base of the New
Providence Shale Member. Same locality as USGS 227 91-PC.
Gnathodus texanus Roundy (15)
Hindeodella sp. (22)

 

113

Polygnathus communis Branson and Mehl (3)

Pn’oniodina acuta Branson (3)

P. pulcher Branson (2)

Pseudopolygnathus prima Branson and Mehl (3)

*Siphonodella sp. (1 fragment)

*Spathognathodus aculeatus (Branson and Mehl) (3)

*8. cf. S. abnormis (Branson and Mehl) (6)

*S. subrecta (Holmes) (1 weathered white)

Cusp fragments, clear with no white matter (7)
Gnathodus texanus has not been reported previously below
the upper Burlington Limestone and first appears in the
Bactrognathus—Taphrognathus zone, according to Rexroad
and Scott (1964). Starred forms probably reworked.

 

REFERENCES CITED

Alvord, DC, 1971, Geologic map of the Hellier quadrangle, Kentucky-
Virginia and parts of the Clintwood quadrangle, Pike County, Ken-
tucky: U.S. Geologic Quadrangle Map GQ-950.

Alvord, DC, and Miller, RC, 1972, Geologic map of the Elkhorn City
quadrangle, Kentucky-Virginia and part of the Harmon
quadrangle, Pike County, Kentucky: US. Geological Survey
Geologic Quadrangle Map GQ-951.

American Commission on Stratigraphic Nomenclature, 1961, Code of
Stratigraphic Nomenclature: American Association of Petroleum
Geologists Bulletin, v. 45, no. 5, p. 645-665.

1970, Code of Stratigraphic Nomenclature, 2d edition: American

Association of Petroleum Geologists, Tulsa, Okla, 45 p.

1983, North American stratigraphic code, the North American
Commission on stratigraphic nomenclature: American Associa-
tion of Petroleum Geologists Bulletin, v. 67 , no. 5, p. 841—875.

Amos. D.H.,1965, Geology of parts of the Shetlerville and Rosiclare
quadrangles. Kentucky: US. Geological Survey Geologic
Quadrangle Map GQ-400.

1967, Geologic map of part of the Smithland quadrangle, Liv-

ingston County, Kentucky: US. Geological Survey Geologic

Quadrangle Map GQ—657.

1971, Geologic map of part of the Leavenworth quadrangle,
Meade County, Kentucky: US. Geological Survey Geologic
Quadrangle Map GQ—941.

_1972, Geologic map of the New Amsterdam quadrangle,
Kentucky-Indiana, and part of the Mauckport quadrangle, Ken-
tucky: US. Geological Survey Geologic Quadrangle Map GQ-990.

1974, Geologic map of the Burna quadrangle, Livingston
County, Kentucky: US. Geological Survey Geologic Quadrangle
Map GQ-l 150.

Amos, D.H., and Finch, W.I., 1968. Geologic map of the Calvert City
quadrangle, Livingston and Marshall Counties, Kentucky: US
Geological Survey Geologic Quadrangle Map GQ-731.

Amos, D.H., and Hays, W.H., 1974, Geologic map of the Dycusburg
quadrangle, western Kentucky: US. Geological Survey Geologic
Quadrangle Map GQ-1149.

Bassler, RS, 1932, The stratigraphy of the Central basin of Tennessee:
Tennessee Division Geology Bulletin 38, 268 p.

Baxter, J .W., 1960. Salem Limestone in southwestern Illinois: Illinois
State Geological Survey Circular 284, 32 p.

Baxter, J .W., Desborough, G.A., and Shaw, C.W., 1967, Areal geology
of the Illinois Fluorspar district, Part 3—Herod and Shetlerville
quadrangles: Illinois State Geological Survey Circular 413, 41 p.

Baxter. J .W., Potter, RE, and Doyle, EL. 1963, Areal geology of
the Illinois Fluorspar district, Part l—Saline Mines. Cave in Rock,
DeKoven. and Repton quadrangles: Illinois State Geological
Survey Circular 342, 43 p.

 

 

 

 

 

114

Benson, D.J., 1976, Lithofacies and depositional environments of
Osagean-Meramecian platform carbonates, southern Indiana.
central and eastern Kentucky: Cincinnati. Ohio, University of Cin-
cinnati Ph. D. thesis, 215 p.

Brenner, J .C.. 1888, Administrative report and introduction to the
report upon western-central Arkansas: Arkansas Geological
Survey Annual Report 1. p. xv—xxxi.

Bristol, H.M., and Howard, R.H.. 1971, Paleogeologic map of the sub-
Pennsylvanian Chesterian (Upper Mississippian) surface in the
Illinois Basin: Illinois State Geological Survey Circular 458. 13 p.

Brown, W.R.. and Osolnik, M.J., 1974, Geolog'c map of the Livingston
quadrangle. southeastern Kentucky: US. Geological Survey
Geologic Quadrangle Map GQ—1179.

Browne. R.G., Baxter, J.W., and Roberts, T.G., 1977, The Archae-
discidae of the Fraileys facies (Mississippian) of central Kentucky:
Bulletins of American Paleontology. v. 72. no. 298, p. 171—228.

Browne, R.G.. and Bryant, A.L., 1970, Plant fossils from the St. Louis
Formation in Kentucky: Journal of Paleontology, v. 44, no. 3,
p. 520—521.

Browne. R.G., and Feb], ER. 1973, Stratigraphy and genera of
calcareous foraminifera of the Fraileys facies (Mississippian) of
central Kentucky: Bulletins of American Paleontology, v. 64, no.
280, p. 173—243.

Bucher, W.H., 1919. On ripples and related sedimentary surface forms
and their paleogeographic interpretation: American Journal of
Science. ser. 4. v. 47, p. 149—210, 241—269.

Burger, A.M.. Rexroad, C.B., Schneider, A.F., and Shaver, RH, 1966.
Excursions in Indiana geology: Association of American State
Geologists, 58th Meeting, Indiana Geological Survey Guidebook
12, 71 p.

Burroughs, W.G., 1923, A Pottsville-filled channel in the Mississip-
pian: Kentucky Geological Survey, ser. 6, v. 10. p. 115—126.

Butts, Charles, 1915. Geology and mineral resources of Jefferson Coun-
ty, Kentucky: Kentucky Geological Survey, ser. 4, v. 3, pt. 2, 270 p.

1917, Descriptions and correlation of the Mississippian forma-
tions of western Kentucky: Kentucky Geological Survey. ser. 4,
v. 5, pt. 1, 119 p.

_1922, The Mississippian series of eastern Kentucky: Kentucky
Geological Survey, ser. 6, v. 7, 188 p.

1940, Geology of the Appalachian valley in Virginia: Virginia
Geological Survey Bulletin 52, pt. 1, 568 p.

Calvert, R.B., 1968, Stratigraphy of a Chester cycle in the Kentucky
part of the Eastern Interior basin: St. Louis, Mo., Washington
University MS. thesis,126 p.

Calvert, W.L.. Bemhagen, R.J., J anssens, Arie, Noger, M.C., and Bran-
son, E.R.. 1968. Geological aspects of the Maysville-Portsmouth
region, southern Ohio and northeastern Kentucky; Ohio Geological
Society and Geological Society of Kentucky joint ﬁeld conference.
May 17—18, 1968: Lexington, Ky., Kentucky Geological Survey,
86 p.

Campbell, Guy, 1946, New Albany Shale: Geological Society of
America Bulletin. v. 57, no. 9, p. 829—908.

Campbell. M.R., 1893, Geology of the Big Stone Gap coal field of
Virginia and Kentucky: US. Geological Survey Bulletin 111,
106 p.

Chaplin, J .R., 1980, Stratigraphy, trace fossil associations, and deposi-
tional environments in the Borden Formation (Mississippian),
northeastern Kentucky: Annual ﬁeld conference of the Geological
Society of Kentucky—lexington, Ky., Kentucky Geological Survey,
1 14 p.

Chaplin. J .R., and Mason. GE, 1979, Upper part of the Borden For-
mation and the lower Newman Limestone. in Ettensohn, F'.R., and
Dever, G.R., J r., eds, Carboniferous geology from the Appalachian
basin to the Illinois basin through eastern Ohio and Kentucky;
Field trip no. 4. 9th International Congress of Carboniferous

 

 

 

MISSISSIPPIAN ROCKS IN KENTUCKY

Stratigraphy and Geology: Lexington, Ky., University of Ken-
tucky, p. 138—143.

Chesnut. D.R., 1980. Echinoderms from the lower part of the Penn-
ington Formation (Chesterian) in south-central Kentucky: Lex-
ington, Ky., University of Kentucky M.S. thesis, 178 p.

Clark. L.D., and Crittenden, M.D., Jr.. 1965, Geology of the Mattingly
quadrangle, Kentucky-Indiana: US. Geological Survey Geologic
Quadrangle Map GQ-361.

Cohee, G.V., and West, W.S.. 1965, Changes in stratigraphic
nomenclature by the US. Geological Survey, 1964: US. Geological
Survey Bulletin 1224-A, p. A1—A22.

Collinson, C.W., 1964, Field guidebook. 28th Annual 'I‘ri-State
Geological Field Conference. western Illinois: Illinois State
Geological Survey Guidebook, ser. 6, 30 p.

__1969, Devonian-Mississippian biostratigraphy of northeastern
Missouri and northwestern Illinois: North American Paleon-
tological Convention Field Trip 3. 23 p.

Collinson, C.W., Rexroad, 0.3., and Thompson. T.L., 1971, Conodont
zonation of the North American Mississippian. in Sweet, W.C.,
and Bergstrom. S.M., eds.. Symposium on conodont biostratig-
raphy: Geological Society of America Memoir 127, p. 353—394.

Collinson, C.W., and Scott, A.J.. 1958, Age of the Springville Shale
(Mississippian) of southern Illinois: Illinois State Geological
Survey Circular 254, 12 p.

Collinson, C.W., Scott. A.J., and Rexroad, 0.3., 1962, Six charts show-
ing biostratigraphic zones and correlations based on conodonts
from the Devonian and Mississippian rocks of the Upper Missis-
sippi Valley: Illinois State Geological Survey Circular 328, 32 p.

Conant, LC, and Swanson, V.E., 1961, Chattanooga shale and related
rocks of central Tennessee and nearby areas: US. Geological
Survey Professional Paper 357, 91 p.

Conkin, J .E.,1954, Hyperammina kentuckyensis n. sp. from the Mis-
sissippian of Kentucky, and discussion of Hyperammina and
Hyperamminoides: Cushman Foundation for Foraminiferal
Research Contributions, v. 5, pt. 4, no. 119. p. 165-169.

1956, Hyalostelia ancora Gutschick in the Mississippian of Indi-

ana and Kentucky: American Midland Naturalist, v. 56, p. 430—433.

1957, Stratigraphy of the New Providence Formation (Missis-

sippian) in Jefferson and Bullitt Counties, Kentucky. and fauna

of the Coral Ridge member: Bulletins of American Paleontology,

v. 38, no. 168, p. 109—157.

1960, Mississippian smaller foraminifera of southern Indiana,

Kentucky. northern Tennessee. and south-central Ohio: Cincin-

nati, Ohio. University of Cincinnati Ph. D. thesis, 413 p.

1961, Mississippian smaller foraminifera of Kentucky, southern
Indiana, northern Tennessee, and south-central Ohio: Bulletins
of American Paleontology, v. 43, no. 196, 368 p.

Conkin. J .E., and Conkin, B.M., eds., 1979, Devonian-Mississippian
boundary in southern Indiana and northwestern Kentucky—Field
trip 7, 9th International Congress of Carboniferous Stratigraphy
and Geology: Sponsors, Graduate School, University of Louisville,
and Bloomington Crushed Stone Company, Bloomington, Indiana,
141 p.

Connor, J.J., 1969, A geochemical discriminant for sandstones of
Mississippian and Pennsylvanian age in Kentucky: Colorado
School of Mines Quarterly. v. 64, no. 3, p. 17-36.

Connor, J .J ., and Trace, R.D., 1970. Outlier of Caseyville Sandstone
near Princeton. Kentucky, may be Bethel Sandstone: U.S. Geo-
logical Survey Professional Paper 700-D, p. D33-D35.

Cooper, B.N.. 1944. Geology and mineral resources of the Burkes
Garden quadrangle, Virginia: Virginia Geological Survey Bulletin
60, 299 p.

Craig, LC. and Connor, C.W., coordinators. 1979, Paleotectonic in-
vestigations of the Mississippian System in the United States:
US. Geological Survey Professional Paper 1010. pt. I, 369 p.

 

 

 

 

REFERENCES CITED

Craig, L.C., and Varnes, KL, 1979, History of the Mississippian
System—An interpretive summary: US. Geological Survey Pro-
fessional Paper 1010, pt. II, p. 371—406.

Cressman, ER, and Noger, M.C., 1981, Geologic mapping of
Kentucky—History and evaluation of the Kentucky Geological
Survey-US. Geological Survey mapping program, 1960—1978:
US. Geological Survey Circular 801, 22 p.

Crider, A.F., 1913, The fire clays and fire clay industries of the Olive
Hill and Ashland districts of northeastern Kentucky: Kentucky
Geological Survey, ser. 4, v. 1, pt. 2, p. 589—711.

Cross, A.T., and Hoskins, J .H., 1951, Paleontology of the Devonian-
Mississippian black shales: Journal of Paleontology, v. 25, no. 6,
p. 713-728.

1952, The Devonian-Mississippian transition ﬂora of east-
central United States: Heerlen, 'I‘roisieme Congres Pour
L Avancement des Etudes de Stratigraphie et de Géologie du
Carbonifere, Compte Rendu, tome 1, p. 113-122.

Csej tey, Bela, J r., 1970, Geologic map of the N olansburg quadrangle.
southeastern Kentucky: US. Geological Survey Geologic Quad-
rangle Map GQ-868.

1971, Geologic map of the Bledsoe quadrangle, southeastern
Kentucky: US. Geological Survey Geologic Quadrangle Map
GQ-889.

Cumings, ER, 1901, The use of Bedford as a formational name: J our-
nal of Geology, v. 9, p. 232-233.

1922, Nomenclature and description of the geological forma-
tions of Indiana, in Handbook of Indiana geology: Indiana Depart-
ment of Conservation Publication 21, pt. 4, p. 403—570.

DeLaney, A.O., and Englund, K.J., 1973, Geologic map of the Ault
quadrangle, northeastern Kentucky: US. Geological Survey
Quadrangle Map GQ-1066.

Denny, CS, 1964, Geology of the Brushart quadrangle, Kentucky:
US. Geological Survey Geologic Quadrangle Map GQ-324.
Dever, G.R., J r., 1971, Guide to the geology along Interstate Highway

64 and the Mountain Parkway, Lexington to Wolfe County, Ken-
tucky. in Smith, GE, and others, Depositional environments of
eastern Kentucky coals: Annual ﬁeld conference, Geological Socie-
ty of America. Coal Division, Lexington, Ky., Kentucky Geological

Survey, p. 2—12.

1973, Stratigraphic relationships in the lower and middle

Newman Limestone (Mississippian), east-central and northeastern

Kentucky: Lexington, Ky., University of Kentucky M.S. thesis,

121 p.

1977, The lower Newman Limestone: Stratigraphic evidence for

Late Mississippian tectonic activity, in Dever, G.R., Jr., and

others, Stratigraphic evidence for late Paleozoic tectonism in

northeastern Kentucky—Field trip guidebook, Eastern Section,

American Association of Petroleum Geologists October 9, 1976:

Lexington, Ky., Kentucky Geological Survey, p. 8-18.

1980a, Stratigraphic relationships in the lower and middle

Newman Limestone (Mississippian), east-central and northeastern

Kentucky: Kentucky Geological Survey, ser. 11, Thesis Series 1,

49 p.

1980b, The Newman Limestone—An indicator of Mississippian
tectonic activity in northeastern Kentucky, in Luther, M.K., ed.,
Proceedings of the technical sessions, Kentucky Oil and Gas
Association 36th and 37th annual meetings, 1972 and 1973: Ken-
tucky Geological Survey, ser. 11, Special Publication 2, p. 42—54.

Dever, G.R., Jr., Ellsworth, G.W., Jr., and Sumartojo, Jojok, 1978,
Evidence for deposition of evaporites in the St. Louis Limestone
(Mississippian) of eastern Kentucky: Geological Society of
America Abstracts with Programs, v. 10, p. 167.

Dever, G.R., Jr., Hester, N.C., Ettensohn, RR, and Moody, J.R., 1979,
Newman Limestone (Mississippian) of east-central Kentucky and
Lower Pennsylvanian slump structures, in Ettensohn, F.R., and

 

 

 

 

 

 

 

 

115

Dever, G.R., J r., eds., Carboniferous geology from the Appalachian
basin to the Illinois basin through eastern Ohio and Kentucky,
Field trip no. 4, 9th International Congress of Carboniferous
Stratigraphy and Geology: Lexington, Ky., University of Ken-
tucky, p. 175—181.

Dever, G.R., Jr., Hoge, H.P., Hester, N.C., and Ettensohn, ER, 1977,
Stratigraphic evidence for late Paleozoic tectonism in northeastern
Kentucky, Field trip guidebook, Eastern Section, American
Association of Petroleum Geologists, October 9, 1976: Lexington,
Ky., Kentucky Geological Survey, ser. 10, 80 p.

Dever, G.R., J r., and McGrain, Preston, 1969, High-calcium and low-
magnesium limestone resources in the region of the lower
Cumberland, Tennessee, and Ohio valleys, western Kentucky: Ken-
tucky Geological Survey, ser. 10, Bulletin 5, 192 p.

Dever, G.R., Jr., McGrain, Preston, and Ellsworth, G.W., Jr., 1979,
Girkin Formation of west-central Kentucky, in Ettensohn, F.R.,
and Dever, G.R., J r., eds., Carboniferous geology from the Ap-
palachian basin to the Illinois basin through eastern Ohio and Ken-
tucky, Field trip no. 4, 9th International Congress of Carboniferous
Stratigraphy and Geology: Lexington, Ky., University of Ken-
tucky, p. 234—238.

Dever, G.R., J r., McGrain, Preston, Ellsworth, G.W., J r., and Moody,
J .R., 1979, Features in the upper Ste. Genevieve of west-central
Kentucky, in Ettensohn, RR, and Dever, G.R., Jr., eds., Car-
boniferous geology from the Appalachian basin to the Illinois basin
through eastern Ohio and Kentucky, Field trip no. 4, 9th Inter-
national Congress of Carboniferous Stratigraphy and Geology:
Lexington, Ky., University of Kentucky, p. 261—263.

Dever, G.R., Jr., McGrain, Preston, and Moody, J .R., 1979, St. Louis
Limestone of eastern Kentucky, in Ettensohn, RR, and Dever,
G.R., J r., eds., Carboniferous geology from the Appalachian basin
to the Illinois basin through eastern Ohio and Kentucky, Field
trip no. 4, 9th International Congress of Carboniferous Stratig-
raphy and Geology: Lexington, Ky., University of Kentucky,
p. 171-174.

1980, Stratigraphic and economic signiﬁcance of the Lithostro—
tion (Siphonodendron) genevieuensis zone in the Ste. Genevieve
Limestone (Mississippian) of western Kentucky: Geological Socie
ty of America Abstracts with Programs, v. 12, p. 223.

Dever, G.R., Jr., and Moody, J .R., 1979a, Review of the lithostratig—
raphy of the Renfro Member of the Borden Formation (Mississip-
pian), east-central Kentucky: Geological Society of America
Abstracts with Programs, v. 11, no. 4, p. 176-177.

1979b, Salem and Warsaw Formations in south-central Ken-
tucky, in Ettensohn, F.R., and Dever, G.R., J r., eds., Carboniferous
geology from the Appalachian basin to the Illinois basin through
eastern Ohio and Kentucky, Field trip no. 4, 9th International Con-
gress of Carboniferous Stratigraphy and Geology: Lexington, Ky.,
University of Kentucky, p. 202-206.

de Witt, Wallace Jr., 1970, Age of the Bedford Shale, Berea Sand-
stone, and Sunbury Shale in the Appalachian and Michigan basins,
Pennsylvania, Ohio, and Michigan: US. Geological Survey
Bulletin 1294-G, p. Gl—Gll.

de Witt, Wallace J r., and McGrew, L.W., 1979, The Appalachian basin
region, in Craig, L.C., and Connor, C.W., coordinators, Paleotectonic
investigations of the Mississippian System in the United States:
US. Geological Survey Professional Paper 1010, pt. I, p. 13—48.

Dutro, J.T., Jr., Gordon, Mackenzie Jr., and Huddle, J.W., 1979,
Paleontological zonation of the Mississippian System, in Craig,
L.C., and Connor, C.W., Paleotectonic investigations of the
Mississippian System in the United States: US. Geological Survey
Professional Paper, pt. II, p. 407-429.

Eames, LE, 1978, A palynological interpretation of the Devonian-
Mississippian boundary from northeastern Ohio, U.S.A.:
Palynology, v. 2, p. 218—219.

 

 

116

Elam, T.D., 1981, Stratigraphy and paleoenvironmental aspects of
the Bedford-Berea sequence and the Sunbury Shale in eastern and
south-central Kentucky: Lexington, Ky., University of Kentucky
M.S. thesis, 155 p.

Elrod, M.N., 1899, The geologic relations of some St. Louis group caves
and sinkholes: Indiana Academy of Science Proceedings, 1898,
v. 8, p. 258—267.

Engelmann, George, 1847, Remarks on the St. Louis Limestone:
American Journal of Science, 2d ser., v. 3, p. 119—120.

Englund, K.J., 1964a, Stratigraphy of the Lee Formation in the
Cumberland Mountains of southeastern Kentucky, in Geological
Survey research 1964: US. Geological Survey Professional Paper
501-B, p. B30—B38.

_1964b, Geology of the Middlesboro South quadrangle,
Tennessee-Kentucky-Virginia: US. Geological Survey Geologic
Quadrangle Map GQ-301.

1969, Geologic map of the Jellico West quadrangle, Kentucky-

Tennessee: US. Geological Survey Geologic Quadrangle Map

GQ—855.

1972, Central Appalachian tectonics as indicated by structural

features in Carboniferous rocks [abs]: American Association of

Petroleum Geologists Bulletin, v. 56, p. 2108.

1974, Sandstone distribution patterns in the Pocahontas For-

mation of southwest Virginia and southern West Virginia, in

Briggs, Garrett, ed., Carboniferous of southeastern United States:

Geological Society of America Special Paper 148, p. 31—45.

1976, Geologic map of the Grahn quadrangle, Carter County,

Kentucky: US. Geological Survey Geologic Quadrangle Map

GQ-1262.

1979, The Mississippian and Pennsylvanian (Carboniferous)
Systems in the United States—Virginia: US. Geological Survey
Professional Paper 1110-C, p. C1—C21.

Englund, K.J., and DeLaney, A.O., 1966, Intertonguing relations of
the Lee Formation in southwestern Virginia: US. Geological
Survey Professional Paper 550-D, p. D47—D52.

Englund, K.J., Gillespie, W.H., Cecil, C.B., Windolph, J .F., Jr., and
Crawford, T.J., 1985, Characteristics of the Mississippian-
Pennsylvanian boundary and associated coal-bearing rocks in the
southern Appalachians: US. Geological Survey Open-File Report
85-577, 83 p.

Englund, K.J., and Henry, T.W., 1979, Mississippian-Pennsylvanian
boundary in the central Appalachians: 9th International Congress
of Carboniferous Stratigraphy and Geology, Compte Rendu, v.
2, p. 330—336.

Englund, K.J., Henry, T.W., and Cecil, 0.8., 1981, Upper Mississip-
pian and Lower Pennsylvanian depositional environments, south-
western Virginia and southern West Virginia, in Roberts. T.G.,
ed., Geological Society of America, Cincinnati, Field trip guide-
books, Volume 1, Stratigraphy, sedimentology: American
Geological Institute, p. 171—175.

Englund, K.J., Landis, ER, and Smith, H.L., 1963, Geology of the
Varilla quadrangle, Kentucky-Virginia: US. Geological Survey
Geologic Quadrangle Map GQ—190.

Englund, K.J., and Randall, A.H., III, 1981, Stratigraphy of the Upper
Mississippian and Lower Pennsylvanian Series in the east—central
Appalachians, in Roberts, T.G., ed., Geological Society of America,
Cincinnati, Field trip guidebooks, Volume 1, Stratigraphy.
sedimentology: American Geological Institute, p. 154—158.

Englund, K.J., Roen, J .B., and DeLaney, A.O., 1964, Geology of the
Middlesboro North quadrangle, Kentucky: US. Geological Survey
Geologic Quadrangle Map GQ-300.

Englund, K.J., and Smith, H.L., 1960, Intertonguing and lateral grada-
tion between the Pennington and Lee Formations in the tri-state
area of Kentucky, Tennessee, and Virginia [abs]: Geological Socie-
ty of America Bulletin, v. 71, p. 2015.

 

 

 

 

 

 

MISSISSIPPIAN ROCKS IN KENTUCKY

Englund, K.J., Smith, H.L., Harris, L.D., and Stephens, J .G., 1963,
Geology of the Ewing quadrangle, Kentucky and Virginia: US.
Geological Survey Bulletin 1142-3, 23 p.

Englund, K.J., and Windolph, J .F., Jr., 1971, Geology of the Carter
Caves Sandstone (Mississippian) in northeastern Kentucky: US.
Geological Survey Professional Paper 750-D, p. D99—D104.

1975, Geologic map of the Olive Hill quadrangle, northeastern
Kentucky: US. Geological Survey Geologic Quadrangle Map
GQ-1270.

Erickson, R.L., 1966, Geologic map of part of the Friendship
quadrangle, Lewis and Greenup Counties, Kentucky: US.
Geological Survey Geologic Quadrangle Map GQ-526.

Ettensohn, ER, 1975, Stratigraphic and paleoenvironmental aspects
of Upper Mississippian rocks (upper Newman Group), east-central
Kentucky: Urbana, 111., University of Illinois Ph. D. thesis, 320 p.

1977, Effects of synsedimentary tectonic activity on the up-

per Newman Limestone and Pennington Formation, in Dever,

G.R., J r., and others, Stratigraphic evidence for late Paleozoic tec-

tonism in northeastern Kentucky, Field trip guidebook, Eastern

Section, American Association of Petroleum Geologists, October

9, 1976: Lexington, Ky., Kentucky Geological Survey, p. 18-29.

1979a, The barrier-shoreline model in northeastern Kentucky,

in Ettensohn, RR, and Dever, G.R., Jr., eds., Carboniferous
geology from the Appalachian basin to the Illinois basin through
eastern Ohio and Kentucky, Field trip no. 4, 9th International Con-
gress of Carboniferous Stratigraphy and Geology: Lexington, Ky.,

University of Kentucky, p. 124—129.

1979b, Radioactive stratigraphy at the Devonian-Mississippian
boundary, in Ettensohn, RR, and Dever, G.R., Jr., eds., Car-
boniferous geology from the Appalachian basin to the Illinois basin
through eastern Ohio and Kentucky, Field trip no. 4, 9th Inter-
national Congress of Carboniferous Stratigraphy and Geology:
Lexington, Ky., University of Kentucky, p. 166-170.

___1980, An alternative to the barrier-shoreline model for deposi-
tion of Mississippian and Pennsylvanian rocks in northeastern
Kentucky: Geological Society of America Bulletin, v. 91, pt. 1,
p. 130—135; pt. 2. p. 934-1056.

1981, Mississippian-Pennsylvanian boundary in northeastern
Kentucky, in Roberts, T.G., ed., Geological Society of America,
Cincinnati, Field trip guidebooks, Volume 1, Stratigraphy, sedi-
mentology: American Geological Institute, p. 195-257.

Ettensohn, ER, and Chesnut, D.R., 1979a, Upper Mississippian car-
bonates and the Pennington Formation south of the basement fault
zone in eastrcentral Kentucky, in Ettensohn, ER., and Dever, G.R.,
J r., eds., Carboniferous geology from the Appalachian basin to the
Illinois basin through eastern Ohio and Kentucky, Field trip no.
4, 9th International Congress of Carboniferous Stratigraphy and
Geology: Lexington, Ky., University of Kentucky, p. 191—194.

1979b, Stratigraphy and depositional environments in the up-
per Hartselle, Bangor and Pennington Formations of south-central
Kentucky, in Ettensohn, RR, and Dever, G.R., Jr., eds., Car-
boniferous geology from the Appalachian basin to the Illinois basin
through eastern Ohio and Kentucky, Field trip no. 4, 9th Inter-
national Congress of Carboniferous Stratigraphy and Geology:
Lexington, Ky., University of Kentucky, p. 194—201.

Ettensohn, RR, and Dever, G.R., Jr., eds., 1979, Carboniferous
geology from the Appalachian basin to the Illinois basin through
eastern Ohio and Kentucky, Field trip no. 4, 9th International Con-
gress of Carboniferous Stratigraphy and Geology: Lexington, Ky.,
University of Kentucky, 293 p.

Ettensohn, F.R., Fulton, LR, and Kepferle, R.C., 1979, The use of
scintillometer and gamma-ray logs for correlation and stratigraphy
in homogeneous black shales—Surnmary: Geological Society of
America Bulletin, Part I, v. 90, p. 421—423 and Part II, v. 90,
p. 828-849.

 

 

 

 

 

 

REFERENCES CITED

Ettensohn, ER, and Peppers, RA, 1979, Palynology and biostratig-
raphy of Pennington shales and coals (Chesterian) at selected sites
in northeastern Kentucky: Journal of Paleontology, v. 53,
p. 453—474.

Ettensohn, F.R., Rice. C.L., Dever, G.R., Jr., and Chesnut, D.R., 1984,
Slade and Paragon Formations—New stratigraphic nomenclature
for Mississippian rocks along the Cumberland Escarpment in Ken-
tucky: U.S. Geological Survey Bulletin 1605-B, 37 p.

Ferm, J .C., 1974, Carboniferous environmental models in eastern
United States and their significance, in Briggs, Garrett, ed., Car-
boniferous of southeastern United States: Geological Society of
America Special Paper 148, p. 79—95.

Ferm, J .C., Horne, J .C., Swinchatt, J .P., and Whaley, P.W., 1971, Car-
boniferous depositional environments in northeastern Kentucky—
Annual spring field conference, Geological Society of Kentucky,
April 23—24, 1971: Lexington, Ky., Kentucky Geological Survey,
30 p.

Fox, K.F., J r., 1965, Geology of the Cadiz quadrangle, Trigg County,
Kentucky: US. Geological Survey Geologic Quadrangle Map
GQ-412.

Fox, K.F., Jr., and Olive, W.W., 1966, Geology of the Birmingham
Point quadrangle, western Kentucky: US. Geological Survey
Geologic Quadrangle Map GQ-471.

Fox, K.F., J r., and Seeland, D.A., 1964, Geology of the Canton quad-
rangle, Kentucky: U.S. Geological Survey Geologic Quadrangle
Map GQ-279.

Freeman, LB, 1951, Regional aspects of Silurian and Devonian
stratigraphy in Kentucky: Kentucky Geological Survey, ser. 9,
Bulletin 6, 565 p.

1953, Regional subsurface stratigraphy of the Cambrian and
Ordovician in Kentucky and vicinity: Kentucky Geological Survey,
ser. 9, Bulletin 12, 352 p.

Froelich, A.J., 1972, Geologic map of the Wallins Creek quadrangle,
Harlan and Bell Counties, Kentucky: US. Geological Survey
Geologic Quadrangle Map GQ-1016.

1973, Geologic map of the Louellen quadrangle, southeastern
Kentucky: US. Geological Survey Geologic Quadrangle Map
GQ-1060.

Froelich, A.J., and Stone, RD, 1973, Geologic map of parts of the
Benham and Appalachia quadrangles, Harlan and Letcher Coun-
ties, Kentucky: U.S. Geological Survey Geologic Quadrangle Map
GQ-1059.

Froelich, A.J., and Tazelaar, J .F., 1974, Geologic map of the Pineville
quadrangle, Bell and Knox Counties, Kentucky: US. Geological
Survey Geologic Quadrangle Map GQ-1129.

Frund, Eugene, 1953, Stratigraphy of the Borden Siltstone in Clin-
ton and adjacent counties, Illinois: Urbana, 111., Illinois Univer-
sity M.S. thesis, 37 p.

Fulton, LR, 1979, Structure and isopach map of the New Albany—
Chattanooga—Ohio Shale (Devonian and Mississippian) in
Kentucky—Eastern sheet: Kentucky Geological Survey, ser. 11,
scale 1:250,000.

Gildersleeve, Benjamin, 1963, Geology of the Bristow quadrangle, Ken-
tucky: U.S. Geological Survey Geologic Quadrangle Map GQ-216.

1966, Geologic map of the Homer quadrangle, Logan County,
Kentucky: US. Geological Survey Geologic Quadrangle Map
GQ-549.

__1971, Geologic map of the N olin Reservoir quadrangle, western
Kentucky: US. Geological Survey Geologic Quadrangle Map
GQ-895.

197 2, Geologic map of the Morgantown quadrangle, Butler and
Warren Counties, Kentucky: US. Geological Survey Geologic
Quadrangle Map GQ-1040.

Gildersleeve, Benjamin, and Johnson, W.D., J r., 1978, Geologic map
of the Caneyville quadrangle, Grayson County, Kentucky: US.
Geological Survey Geologic Quadrangle Map GQ-1472.

 

 

 

 

 

117

Gordon, Mackenzie, Jr., and Henry, T.W., 1981, Late Mississippian
and Early Pennsylvanian invertebrate faunas, east-central
Appalachians—A preliminary report, in Roberts, T.G., ed.,
Geological Society of America Cincinnati, Field trip guidebooks,
Volume 1, Stratigraphy, sedimentology: American Geological In-
stitute, p. 165-171.

Gordon, Mackenzie, Jr., and Mason, GE, 1985, Progradation of the
Borden Formation in Kentucky, U.S.A., demonstrated by suc-
cessive Early Mississippian (Osagean) ammonoid faunas: Field
Trip no. 4, 9th International Congress of Carboniferous Stratig-
raphy and Geology, p. 191—198.

Greenfield, R.E., 1957, Mississippian subsurface stratigraphy in Clay,
Harlan, and Leslie Counties, Kentucky: Lexington, Ky., Univer-
sity of Kentucky M.S. thesis, 46 p.

Gualtieri, J .L., 1967a, Geologic map of the Crab Orchard quadrangle,
Lincoln County, Kentucky: US. Geological Survey Geologic
Quadrangle Map GQ-571.

1967b, Geologic map of the Brodhead quadrangle, east-central

Kentucky: US. Geological Survey Geologic Quadrangle Map

GQ-662.

1968a, Geologic map of the Wildie quadrangle, Garrard and
Rockcastle Counties, Kentucky: US. Geological Survey Geologic
Quadrangle Map GQ-684.

_1968b, Geologic map of the J ohnetta quadrangle, Rockcastle
and Jackson Counties Kentucky: US. Geological Survey Geologic
Quadrangle Map GQ-685.

Gutschick, RC, 1954, Sponge spicules from the Lower Mississippian
of Indiana and Kentucky: American Midland Naturalist, v. 52,
no. 2, p. 501-509.

Haney, D.C., 1976, Structural control on Paleozoic sedimentation in
eastern Kentucky [abs.]: American Association of Petroleum
Geologists Bulletin, v. 60, p. 1620.

_1979, The Mississippian-Pennsylvanian systemic boundary in
eastern Kentucky: Southeastern Geology, v. 21, p. 53-62.
1980. The Mississippian-Pennsylvanian systemic boundary in

eastern Kentucky—Reply: Southeastern Geology, v. 21,11 301-303.

Haney, DC, and Hester, N.C., 1978, Geologic map of the Scranton
quadrangle, Menifee County, Kentucky: US. Geological Survey
Geologic Quadrangle Map GQ-1488.

Haney, D.C., Hester, N.C., and Hoge, ER, 1975, Evidence for fault-
controlled deposition of Pennsylvanian sediments in central
eastern Kentucky: Geological Society of America Abstracts with
Programs, v. 7, p. 495—496.

Haney, DC, and Rice, C.L., 1978, Geologic map of the Leighton quad-
rangle, east-central Kentucky: US. Geological Survey Geologic
Quadrangle Map GQ-1495.

Harman, A.E., 1975, A sedimentary and petrographic study of the
Fort Payne Formation, Lower Mississippian (Osage), in south-
central Kentucky: Cincinnati, Ohio, University of Cincinnati M.S.
thesis, 111 p.

Hansen, DE, 1975, Geologic map of the Sebree quadrangle, Webster
County, Kentucky: US. Geological Survey Geologic Quadrangle
Map GQ-1238.

Hass, W.H., 1956, Age and correlation of the Chattanooga Shale and
the Maury Formation: US. Geological Survey Professional Paper
286, 47 p.

Hasson, K.O., 1973, Type and standard reference of the Grainger For-
mation (Mississippian), northeast Tennessee: Journal of Tennessee
Academy of Science, v. 48, no. 1, p. 17—22.

Hatch, N .L., J r., 1964, Geology of the Shopville quadrangle, Kentucky:
US. Geological Survey Geologic Quadrangle Map GQ—282.
Hauser, R.E., Walker, F.H., and Nelson, V.E., 1957, Itinerary—Some
stratigraphic and structural features of the Middlesboro basin;
Road log for field conference, Geological Society of Kentucky and
Appalachian Geological Society: Lexington, Ky., Kentucky Geo-

logical Survey, 35 p.

 

 

 

118

Haynes, DD, 1966, Geologic map of the Horse Cave quadrangle, Bar-
ren and Hart Counties, Kentucky: U.S. Geological Survey Geologic
Quadrangle Map GQ-558.

Hays, W.H., 1964, Geology of the Grand Rivers quadrangle, Kentucky:
U.S. Geological Survey Geologic Quadrangle Map GQ-328.
Hetherington, P.A., 1981, Stratigraphic relationships and depositional
controls of the Mississippian (Meramecian and Lower Chesterian)
“Big Lime” in southeastern Kentucky-A subsurface petroliferous
equivalent of the lower and middle Newman Group: Lexington,

Ky., University of Kentucky M.S. thesis, 135 p.

Heyl, A.V., 1972, The 38th parallel lineament and its relationship to
ore deposits: Economic Geology, v. 67 , p. 879—894.

Hicks, L.E., 1878, The Waverly Group in central Ohio: American Jour-
nal of Science, 3rd ser., v. 16, p. 216-224.

Hoge, H.P.,1977a, Stop 2,in Dever, G.R., Jr., and others. Stratigraphic
evidence for late Paleozoic tectonism in northeastern Kentucky;
Field trip guidebook, Eastern Section, American Association of
Petroleum Geologists October 9, 1976: Lexington, Ky., Kentucky
Geological Survey, p. 57—59.

1977b,Geologic map of the Frenchburg quadrangle, east-central
Kentucky: U.S. Geological Survey Geologic Quadrangle map
GQ-1390.

Hoge, H.P., and Chaplin, J .R., 1972, Geologic map of the Morehead
quadrangle, Rowan County, Kentucky: U.S. Geological Survey
Geologic Quadrangle Map GQ-1022.

Hopkins, T.C., and Siebenthal, GE, 1897, The Bedford oolitic lime-
stone of Indiana: Indiana Department of Geology and Natural
Resources, let Annual Report, p. 291—427.

Home, J .C., and Perm, J .C., 1970, Facies relationships of the
Mississippian-Pennsylvanian contact in northeastern Kentucky:
Geological Society of America Abstracts with Programs, v. 2,
p. 217.

__1978, Carboniferous depositional environments—eastern Ken-
tucky and southern West Virginia: Columbia, SC, University of
South Carolina Department of Geology, 151 p.

Horne, J .C., Ferm, J .C., and Swinchatt, J .P., 1974, Depositional model
for the Mississippian-Pennsylvanian boundary in northeastern
Kentucky, in Briggs, Garrett, ed., Carboniferous of southeastern
United States: Geological Society of America Special Paper 148,
p. 97-114.

Horne, J .C., Swinchatt, J .P., and Ferm, J .C., 1971, LeeNewman bar-
rier shoreline model. in Ferm, J .C., and others, Carboniferous
depositional environments in northeastern Kentucky—Annual
Spring Field Conference, Geological Society of Kentucky, April
23, 24, 1971: Lexington, Ky., Kentucky Geological Survey, p. 5—9.

Horowitz, AS, 1965, Crinoids from the Glen Dean Limestone (Middle
Chester) of southern Indiana and Kentucky: Indiana Geological
Survey Bulletin 34, 52 p.

Horowitz, A.S., Mamet, B.L., Neves, Roger, Potter, RE, and Rex-
road, 03., 1979, Carboniferous paleontological zonation and in-
tercontinental correlation of the Fowler no. 1 Traders core, Scott
County, Tennessee, U.S.A.: Southeastern Geology, v. 20,
p. 205-228.

Horowitz, A.S., and Rexroad, 0.13., 1972, Conodont biostratigraphy
of some United States Mississippian sites: Journal of Paleon-
tology, v. 46, no. 6, p. 884—891.

Horowitz, AS, and Strimple, H.L., 1974, Chesterian echinoderm zona-
tion in eastern United States: Septieme Congres International de
Stratigraphie et de Géologie du Carbonifere, Compte Rendu, band
III, p. 207—220.

Huddle. J .W., 1934, Conodonts from the New Albany Shale of Indiana:
Bulletins of American Paleontology, v. 21, no. 72, p. 1-136.

Hyde, J .E., 1911, The ripples of the Bedford and Berea formations
of central and southern Ohio, with notes on the paleogeography
of that epoch: Journal of Geology, v. 19, p. 257—269.

 

 

MISSISSIPPIAN ROCKS IN KENTUCKY

 

1915, Stratigraphy of the Waverly formations of central and
southern Ohio: Journal of Geology. v. 23, no. 7, p. 655-682; no.
a, p. 757—779.

Hylbert, D.K., and Philley, J .C., 1971, Geologic map of the Bangor
quadrangle, east-central Kentucky: U.S. Geological Survey
Geologic Quadrangle Map GQ-947.

Jennings, J .R.. 1977, Recognition of Upper Mississippian strata by
means of plant megafossils: Geological Society of America
Abstracts with Programs, v. 9, p. 610.

Johnson, W.D., Jr., 1977, Geologic map of the Falls of Rough quad-
rangle, western Kentucky: U.S. Geological Survey Geologic Quad-
rangle Map GQ—1360.

Jorgensen, D.B., and Carr, DD, 1973, Inﬂuence of cyclic deposition,
structural features, and hydrologic controls on evaporite deposits
in the St. Louis Limestone in southwestern Indiana, in Pro-
ceedings, 8th forum on geology of industrial minerals: Iowa
Geological Survey Public Information Circular 5, p. 43—65.

Kammer, T.W., 1982, Fossil communities of the prodeltaic New Prov-
idence Shale Member of the Borden Formation (Mississippian).
north-central Kentucky and southern Indiana: Bloomington, 1nd,,
Indiana University Ph. D. dissertation, 301 p.

Keith, Arthur, 1895, Geology of Chilhowee Mountain, Tennessee:
Philosophical Society of Washington Bulletin, v. 12, p. 71—88.

Kepferle, R.C., 1963a, Geology of the Howe Valley quadrangle, Ken-
tucky: U.S. Geological Survey Geologic Quadrangle Map GQ-232.

_1963b, Geology of the Cecilia quadrangle, Kentucky: U.S.
Geological Survey Geologic Quadrangle Map GQ—263.

1966a, Geologic map of the Howardstown quadrangle, central
Kentucky: U.S. Geological Survey Geologic Quadrangle Map
GQ—505.

_1966b, Geologic map of the Elizabethtown quadrangle, Hardin
and Larue Counties, Kentucky: U.S. Geological Survey Geologic
Quadrangle Map GQ—559.

1967 , Geologic map of the Colesburg quadrangle, Hardin and
Bullitt Counties, Kentucky: U.S. Geological Survey Geologic
Quadrangle Map GQ-602.

_1968, Geologic map of the Shepherdsville quadrangle, Bullitt
County, Kentucky: U.S. Geological Survey Geologic Quadrangle
Map GQ—740.

_1971, Members of the Borden Formation (Mississippian) in
north-central Kentucky: U.S. Geological Survey Bulletin 1354-3,
p. B1-B18.

1972a, Stratigraphy, petrology, and depositional environment

of the Kenwood Siltstone Member, Borden Formation (Mississip-

pian), Kentucky and Indiana: U.S. Geological Survey Open-file
report; Cincinnati, Ohio, University of Cincinnati Ph. D. thesis,

233 p.

1972b, Geologic map of the Brooks quadrangle, Bullitt and J ef-

ferson Counties, Kentucky: U.S. Geological Survey Geologic Quad-

rangle Map GQ-961.

1973, Geologic map of the Raywick quadrangle, central Ken-

tucky: U.S. Geological Survey Geologic Quadrangle Map GQ-1048.

1974, Geologic map of parts of the Louisville West and Lanes-
ville quadrangles, Jefferson County, Kentucky: U.S. Geological
Survey Geologic Quadrangle Map GQ-1202.

_1977a, Stratigraphy, petrology, and depositional environment
of the Kenwood Siltstone Member, Borden Formation (Mississip-
pian), Kentucky and Indiana: U.S. Geological Survey Professional
Paper 1007, 49 p.

1977b, Geologic map of the Pitts Point quadrangle, Bullitt and
Hardin Counties, Kentucky: U.S. Geological Survey Geologic
Quadrangle Map GQ—1376.

____1979a, Delta front of the Borden Formation in central Ken-
tucky, in Ettensohn, F.R., and Dever, G.R., Jr., eds., Carboniferous
geology from the Appalachian basin to the Illinois basin through

 

 

 

 

 

 

REFERENCES CITED

eastern Ohio and Kentucky, Field trip no. 4, 9th International

Congress of Carboniferous Stratigraphy and Geology: Lexington,

Ky., University of Kentucky, p. 239—251.

1979b, Glauconite laminae in Floyds Knob Bed of Muldraugh
Member of Borden Formation show convergence along Borden
delta front, in Ettensohn, RR, and Dever, G.R., Jr., eds., Car-
boniferous geology from the Appalachian basin to the Illinois basin
through eastern Ohio and Kentucky, Field trip no. 4, 9th Inter-
national Congress of Carboniferous Stratigraphy and Geology:
Lexington, Ky., University of Kentucky, p. 253—254.

Kepferle, R.C., Lewis. R.Q., Sr., and Taylor, A.R., 1974, Kniﬂey Sand-
stone and Cane Valley Limestone—Two new members of the Fort
Payne Formation (Lower Mississippian) in south-central Ken-
tucky: U.S. Geological Survey Bulletin 1394-A, p. A63-A70.

Kepferle, R.C., and Peterson, W.L., 1964, Lithologic and radioactivi-
ty log of Bell Scott quarry drill hole, Meade County, Kentucky:
U.S. Geological Survey Open-file report, 1 sheet.

Kepferle, R.C., Peterson, W.L., and Sable, E.G., 1964, Road log, field
trip no. 2, in Pohl, E.R., and others, Itinerary—Geologic features
of the Mississippian Plateaus in the Mammoth Cave and Eliza-
bethtown areas, Kentucky—Annual spring field conference of
Geological Society of Kentucky: Lexington, Ky., Kentucky Geo-
logical Survey, p. 24—32.

Kepferle, R.C., Pryor, W.A., Maynard, J .B., and Harrell, J ames,'1980,
J abez Sandstone Member—A new member of the Fort Payne For-
mation (Mississippian), south-central Kentucky: U.S. Geological
Survey Bulletin 1502-A, p. A104—A109.

Kepferle, R.C., and Sable, E.G., 1977, Geologic map of the Fort Knox
quadrangle, north-central Kentucky: U.S. Geological Survey Geo-
logic Quadrangle Map GQ-1375.

Kepferle, R.C., Wilson, E.N., and Ettensohn, FR, 1978, Preliminary
stratigraphic cross section showing radioactive zones in the Devo-
nian black shales in the southern part of the Appalachian basin:
U.S. Geological Survey Oil and Gas Investigations Chart OC-85.

Kindle, E.M., 1899, The Devonian and Lower Carboniferous faunas
of southern Indiana and central Kentucky: Bulletins of American
Paleontology, v. 3, no. 12, p. 3—111.

King, RB, 1969, Tectonic map of North America: U.S. Geological
Survey, scale 1:5,000,000.

Klein, H.M., 1974, The Beaver Creek Limestone sediment gravity ﬂow
fan deposits, Fort Payne Formation, Lower Mississippian
(Osagean), south-central Kentucky, U.S.A.: Cincinnati, Ohio,
University of Cincinnati M.S. thesis, 121 p.

Klekamp, C.T., 1971, Petrology and paleocurrents of the Ste. Gene—
vieve Member of the Newman Limestone (Mississippian) in Carter
County, northeastern Kentucky: Cincinnati, Ohio, University of
Cincinnati M.S. thesis, 38 p.

Klemic, Harry, 1966a, Geologic map of the Oak Grove quadrangle,
Kentucky-Tennessee: U.S. Geological Survey Geologic Quadrangle
Map GQ—565.

1966b, Geologic map of the Herndon quadrangle, Kentucky-

Tennessee: U.S. Geological Survey Geologic Quadrangle Map

GQ-572.

1967, Geologic map of the Hopkinsville quadrangle, Christian
County, Kentucky: U.S. Geological Survey Geologic Quadrangle
Map GQ-651.

Klemic, Harry, and Ulrich, GE, 1967, Geologic map of the Roaring
Spring quadrangle, Kentucky-Tennessee: U.S. Geological Survey
Geologic Quadrangle Map GQ-658.

Lambert, T.W., and MacCary, L.M., 1964. Geology of the Briensburg
quadrangle, Kentucky: U.S. Geological Survey Geologic Quad-
rangle Map GQ-327.

Laudon, LR, 1948, OsageMeramec contact, in Weller, J .M., ed., Sym-
posium on problems of Mississippian stratigraphy and correla-
tion: Journal of Geology, v. 56. no. 4, p. 288-302.

 

 

 

 

119

Lewis, R.Q., Sr., 1963, Geologic features of the Mississippian Plateau,
south-central Kentucky—Annual Spring Meeting of the Geological
Society of Kentucky: Lexington, Ky., Kentucky Geological
Survey, 28 p.

1967 , Geologic map of the Dubre quadrangle, southern Ken-

tucky: U.S. Geological Survey Geologic Quadrangle Map GQ-67 6.

1971a, Geologic map of the Delmer quadrangle, Pulaski Coun-

ty, Kentucky: U.S. Geological Survey Geologic Quadrangle Map

GQ-909.

1971b, The Monteagle Limestone of south-central Kentucky:

U.S. Geological Survey Bulletin 1324-E, 10 p.

1975, Geologic map of the Frazer quadrangle, Pulaski and
Wayne Counties, Kentucky: U.S. Geological Survey Geologic
Quadrangle Map GQ-1223.

Lewis, R.Q., Sr., and Potter, P.E., 1978, Surface rocks in the western
Lake Cumberland area. Clinton, Russell, and Wayne Counties,
Kentucky—Annual field conference of the Geological Society of
Kentucky: Lexington, Ky., Kentucky Geological Survey, 41 p.

Lewis, R.Q., Sr., and Taylor, A.R., 1971, Geologic map of the Houston-
ville quadrangle, Casey and Lincoln Counties, Kentucky: U.S.
Geological Survey Geologic Quadrangle Map GQ-916.

1975, The Science Hill Sandstone Member of the Warsaw For-

mation, south-central Kentucky: U.S. Geological Survey Bulletin

1405-A, p. A28-A30.

1976, Geologic map of the Coopersville quadrangle, Wayne and

McCreary Counties, Kentucky: U.S. Geological Survey Geologic

Quadrangle Map GQ-1315.

1979, The Science Hill Sandstone Member of the Warsaw For-
mation and its relation to other elastic units in south-central Ken-
tucky: U.S. Geological Survey Bulletin 1435-D, 15 p.

Lewis, R.Q., Sr., Taylor, A.R., and Weir, G.W., 1973, Geologic map
of the Eubank quadrangle, south-central Kentucky: U.S. Geolog-
ical Survey Geologic Quadrangle Map GQ-1096.

Lewis, R.Q., Sr., and Thaden, R.E., 1962, Geology of the Wolf Creek
Dam quadrangle, Kentucky: U.S. Geological Survey Geologic
Quadrangle Map GQ-177.

1965, Geologic map of the Cumberland City quadrangle,

southern Kentucky: U.S. Geological Survey Geologic Quadrangle

Map GQ-475.

1966, Geologic map of the Albany quadrangle. Kentucky-
Tennessee: U.S. Geological Survey Geologic Quadrangle Map
GQ—550.

Lineback, J .A., 1964, Stratigraphy and depositional environments of
the New Albany Shale in Indiana: Bloomington, Ind., Indiana Uni-
versity Ph. D. thesis, 136 p.

_1966, Deep-water sediments adjacent to the Borden Siltstone
(Mississippian) delta in southern Illinois: Illinois State Geological
Survey Circular 401, 48 p.

_1968a. Subdivisions and depositional environments of New
Albany Shale (Devonian-Mississippian) in Indiana: American Asso-
ciation of Petroleum Geologists Bulletin, v. 52, no. 7, p. 1291-1303.

1968b, Turbidites and other sandstone bodies in the Borden

Siltstone (Mississippian) in Illinois: Illinois State Geological

Survey Circular 425, 29 p.

1969, Illinois Basin—Sediment—starved during Mississippian:
American Association of Petroleum Geologists Bulletin, v. 53, no.
1, p. 112—126.

_1970, Stratigraphy of the New Albany Shale in Indiana: Indiana
Geological Survey Bulletin 44, 73 p.

1972, Lateral gradation of the Salem and St. Louis Limestone
(Middle Mississippian) in Illinois: Illinois State Geological Survey
Circular 474, 23 p.

Livesay, Ann, 1953, Geology of the Mammoth Cave National Park
area: Kentucky Geological Survey, ser. 10, Special Publication 7,
40 p.; revised 1962 by Preston McGrain.

 

 

 

 

 

 

 

 

 

 

 

 

120

Lobeck, A.K., 1929, The geology and physiography of the Mammoth
Cave National Park: Kentucky Geological Survey, ser. 6, v. 31,
p. 327—399.

Malott, CA, 1919, The “America Bottoms” region of eastern Greene
County, Indiana—A type unit in southern Indiana physiography:
Indiana University Studies, v. 6, no. 40, p. 7-20.

1952, Stratigraphy of the Ste. Genevieve and Chester Forma-
tions of southern Indiana: Ann Arbor, Mich., The Edwards Letter
Shop, 105 p.

Malott, CA, and McGrain, Preston, 1977, A geologic profile of Sloans
Valley, Pulaski County, Kentucky: Kentucky Geological Survey,
ser. 10, Report of Investigations 20, 11 p.

Mason, CE, 1979, A new Lower Mississippian cephalopod fauna from
the Borden Formation, northeastern Kentucky: Urbana, 111.,
University of Illinois at Urbana-Champaign, 9th International
Congress of Carboniferous Stratigraphy and Geology Abstracts
of Papers, p. 131.

Mason, CE, and Chaplin, J .R., 1979, Nancy and Cowbell Members
of the Borden Formation, in Ettensohn, F.R., and Dever, G.R.,
J r., eds., Carboniferous geology from the Appalachian basin to
the Illinois basin through eastern Ohio and Kentucky, Field trip
no. 4, 9th International Congress of Carboniferous Stratigraphy
and Geology: Lexington, Ky., University of Kentucky, p. 147—151.

Maughan, E.K., 1976, Geologic map of the Roxana quadrangle,
Letcher and Harlan Counties, Kentucky: U.S. Geological Survey
Geologic Quadrangle Map GQ-1299.

Maxwell, CH, 1964, Geology of the Kniﬂey quadrangle, Kentucky:
U.S. Geological Survey Geologic Quadrangle Map GQ-294.
Maxwell, C.H., and Turner, W.B., 1964, Geology of the Cane Valley
quadrangle, Kentucky: U.S. Geological Survey Geologic Quad-

rangle Map GQ-369.

McDowell, RC, 1975, Geologic map of the Farmers quadrangle, east-
central Kentucky: U.S. Geological Survey Geologic Quadrangle
Map GQ-1236.

1978, Geologic map of the Levee quadrangle, east-central Ken-
tucky: U.S. Geological Survey Geologic Quadrangle Map GQ—1478.

McDowell, R.C., ed., 1986, The geology of Kentucky—A text to ac-
company the geologic map of Kentucky: U.S. Geological Survey
Professional Paper 1151-H, p. H1-H76.

McDowell, R.C., Grabowski, G.J., Jr., and Moore, S.L., 1981, Geologic
map of Kentucky: U.S. Geological Survey, 4 sheets, 1:250,000
scale.

McDowell, RC, and Weir, G.W., 1977, Geologic map of the Olympia
quadrangle, Bath and Menifee Counties, Kentucky: U.S. Geolog-
ical Survey Geologic Quadrangle Map GQ-1406.

McFarlan, A.C., 1943 (reprinted 1961b, Geology of Kentucky: Lex-
ington, Ky., University of Kentucky, 531 p.

1958, Behind the scenery in Kentucky: Kentucky Geological
Survey, ser. 9, Special Publication 10, 144 p.

McFarlan, A.C., Swann, D.H., Walker, F.H., and Nosow, Edmund,
1955, Some old Chester problems—Correlations of lower and mid-
dle Chester formations of western Kentucky: Kentucky Geological
Survey, ser. 9, Bulletin 16, 37 p.

McFarlan, A.C., and Walker, F.H., 1956, Some old Chester problems—
Correlations along the eastern belt of outcrop: Kentucky Geolog-
ical Survey, ser. 9, Bulletin 20, 36 p.

McGrain, Preston, 1954, Geology of the Carter and Cascade Caves
area: Kentucky Geological Survey, ser. 9, Special Publication 5,
32 p.; repr. 1966 in ser. 10 as Special Publication 12.

1967, The geologic story of Bernheim Forest: Kentucky

Geological Survey, ser. 10, Special Publication 13, 26 p.

1969, Extension of Lost River Chert across parts of Kentucky:
American Association of Petroleum Geologists Bulletin, v. 53, no.
7, p. 1506-1507.

McGrain, Preston, and Dever, G.R., J r., 1967, Limestone resources

 

 

 

 

 

 

MISSISSIPPIAN ROCKS IN KENTUCKY

in the Appalachian region of Kentucky: Kentucky Geological
Survey, ser. 10, Bulletin 4, 12 p.

McGrain, Preston, and Helton, W.L., 1964, Gypsum and anhydrite
in the St. Louis Limestone in northwestern Kentucky: Kentucky
Geological Survey, ser, 10, Information Circular 13, 26 p.

McGregor, D.J., 1954, Gypsum and anhydrite deposits in south-
western Indiana: Indiana Geological Survey Report of Progress
8, 24 p.

Meek, F.B., and Worthen, AH, 1861, Remarks on the age of the
Goniatite limestone at Rockford, Indiana, and its relation to the
black slate of the western states and to some of the succeeding
rocks above the latter: American Journal of Science, 2d ser., v.
32, p. 167-177.

Miller, A.M., 1919, The geology of Kentucky: Kentucky Department
of Geology and Forestry, ser. 5, Bulletin 2, 392 p.

Miller, M.S., 1974, Stratigraphy and coal beds of Upper Mississippian
and Lower Pennsylvanian rocks in southwestern Virginia: Virginia
Division of Mineral Resources Bulletin 84, 211 p.

Miller, R.L., 1964, The Little Stone Gap Member of the Hinton For-
mation (Mississippian) in southwest Virginia: U.S. Geological
Survey Professional Paper 501-B, p. BBQ-B42.

Miller, TR, 1963, Geologic map of the Sugar Grove quadrangle,
Kentucky: U.S. Geological Survey Geologic Quadrangle Map
GQ-225.

Moore, B.R., and Clarke, M.K., 1970, The significance of a turbidite
sequence in the Borden Formation (Mississippian) of eastern Ken-
tucky and southern Ohio, in Lajoie, J ., ed., Flysch sedimentology
in North America: Geological Association of Canada Special Paper
7, p. 211—218.

Moore, F.B.,1964a, Geologyof the Summit quadrangle, Kentucky: U.S.
Geological Survey Geologic Quadrangle Map GQ-298.

_____1964b, Lithologic and radioactivity log of Summit area drill hole,
Hardin County, Kentucky: U.S. Geological Survey Open-file
report, 1 sheet.

1976, Geology of the Hibernia quadrangle, central Kentucky:
U.S. Geological Survey Geologic Quadrangle Map GQ-1352.

Moore, RC, 1928, Early Mississippian formations in Missouri:
Missouri Bureau of Geology and Mines, 2d sen, v. 21, 283 p.

Moore, S.L., 1961, Geology of the Austin quadrangle, Kentucky: U.S.
Geological Survey Geologic Quadrangle Map GQ-173.

1976, Geologic Map of the Saloma quadrangle, central Ken—

tucky: U.S. Geological Survey Geologic Quadrangle Map GQ—1351.

1977, Geologic map of the Bradfordsville NE quadrangle, cen-
tral Kentucky: U.S. Geological Survey Geologic Quadrangle Map
GQ—1396.

Morris, R.H., 1965a, Geology of the Charters quadrangle, Kentucky:
U.S. Geological Survey Geologic Quadrangle Map GQ-293.
1965b, Geologic map of the Stricklett quadrangle, northeastern
Kentucky: U.S. Geological Survey Geologic Quadrangle Map

GQ-394.

1966a, Geologic map of the Head of Grassy quadrangle, Lewis

County, Kentucky: U.S. Geological Survey Geologic Quadrangle

Map GQ-484.

1966b, Geologic map of parts of the Concord and Buena Vista
quadrangles, Lewis County, Kentucky: U.S. Geological Survey
Geologic Quadrangle Map GQ-525.

Morris, R.H., and Pierce, KL, 1967 , Geologic map of the Vanceburg
quadrangle, Kentucky-Ohio: U.S. Geological Survey Geologic
Quadrangle Map GQ-598.

Myers, W.B., 1964, Geology of the Petroleum quadrangle, Kentucky-
Tennessee: U.S. Geological Survey Geologic Quadrangle Map
GQ-352.

Nelson, W.H., 1962, Geology of the Holland quadrangle, Kentucky-
Tennessee: U.S. Geological Survey Geologic Quadrangle Map
GQ-174.

 

 

 

 

 

 

REFERENCES CITED

N ewberry, J .S., 1870, Report on the progress of the Geological Survey
of Ohio in 1869: Ohio Geological Survey Report of Progress 1869,
pt. 1, p. 3-51.

Newell. W.L., 1975, Geologic map of the Frakes quadrangle and part
of the Eagan quadrangle. southeastern Kentucky: U.S. Geological
Survey Geologic Quadrangle Map GQ-1249.

Nicoll, RS, 1971, Stratigraphy and conodont paleontology of the
Sanders Group (Mississippian) in Indiana and adjacent Kentucky:
Iowa City, Iowa, University of Iowa Ph. D. thesis, 81 p.

Nicoll, KS, and Rexroad, CD. 1975, Stratigraphy and conodont
paleontology of the Sanders Group (Mississippian) in Indiana and
adjacent Kentucky: Indiana Geological Survey Bulletin 51, 36 p.

Olive, W.W.. 1965, Geology of the Rice quadrangle, Kentucky: U.S.
Geological Survey Geologic Quadrangle Map GQ-332.

Palmer. J .E., 1978, Geologic map of the Guston quadrangle, Meade
and Breckenridge Counties. Kentucky: U.S. Geological Survey
Geologic Quadrangle Map GQ-1481.

Patterson, SH, and Hostel-man, J .W., 1961, Geology of the Haldeman
quadrangle, Kentucky: U.S. Geological Survey Geologic Quad-
rangle Map GQ-169.

Patton. J.B., 1949. Crushed stone in Indiana: Indiana Geological
Survey Report of Progress 3, 47 p.

Pear, J .L.. 1980. Factors controlling stratigraphic variation within
the Mississippian “Big Lime” of eastern Kentucky—A petrolifer-
ous equivalent of the lower Newman Group (Meramecian. Lower
Chesterian): Lexington, Ky., University of Kentucky M.S. thesis.
98 p.

Pepper, J .F., de Witt. Wallace. Jr., and Demarest, D.F., 1954, Geology
of the Bedford Shale and Berea Sandstone in the Appalachian
basin: U.S. Geological Survey Professional Paper 259, 111 p.

Perry, T.G., and Smith. N .M.. 1958, The Meramec-Chester and intra-
Chester boundaries and associated strata in Indiana: Indiana
Geological Survey Bulletin 12, 110 p.

Peterson, M.N.A., 1956, Stratigraphy and petrog'raphy of the (Up-
per Mississippian) Big Clifty Formation of Tennessee: Evanston.
11].. Northwestern University Ph. D. dissertation.

Peterson, W.L., 1964, Geologic map of the Big Spring quadrangle,
Kentucky: U.S. Geological Survey Geologic Quadrangle Map
GQ-261.

___1966. Geologic map of the New Haven quadrangle, Nelson and
Larue Counties, Kentucky: U.S. Geological Survey Geologic Quad-
rangle Map GQ—506.

Peterson, W.L., and Kepferle. R.C., 1970, Deltaic deposits of the
Borden Formation in central Kentucky: U.S. Geological Survey
Professional Paper 700-D, p. D49-D54.

Philley, J .C., 1970. Paleozoic section on east ﬂank of Cincinnati Arch
along Interstate 64, Lexington to Olive Hill. Kentucky. Part 11:
Valley of Licking River eastward to Olive Hill Interchange;
Guidebook for ﬁeld trips, 18th annual meeting, Southeastern Sec-
tion, Geological Society of America: Lexington, Ky., Kentucky
Geological Survey, p. 56—70.

1978, Geologic map of the Salt Lick quadrangle, east-central
Kentucky: U.S. Geological Survey Geologic Quadrangle Map
GQ-1499.

Philley. J .C., Hylbert, D.K., and Hoge, HR, 1974, Geologic map of
the (kanston quadrangle, northeastern Kentucky: U.S. Geological
Survey Geologic Quadrangle Map GQ-1212.

1975, Geologic map of the Soldier quadrangle, northeastern Ken-
tucky: U.S. Geological Survey Geologic Quadrangle Map GQ-1233.

Pinsak, A.P., 1957. Subsurface stratigraphy of the Salem Limestone
and associated formations in Indiana: Indiana Geological Survey
Bulletin 11. 62 p.

Pipiiingos, G.N., Bergman, 8.0., and Trent, V.A.. 1968, Geologic map
of the Ezel quadrangle, Morgan and Menifee Counties, Kentucky:
U.S. Geological Survey Geologic Quadrangle Map GQ—721.

 

 

 

121

Pohl, E.R.. 1970. Upper Mississippian deposits of south-central Ken-
tucky, a project report: Transactions of the Kentucky Academy
of Science, v. 31. p. 1—15.

Pohl, E.R., Browne. R.G.. and Chaplin, J .R., 1968, Foraminifera of
the Fraileys Member (Upper Mississippian) of central Kentucky:
Journal of Paleontology. v. 42, no. 2, p. 581-582.

Pohl, E.R., and Philley, J .C., 1971, Age and stratigraphy of Upper
Mississippian carbonates of northeastern Kentucky: Geological
Society of America Abstracts with Programs, v. 3, p. 340.

Potter, RE, 1962, Late Mississippian sandstones of Illinois: Illinois
State Geological Survey Circular 340. 36 p.

_1963. Late Paleozoic sandstones of the Illinois basin: Illinois
State Geological Survey Report of Investigations 217, 92 p.

1978. Structure and isopach map of the New Albany-
Chattanooga—Ohio Shale (Devonian and Mississippian) in
Kentucky—Central sheet: Kentucky Geological Survey, ser. 10,
scale 1:250,000.

Potter. P.E., Nosow, Edmund, Smith, N .M.. Swann. D.H., and Walker,
ER. 1958. Chester cross-bedding and sandstone trends in the
Illinois basin: American Association of Petroleum Geologists
Bulletin, v. 42. no. 5, p. 1013-1046.

Potter, RE, and Pryor, W.A., 1961, Dispersal centers of Paleozoic
and later clastics of the upper Mississippi Valley and adjacent
areas: Geological Society of America Bulletin. v. 72, no. 8,
p. 1195—1249.

Pryor, W.A., and Sable, E.G.. 1974, Carboniferous of the Eastern In-
terior basin, in Briggs, Garrett. ed., Carboniferous of the south-
eastern United States: Geological Society of America Special
Paper 148, p. 281—313.

Rainey, H.C., III, 1963, Geology of the Hadley quadrangle, Kentucky:
U.S. Geological Survey Geologic Quadrangle Map GQ-237.
1964, Geology of the Rockfield quadrangle, Kentucky: U.S.

Geological Survey Geologic Quadrangle Map GQ-309.

Reger, D.B., 1926, Mercer. Monroe, and Summers Counties: West
Virginia Geological Survey (County Report), 963 p.

Rexroad. C.B.. 1957 . Conodonts from the Chester series in the type
area of southwestern Illinois: Illinois State Geological Survey
Report of Investigations 199, 43 p.

1958, Conodonts from the Glen Dean Formation (Chester) of

the Illinois basin: Illinois State Geological Survey Report of In-

vestigations 209. 27 p.

1969. Conodonts from the Jacobs Chapel Bed (Mississippian)
of the New Albany Shale in southern Indiana: Indiana Geological
Survey Bulletin 41. 55 p.

Rexroad, C.B., and Collinson, Charles, 1963, Conodonts from the St.
Louis Formation (Valmeyeran Series) of Illinois, Indiana, and
Missouri: Illinois Geological Survey Circular 355. 28 p.

Rexroad. C.B., and Liebe. R.M., 1962, Conodonts from the Paoli and
equivalent formations in the Illinois basin: Micropaleontology, v.
8, no. 4. p. 509—514.

Rexroad, C.B., and Scott. A.J., 1964. Conodont zones in the Rockford
Limestone and the lower part of the New Providence Shale (Mis-
sissippian) in Indiana: Indiana Geological Survey Bulletin 30. 54 p.

Reynolds, D.W., and Vincent. J .K., 1967, Western Kentucky’s Bethel
Channel—The largest continuous reservoir in the Illinois Basin,
in Rose, W.D., ed., Proceedings of the technical sessions. Kentucky
Oil and Gas Association 29th annual meeting. June 3—4, 1965:
Kentucky Geological Survey, ser. 10. Special Publication 14,
p. 19-30.

Rice. CL, 1972, Geologic map of the Alcorn quadrangle. east-central
Kentucky: U.S. Geological Survey Geologic Quadrangle Map
GQ-963.

1973. Geologic map of the Jenkins West quadrangle,
Kentucky—Virginia: U.S. Geological Survey Geologic Quadrangle
Map GQ-1126.

 

 

 

 

122

1980, The Mississippian-Pennsylvanian systemic boundary in
eastern Kentucky—Discussion: Southeastern Geology, v. 21,
p. 299-301.

1984, Sandstone units of the Lee Formation and related strata
in eastern Kentucky: U.S. Geological Survey Professional Paper
1151-G, 53 p.

_1985, Terrestrial vs. marine depositional model—A new assess-
ment of subsurface Lower Pennsylvanian rocks of southwestern
Virginia: Geology, v. 13, no. 11, p. 786—789.

Rice, C.L., and Maughan, E.K., 1978, Geologic map of the Kayjay
quadrangle and part of the Fork Ridge quadrangle, Bell and Knox
Counties, Kentucky: U.S. Geological Survey Geologic Quadrangle
Map GQ-1505.

Rice, CL, and Newell, W.L., 1975, Geologic map of the Saxton quad-
rangle and the J ellico East quadrangle, Kentucky: U.S. Geolog-
ical Survey Geologic Quadrangle Map GQ-1264.

Rice, C.L., Sable, E.G., Dever, G.R., Jr., and Kehn, T.M., 1979, The
Mississippian and Pennsylvanian (Carboniferous) Systems in the
United States—Kentucky: U.S. Geological Survey Professional
Paper 1110-F, 32 p.

Rice, CL, and Weir, G.W., 1984, Lee and Breathitt Formations along
the northwestern part of the eastern Kentucky coal ﬁeld, in Rice,
C.L., Sandstone units of the Lee Formation and related strata in
eastern Kentucky: U.S. Geological Survey Professional Paper
ll51-G, p. G27—G53.

Rice, CL, and Wolcott, DE, 1973, Geologic map of the Whitesburg
quadrangle, Kentucky-Virginia, and part of the Flat Gap
quadrangle, Letcher County, Kentucky: U.S. Geological Survey
Geologic Quadrangle Map GQ-1119.

Rich, J .L., 1951, Probable fondo origin of Marcellus—Ohio-New
Albany-Chattanooga bituminous shales: American Association
of Petroleum Geologists Bulletin, v. 35, no. 9, p. 2017—2040.

Rich, J .L., and Wilson, J .W., 1950, Paleogeographic and stratigraphic
significance of subaqueous ﬂow markings in Lower Mississippian
strata of south-central Ohio and adjacent parts of Kentucky [abs]:
Geological Society of America Bulletin, v. 61, no. 12, p. 1496.

Rogers, W.B., 1963, Geology of the Eddyville quadrangle, Kentucky:
US. Geological Survey Geologic Quadrangle Map GQ-255.

Rogers, W.B., and Hays, W.H., 1967, Geologic map of the Fredonia
quadrangle, western Kentucky: U.S. Geological Survey Geologic
Quadrangle Map GQ-607.

Ross, C.A., and Ross, J .R.P., 1985, Late Paleozoic depositional se-
quences are synchronous and worldwide: Geology, v. 13,
p. 194-197.

Rubey, W.W., 1952, Geology and mineral resources of the Hardin and
Brussels quadrangles (in Illinois): U.S. Geological Survey Profes-
sional Paper 218, 179 p.

Sable, E.G., 1964, Geology of the Constantine quadrangle, Kentucky:
U.S. Geological Survey Geologic Quadrangle Map GQ-302.
1970, Source areas of Lower Mississippian red beds in eastern
Midcontinent: U.S. Geological Survey Professional Paper 700-D,

p. D88—D91.

1979a, Eastern Interior Basin region, in Craig, L.C., and Connor,

C.W., coordinators, Paleotectonic investigations of the Missis-

sippian System in the United States: U.S. Geological Survey Pro-

fessional Paper 1010, pt. I, p. 59—106.

1979b, Evaporites in the Eastern Interior Basin, in de Witt,
Wallace, J r., Sable, E.G., and Cohee, G.V., Evaporite deposits in
Mississippian rocks of the eastern United States: U.S. Geological
Survey Professional Paper 1010-T, p. 432—437.

Sable, E.G., Kepferle, RC, and Peterson, W.L., 1966, Harrodsburg
Limestone in Kentucky: U.S. Geological Survey Bulletin 1224-1,
12 p.

Sable, E.G., and Peterson, W.L., 1966, Road log, Field trip No. 2,
Saturday, April 23, 1966, in Geological features of selected

 

 

 

 

 

 

MISSISSIPPIAN ROCKS IN KENTUCKY

Pennsylvanian and Mississippian channel deposits along the
eastern rim of the western Kentucky coal basin; Annual Spring
Field Conference of Geological Society of Kentucky [Guidebook]:
Lexington, Ky., Kentucky Geological Survey, p. 20—32.

Safford, J .M., and Killebrew, J .B., 1900, The elements of the geology
of Tennessee: Nashville, Tenn., Foster and Webb, 264 p.

Sample, RD, 1965, Geology of the Princeton West quadrangle, Ken-
tucky: U.S. Geological Survey Geologic Quadrangle Map GQ—385.

Sandberg, C.A., 1981, The Devonian—Carboniferous boundary in North
America: Geological Society of America Abstracts with Programs,
v. 13, p. 315.

Sandberg, C.A., and Bowles, CG, 1965, Geology of the Cub Run quad-
rangle, Kentucky: U.S. Geological Survey Geologic Quadrangle
Map GQ-386.

Sando, W.J., 1983, Revision of Lithostrotionella (Coelenterata, Rugosa)
from the Carboniferous and Permian: U.S. Geological Survey Pro-
fessional Paper 1247 , 47 p.

Sando, W.J., Mamet, B.L., and Dutro, J .T., Jr., 1969, Carboniferous
megafaunal and microfaunal zonation in the northern Cordillera
of the United States: U.S. Geological Survey Professional Paper
613-E, 29 p.

Savage, T.E., 1920, The Devonian formations of Illinois: American
Journal of Science, 4th ser., v. 49, p. 169-182.

Savage, TE, and Sutton, A.H., 1931, Age of the black shale in south-
central Kentucky: American Journal of Science, ser. 5, v. 22,
p. 441-448.

Saxby, D.B., and Lamar, J .E., 1957, Gypsum and anhydrite in Illinois:
Illinois State Geological Survey Circular 226, 26 p.

Schlanger, 8.0., 1965, Geology of the Maretburg quadrangle, Ken-
tucky: U.S. Geological Survey Geolog'c Quadrangle Map GQ~338.

Schlanger, 8.0., and Weir, G.W., 1971, Geologic map of the Mount
Vernon quadrange, Rockcastle County, Kentucky: U.S. Geological
Survey Geologic Quadrangle Map GQ-902.

Schnabel, R.W., and MacKallor, J .A., 1964, Geology of the Fenton
quadrangle, Kentucky: U.S. Geological Survey Geology Quad-
rangle Map GQ—317.

Schwalb, H.R., 1969, Paleozoic geology of the Jackson Purchase
region, Kentucky, with reference to petroleum possibilities: Ken-
tucky Geological Survey, ser. 10, Report of Investigations 10, 40 p.

Schwalb, H.R., and Potter, RE, 1978, Structure and isopach map
of the New Albany-Chattanooga—Ohio Shale (Devonian and
Mississippian) in Kentucky—Western sheet: Kentucky Geological
Survey, ser. 10, scale 1:250,000.

Sedimentation Seminar, 1966, Crossbedding in the Salem Limestone
of central Indiana: Sedimentology, v. 6. p. 95—114.

___1969, Bethel Sandstone (Mississippian) of western Kentucky
and south-central Indiana, a submarine-channel fill: Kentucky
Geological Survey, set. 10, Report of Investigations 11, 24 p.

___1972, Sedimentology of the Mississippian Kniﬂey Sandstone
and Cane Valley Limestone in south-central Kentucky: Kentucky
Geological Survey, ser. 10, Report of Investigations 13, 30 p.

__1978, Sedimentology of the Kyrock Sandstone (Pennsylvanian)
in the Brownsville paleovalley, Edmonson and Hart Counties, Ken-
tucky: Kentucky Geological Survey, ser. 10, Report of Investiga-
tions 21, 24 p.

1981, Mississippian and Pennsylvanian section on Interstate
75 south of J ellico, Campbell County, Tennessee: Tennessee Divi-
sion of Geology Report of Investigations 38, 42 p.

Seeland, D.A., 1968, Geologic map of the Cobb quadrangle, 'I‘rigg and
Caldwell Counties, Kentucky: U.S. Geological Survey Geologic
Quadrangle Map GQ-710.

Sergeant, RE, and Haney, D.C., 1980, Stratigraphic evidence for late
Paleozoic reactivation of subsurface tectonic elements in north-
eastern Kentucky: Geological Society of America Abstracts with
Programs, v. 12, p. 208.

 

REFERENCES CITED

Shaver, R.H., and others, 1970, Compendium of rock-th stratigraphy
in Indiana: Indiana Geological Survey Bulletin 43, 229 p.
Shawe, F.R., and Gildersleeve, Benjamin, 1969, An anastomosing
channel complex at the base of the Pennsylvanian System in
western Kentucky: U.S. Geological Survey Professional Paper

650—D, p. D206—D209.

Shepherd, R.G., Pashin, J .C., Greb. SE, and Rice, C.L., 1986, Com-
ment and reply on “Terrestrial vs. marine depositional model—
A new assessment of subsurface Lower Pennsylvanian rocks of
southwestern Virginia”: Geology, v. 14, no. 9, p. 800—802.

Sheppard, R.A., 1964, Geology of the Portsmouth quadrangle,
Kentucky-Ohio. and parts of the Wheelersburg and New Boston
quadrangles, Kentucky: U.S. Geological Survey Geologic Quad-
rangle Map GQ-312.

1964, Geology of the Tygarts Valley quadrangle, Kentucky:
U.S. Geological Survey Geologic Quadrangle Map GQ-289.
Short, M.R., 1978, Petrology of the Pennington and Lee Formations
of northeastern Kentucky and the Sharon Conglomerate of south-
eastern Ohio: Cincirmati, Ohio, University of Cincinnati Ph. D.

thesis, 219 p.

Shumard, BR, 1860, Observations on the geology of the County of
St. Genevieve: St. Louis Academy of Science Transactions, v. 1,
p. 404—415.

Siever, Raymond, 1951, The Mississippian-Pennsylvanian uncon-
formity in southern Illinois: American Association of Petroleum
Geologists Bulletin, v. 35, p. 542—581. (Reprinted as Illinois State
Geological Survey Report of Investigations 152.?

Simmons, G.C., 1967, Geologic map of the Clay City quadrangle,
Powell and Estill Counties, Kentucky: U.S. Geological Survey
Geologic Quadrangle Map GQ-663.

Sloss, L.L., 1963, Sequences in the cratonic interior of North America:
Geological Society of America Bulletin, v. 74, p. 93—113.

Sloss, L.L., Krumbein, W.C., and Dapples, E.C., 1949, Integrated fades
analysis, in Longwell, C.R., chm., Sedimentary facies in geologic
history [symposium]: Geological Society of America Memoir 39,
p. 91-123.

Smith, E.A., 1890, Geological structure and description of the valley
regions adjacent to the Cahaba coal field: Alabama Geological
Survey Special Report 2, pt. 2, p. 137-180.

1894, Geological map of Alabama: Alabama Geological Survey
Map 1.

Smith, J .H., 1978, Geologic map of the Bell Farm quadrangle and part
of the Barthel] SW quadrangle, McCreary and Wayne Counties,
Kentucky: U.S. Geological Survey Geologic Quadrangle Map
GQ-1496.

Smith, M.O., Noger, M.O., and Smith, G.E., 1967, Some aspects of
the stratigraphy of the Pine Mountain front near Elkhorn City,
Kentucky, with notes on pertinent structural features, Annual
Spring Field Conference, Geological Society of Kentucky: Lex-
ington, Ky., Kentucky Geological Survey, 24 p.

Smith, N.M., 1965, The Sanders Group and subj acent Muldraugh For-
mation (Mississippian) in Indiana: Indiana Geological Survey
Report of Progress 29, 20 p.

Stearns, R.G., 1963, Monteagle Limestone, Hartselle Formation, and
Bangor Limestone—A new Mississippian nomenclature for use
in middle Tennessee, with a history of its development: Tennessee
Division of Geology Information Circular 11, 18 p.

Stockdale, RB, 1929, Stratigraphic units of the Harrodsburg Lime
stone: Indiana Academy of Science Proceedings, v. 38, p. 233-242.

___.1931, The Borden (Knobstone) rocks of southern Indiana: In-
diana Department of Conservation, Division of Geology, Publica-
tion 98, 330 p.

1939. Lower Mississippian rocks of the east-central interior:
Geological Society of America Special Paper 22, 248 p.

Stokley, J.A., and McFarlan, A.C., 1952, Industrial limestones of

 

 

 

 

123

Kentucky no. 2: Kentucky Geological Survey, ser. 9, Report of
Investigations 4, 95 p.

Stouder, R.E., 1938, Chester rocks of Meade, Hardin. and Breckinridge
Counties, Kentucky: American Association of Petroleum Geolo-
gists Bulletin, v. 22, p. 267-284.

.______1941, Geology of the Big Clifty quadrangle: Kentucky Depart-
ment of Mines and Minerals, Geological Division, ser. 8, Bulletin
7, 72 p.

Sutton, A.H., and Weller, J .M., 1932, Lower Chester correlations in
western Kentucky and Illinois: Journal of Geology, v. 40,
p. 430-442.

Swadley, WC, 1962, Geology of the Big Clifty quadrangle, Kentucky:
U.S. Geological Survey Geologic Quadrangle Map GQ-192.
Swann, D.H., 1963, Classiﬁcation of Genevievian and Chesterian (Late
Mississippian) rocks of Illinois: Illinois State Geological Survey

Report of Investigations 216, 91 p.

_1964, Late Mississippian rhythmic sediments of Mississippi
Valley: American Association of Petroleum Geologists Bulletin,
v. 48, p. 637-658.

Swann, D.H., and Bell, A.H., 1958, Habitat of oil in the Illinois Basin,
in Weeks, L.G., ed., Habitat of oil—A symposium: American
Association of Petroleum Geologists Bulletin, p. 447-472.

Swann, D.H., Lineback, J .A., and Frund, Eugene, 1965, The Borden
Siltstone (Mississippian) delta in southwestern Illinois: Illinois
State Geological Survey Circular 386, 20 p.

Swann, D.H., and Willman, H.B., 1961, Megagroups in Illinois: Amer-
ican Association of Petroleum Geologists Bulletin, v. 45,
p. 471—483.

Swinchatt, J .P., 1970, Mississippian carbonate facies in northeastern
Kentucky—Tidal ﬂat-island complexes on a tenigenous—carbonate
shelf: Geological Society of America Abstracts with Programs,
v. 2, p. 243.

Taylor, A.R., 1962, Geology of the Amandaville quadrangle, Kentucky:
U.S. Geological Survey Geologic Quadrangle Map GQ-186.
_1964, Geology of the Breeding quadrangle, Kentucky: U.S.

Geological Survey Geologic Quadrangle Map GQ—287.

1966, Geologic map of the Mannsville quadrangle, south-central

Kentucky: U.S. Geological Survey Geologic Quadrangle Map

GQ-562.

1976, Geologic map of the Monticello quadrangle, Wayne Coun-

ty, Kentucky: U.S. Geological Survey Geologic Quadrangle Map

GQ-1319.

1977 , Geologic map of the Parmleysville quadrangle and part
of the Sharp Place quadrangle, Wayne and McCreary Counties,
Kentucky: U.S. Geological Survey Quadrangle Map GQ-1405.

Taylor, A.R., and Lewis, R.Q., Sr., 1973, Geologic map of the Science
Hill quadrangle, south-central Kentucky: U.S. Geological Survey
Geologic Quadrangle Map GQ-1105.

Taylor, A.R., Lewis, R.Q., Sr., and Smith, J .H., 1975, Geologic map
of the Burnside quadrangle, south-central Kentucky: U.S. Geolog-
ical Survey Geologic Quadrangle Map GQ-1253.

Taylor, A.R., Luft, S.J., and Lewis, R.Q., Sr., 1968, Geologic map of
the Greensburg quadrangle, Green and Taylor Counties, Ken-
tucky: U.S. Geological Survey Geologic Quadrangle Map GQ-7 39.

Tazelaar, J .F., and Newell, W.L., 1974. Geologic map of the Evarts
quadrangle and part of the Hubbard Springs quadrangle,
southeastern Kentucky and Virginia: U.S. Geological Survey
Geologic Quadrangle Map GQ-914.

Thaden, RE, and Lewis, R.Q., Sr., 1962, Geology of the Jamestown
quadrangle, Kentucky: U.S. Geological Survey Geologic Quad-
rangle Map GQ-182.

1963, Geology of the Creelsboro quadrangle, Kentucky: U.S.

Geological Survey Geologic Quadrangle Map GQ-204.

1965, Geology of the Eli quadrangle, Kentucky: U.S. Geological

Survey Geologic Quadrangle Map GQ-393.

 

 

 

 

 

124

_1966, Geologic map of the J abez quadrangle, Russell and Wayne
Counties, Kentucky: US. Geological Survey Geologic Quadrangle
Map GQ-483.

_1969, Geologic map of the Faubush quadrangle, Pulaski and
Russell Counties, Kentucky: US. Geological Survey Geologic
Quadrangle Map GQ—802.

Thaden, R.E., Lewis, R.Q., Sr., Cattermole, J .M., and Taylor, A.R.,
1961, Reefs in the Fort Payne Formation of Mississippian age,
south-central Kentucky: US. Geological Survey Professional
Paper 424-B, p. B88—B90.

Thomas, W.A., 1972, Mississippian stratigraphy of Alabama: Alabama
Geological Survey Monograph 12, 121 p.

_1974, Converging clastic wedges in the Mississippian of
Alabama, in Briggs. Garrett, 'ed., Carboniferous of southeastern
United States: Geological Society of America Special Paper 148,
p. 187—207.

Thomas, W.A., and Mack, G.H., 1982, Paleogeographic relationship
of a Mississippian barrier-island and shelf-bar system (Hartselle
Sandstone) in Alabama to the Appalachian-Ouachita orogenic belt:
Geological Society of America Bulletin, v. 93, p. 6—19.

Thompson, TL, 1967, Conodont zonation of lower Osagean rocks
(Lower Mississippian) of southwestern Missouri: Missouri Geo-
logical Survey and Water Resources Report of Investigations 39,
88 p.

Thompson, TL, and Fellows, L.D., 1969, Stratigraphy and conodont
biostratigraphy of Kinderhookian and Osagean (Lower Mississip-
pian) rocks of southwestern Missouri and adjacent areas: Missouri
Geological Survey and Water Resources Report of Investigations
45, 263 p.

Tippie, RE, 1945, Rosiclare-Fredonia contact in and adjacent to
Hardin and Pope Counties, Illinois: American Association
Petroleum Geologists Bulletin, v. 29, p. 1654—1663. (Reprinted
as Illinois State Geological Survey Report of Investigations 1 12.)

Trace, R.D., 1962, Geology of the Salem quadrangle, Kentucky: US.
Geological Survey Geologic Quadrangle Map GQ-206.

1972, Geologic map of the Princeton East quadrangle, Caldwell

County, Kentucky: US. Geological Survey Geologic Quadrangle

Map GQ-1032.

1974, Geologic map of part of the Cave In Rock quadrangle,

Crittenden County, Kentucky: US. Geological Survey Geologic

Quadrangle Map GQ-1201.

1981, Middle Chesterian rocks in the Stevens Hill cut, Caldwell
County, Kentucky: Kentucky Geological Survey, ser. 11 (chart-
like format).

Trace, RD, and Amos, D.H., 1984, Stratigraphy and stmcture of the
western Kentucky ﬂuorspar district: US. Geological Survey Pro-
fessional Paper 1151-D, 41 p.

Trace, RD, and Kehn, T.M., 1968, Geologic map of the Olney quad-
rangle, Caldwell and Hopkins Counties, Kentucky: US Geolog-
ical Survey Geologic Quadrangle Map GQ—742.

Ulrich, E.O., 1904, Determination and correlation of formations (of
northern Arkansas): US. Geological Survey Professional Paper
24, p. 90—113.

1917, The formations of the Chester Series in western Kentucky
and their correlates elsewhere: Kentucky Geological Survey, ser.
4, v. 5, pt. 12, 272 p.

Ulrich, E.O., and Smith. W.S.T., 1905, The lead, zinc, and ﬂuorspar
deposits of western Kentucky: US. Geological Survey Professional
Paper 36, 218 p.

Ulrich, GE, 1966, Geologic map of the Church Hill quadrangle, Chris-
tian County, Kentucky: US. Geological Survey Geologic
Quadrangle Map GQ—556.

Ulrich, G.E., and Klemic, Harry, 1966, Geologic map of the Caledonia
quadrangle, Trigg and Christian Counties, Kentucky: US. Geo-
logical Survey Geologic Quadrangle Map GQ—604.

 

 

 

 

 

MISSISSIPPIAN ROCKS IN KENTUCKY

Vail, P.R., 1959, Stratigraphy and lithofacies of Upper Mississippian
rocks in the Cumberland Plateau: Evanston, Ill., Northwestern
University Ph. D. thesis. 143 p.

Van Tuyl, F.M., 1925, The stratigraphy of the Mississippian forma-
tions of Iowa: Iowa Geological Survey, v. 30, p. 33—374.

Varnes, KL, 1979, Index to localities and sources, in Craig, L.C.,
Connor, C.W., and others, Paleotectonic investigations of the
Mississippian System in the United States, Part II, Interpretive
Summary and Special Features of the Mississippian System: US.
Geological Survey Professional Paper 1010, p. 461—546.

Vincent, J .W., 1975, Lithofacies and biofacies of the Haney Limestone
(Mississippian), Illinois. Indiana, and Kentucky: Kentucky Geo-
logical Survey, ser. 10, Thesis Series 4, 64 p.

Walker, KR, 1962, Lithofacies map of Lower Mississippian clastics
of eastern and east-central United States: American Association
of Petroleum Geologists Bulletin, v. 46, p. 105—111.

Wanless, H.R., 1957, Geology and mineral resources of the Beards-
town, Glasford, Havana, and Vermont quadrangles [Illinois]: 11-
linois State Geological Survey Bulletin 82, 233 p.

Webster, GD, 1969, Chester trough Derry conodonts and stratig-
raphy of northern Clark and southern Lincoln Counties, Nevada:
California University Publications of Geological Science, v. 79,
121 p.

Weir, G.W., 1967, Geologic map of the Berea quadrangle, east-central
Kentucky: US. Geological Survey Geologic Quadrangle Map
GQ—649.

1970, Borden Formation (Mississippian) in southeast-central

Kentucky, in Guidebook for field trips, 18th annual meeting,

Southeastern Section, The Geological Society of America: Lex-

ington, Ky., Kentucky Geological Survey, p. 29—48.

1972, Geologic map of the Halls Gap quadrangle, Lincoln Coun-

ty, Kentucky: US. Geological Survey Geologic Quadrangle Map

GQ-1009.

1974a, Geologic map of the Stanton quadrangle, Powell and
Estill Counties, Kentucky: US. Geological Survey Geologic Quad-
rangle Map GQ-1182.

___1974b, Geologic map of the Slade quadrangle, east-central
Kentucky: US Geological Survey Geologic Quadrangle Map
GQ-1183.

1976, Geologic map of the Means quadrangle, east-central Ken-
tucky: US. Geological Survey Geologic Quadrangle Map GQ-1324.

Weir, G.W., Gualtieri, J .L., and Schlanger, 8.0., 1966, Borden For-
mation (Mississippian) in south- and southeast-central Kentucky:
US. Geological Survey Bulletin 1224-F, 38 p.

Weir, G.W., Lee, K.Y., and Cassity, RE, 1971, Geologic map of the
Bighill quadrangle, east-central Kentucky: US. Geological Survey
Geologic Quadrangle Map GQ-900.

Weir, G.W., and McDowell, RC, 1976, Geologic map of the Preston
quadrangle, Bath and Montgomery Counties, Kentucky: U.S. Geo-
logical Survey Geologic Quadrangle Map GQ-1334.

Weir, G.W., and Richards, P.W., 197 4, Geologic map of the Pomeroy-
ton quadrangle, east-central Kentucky: US. Geological Survey
Geologic Quadrangle Map GQ-1184.

Weir, G.W., and Schlanger, 8.0., 1969, Geologic map of the Woodstock
quadrangle, south-central Kentucky: US. Geological Survey
Geologic Quadrangle Map GQ-776.

Weller, J .M., 1931. The Mississippian fauna of Kentucky, in J illson,
W.R., The paleontology of Kentucky: Kentucky Geological
Survey, ser. 6, v. 36, p. 251—291.

1956, Argument for diastrophic control of late Paleozoic cyclo-
thems: American Association of Petroleum Geologists Bulletin,
v. 40, p. 17-50.

Weller, J .M., and others (Mississippian subcommittee), 1948, Correla-
tion of the Mississippian formations of North America: Geological
Society of America Bulletin, v. 59, p. 91-196.

 

 

 

 

 

REFERENCES CITED

Weller, J .M.. and Sutton. A.H., 1940. Mississippian border of Eastern
Interior Basin: American Association of Petroleum Geologists
Bulletin, v. 24, p. 765—858.

Weller, Stuart, 1909, Correlation of the Middle and Upper Devonian
and the Mississippian faunas of North America: Journal of Geol-
ogy, v. 17, p. 257—285.

1920, The Chester Series in Illinois: Journal of Geology, v. 28,

p. 281-303, 395—416.

1926, Fauna] zones in the standard Mississippian sections: J our-

nal of Geology, v. 34, p. 320—335.

1927, The story of the geologic making of southern Illinois: I1-
linois State Geological Survey Educational Series 1. 41 p.
Weller, Stuart, and Sutton, A.H., 1951, Geologic map of the western
Kentucky fluorspar district: US. Geological Survey Mineral In-

vestigations Field Studies Map MF-2, scale 1:62.500.

Wells, Dana, 1950. Lower middle Mississippian of southeastern West
Virginia: American Association of Petroleum Geologists Bulletin,
v. 34, p. 882—922.

Whaley, P.W., Hester, N.C., Williamson, A.D., Beard, J.G.. Pryor,
W.A., and Potter, RE, 1979, Depositional environments of Penn-
sylvanian rocks in western Kentucky—Annual field conference
of the Geological Society of Kentucky: Lexington, Ky., Kentucky
Geological Survey, 48 p.

Whitehead, N.H., III. 1976. The stratigaphy, sedimentology, and con-
odont paleontology of the Floyds Knob Bed and Edwardsville
Member of the Muldraugh Formation (Valmeyeran), southern In-
diana and north-central Kentucky: Urbana, 111., University of 11-
linois M.S. thesis, 443 p.

_1978a, Lithostratigaphy, depositional environments, and cone-
dont biostratigraphy of the Muldraugh Formation (Mississippian)
in southern Indiana and north-central Kentucky: Southeastern
Geology, v. 19, p. 83—109.

__1978b, Lithostratigraphy and depositional history of the Maury,
Grainger, Borden, Price, and Maccrady Formations (Lower Missis-
sippian) of the east-central United States (Illinois, Indiana, Ken-
tucky, Tennessee, and Virginia): Geological Society of America
Abstracts with Programs, v. 10, p. 201—202.

 

 

 

rims. GOVERNMENT PRINTING OFFICE: 1990-573-049/26001

 

125

Williams, J .S., 1957, Paleoecology of the Mississippian of the Upper
Mississipi Valley region, in Ladd, H.S., ed.. Paleoecology: Geo-
logical Society of America Memoir 67, p. 279—324.

Williamson, A.D.. and McGrain, Preston, 1979, Hawesville, Kentucky,
to Rough River State Resort Park, in Palmer, J .E., and Dutcher,
R.R., eds., Depositional and structural history of the Pennsylva-
nian System of the Illinois basin. Part I. Road log and descrip-
tions of stops; Field Trip 9, 9th International Congress of
Carboniferous Stratigraphy and Geology: Illinois State Geological
Survey, Guidebook Series 15, p. 28-37.

Willman, H.B., and others, 1975, Handbook of Illinois stratigraphy:
Illinois State Geological Survey Bulletin 95, 261 p.

Wilpolt. R.H., and Marden, D.W., 1959, Geology and oil and gas
possibilities of Upper Mississippian rocks of southwestern
Virginia, southern West Virginia, and eastern Kentucky: US.
Geological Survey Bulletin 1072-K, p. 587-656.

Wilshire, H.G., 1964, Geology of the New Concord quadrangle and
part of the Buchanan quadrangle, Kentucky: US. Geological
Survey Geologic Quadrangle Map GQ-313.

Wilson, W.J., 1950, Cincinnati sub-aqueous ﬂow-marking in the
Lower Mississippian strata of south-central Ohio and adjacent
parts of Kentucky: Cincinnati, Ohio, University of Cincinnati
M.S. thesis.

Withington. CE, and Sable, E.G., 1969, Geologic map of the Rock
Haven quadrangle. Kentucky-Indiana. and part of the Laconia
quadrangle, Kentucky: US. Geological Survey Geologic Quad-
rangle Map GQ—780.

Wolfe, E.W., 1964, Geology of the Fairdealing quadrangle, Kentucky:
US. Geological Survey Geologic Quadrangle Map GQ-320.
Woodward, HP, 1961, Preliminary subsurface study of southeastern
Appalachian Interior Plateau: American Association of Petroleum

Geologists Bulletin, v. 45, p. 1634-1655.

Worthen, A.H., 1860, Remarks on the discovery of a terrestrial ﬂora
in the Mountain Limestone of Illinois [abs.]: American Associa-
tion for the Advancement of Science Proceedings, v. 13,
p. 312-313.

 

 

 

 

Earthquake-Induced Liquefaction
Features in the Coastal Setting of South
Carolina and in the Fluvial Setting of the
New Madrid Seismic Zone

By S.F. OBERMEIER, R.B. JAGOBSON, ].P. SMOOT, R.E. WEEMS, G.S. GOHN,
].E. MONROE, and BS. POWARS

 

U.S. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1504

 

 

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON: 1990

DEPARTMENT OF THE INTERIOR

MANUEL LUJAN, Jr., Secretary

U.S. GEOLOGICAL SURVEY

Dallas L. Peck, Director

Any use of trade, product, or ﬁrm names in this ublication is for
descriptive purposes onlyT and does not imply en orsement by the
.S. Government

 

Library of Congress Cataloging in Publication Data

Earthquake-induced liquefaction features in the coastal setting of South Carolina and in the ﬂuvial setting of the New Madrid
seismic zone / by Stephen F. Obermeier [et 3.1.].

p. cm.——(U.S. Geological Survey professional paper ; 1504)

Bibliography: p.

Supt. of Docs. no. : I 19.16 : 1504

1. Geology—South Carolina—Atlantic Coast. 2. Geology—Missouri-New Madrid Region. 3. Soil liquefaction. 4. Geology.
Structural. I. Obermeier, Stephen F. II. Series: Geological Survey professional paper ; 1504.

QE162.A85E37 1990

557.57—dc20 89—600177

CIP

 

For sale by the Books and Open-File Reports Section, US. Geological Survey,
Federal Center, Box 25425, Denver, CO 80225

CONTENTS

 

Page Page
Abstract ............................................................................ 1 New Madrid Seismic Zone—Continued
Introduction.. 1 Characteristics of and Criteria for Earthquake-Induced
Coastal South Carolina ........................................................ 2 Liquefaction Features—Continued
Regional Setting—Geology and Liquefaction Susceptibilityn 3 Fissures ............................................................ 26
Characteristics 0f and Criteria for Sand-Blow Formation ----- 4 Intruded Features .............................................. 27
Filled Sand-Blow Craters ......................................... 5 Features of Unknown or Nonearthquake Origin ............... 34
Vented-Sand Volcanoes ........................................... 8 Sand Boils ..................... 34
Featgres1 1Showmg Ev1dence of Lateral Spreads or Ground 11 Mima Mounds ................. 34
sc1 9. ions ...........................................................
Regional Distribution of Sand Blows ........ 12 332d Structures .................................................... 35
. . . er Features ..................................................... 38
Other FOSSIL-’18 01191.15 """ """"""""" 13 Local Geologic Controls on Production of Vented Sand ...... 38
Features of Weathering Origin .................................... 14 T t t Thi kn 38
Overview of the South Carolina Liquefaction Study ......... 18 ops ra um , c ess """
Earthquake Ages ................................................... 19 Topstratum ”“9108? """" 38
Shaking Severity Estimation ____________ 19 Substratum Grain Size ............ : .......... ~ ..................... 3 9
Holocene Earthquake Shaking __________ 19 Resistance of Source Sand to Liquefactlon ................. 39
New Madrid Seismic Zone ............................................. 20 Overview Of Studies ........................................................... 39
Regional Geologic and Seismotectonic Setting ___________________ 22 Geologic Criteria ......................................................... 39
Characteristics of Quaternary Alluvium ..................... 23 Types of Earthquake Features ..................................... 40
Seismotectonic Setting ............................................ 24 Geologic Controls ......................................................... 4O
Characteristics of and Criteria for Earthquake-Induced Suggestions for Future Research ................................. 41
Liquefaction Features ......................................... 24 Relevance of Liquefaction Features ...................................... 42
Vented-Sand Volcanoes .......................................... 24 References Cited ............................................................... 42
ILLUSTRATIONS
[Plate in pocket]
Page
PLATE 1. Typical load structures produced by rapid sedimentation, observed in trench near Marked Tree, Arkansas .................. In pocket
FIGURE 1. Map showing 1886 and older sand-blow sites ................................................................................................................ 2
2. Schematic cross section of representative barrier showing sediment types, ground-water table locations, ﬁlled sand-blow
craters, and Bh soil horizons ................................................................................................................................ 4
3. Sketch and photograph of craters produced by the 1886 earthquake ................................................................................. 6
4. Schematic cross section of normal type of ﬁlled sand-blow crater. .................................................................................... 7
5. Ternary diagram showing grain sizes of paired samples at identical depths inside and outside liquefaction craters ................... 7
6. Schematic cross sections showing sand-ﬁlled ﬁssures interpreted as resulting from liquefaction and ﬂowage during the
1886 earthquake ................................................................................................................................................ 9
7. Schematic vertical section of representative vented-sand volcano ..................................................................................... 10
8. Photograph of vented-sand volcanoes that have coalesced to form a continuous sheet of sand on the ground surface ................. 11
9. Schematic cross section of ﬁlled sand-blow crater, illustrating aspects associable with downslope movement ........................... 12
10. Schematic cross section showing pedogenic tonguing of BE- and E-horizon sand into underlying B horizon and a graph
showing particle-size data for a fractured pedogenic tongue ...................................................................................... 15
11. Sketch of cross section through a white, pedogenic sand tongue at site HW, near Hollywood, 8.0. ....................................... 16
12. Sketch of vertical section through BE’ horizon at site HW, near Hollywood, S.C. .............................................................. 18
13. Map showing late Quaternary alluvial deposits of St. Francis and Western Lowlands Basins ................................................ 21
14. Map showing area covered by vented sand, estimated epicenters of strongest 1811—12 earthquakes, and faults and fault zones.. 22
15. Schematic east-west cross section showing geologic and ground-water setting of St. Francis Basin ....................................... 23
16. Block diagram showing the conﬁguration of the buried New Madrid Rift Complex ............................................................. 24

   

 
 
 

  
 

 
 
 

 

 

III

IV

17.
18.
19.
20.

21.

23.

25.

Table 1.

CONTENTS

   
 

Map of northern Mississippi Embayment showing earthquake epicenters, plutons, rift boundaries, and faults ......................... 25
Photograph showing stratigraphy of vented-sand volcano with organic-rich silt between two ﬁning-upward sequences of sand 26
Photographs and line drawings of an “eruptive vent” that cut stratiﬁed deposits of a vented-sand volcano .............................. 27
Aerial photographs showing vented sand, interpreted as the product of liquefaction and ﬂowage during the 1811—12

earthquakes ...................................................................................................................................................... 28
Photographs showing sand dikes and sills, interpreted as having originated by liquefaction and ﬂowage during the

1811—12 earthquakes ........................................................................................................................................... 31
Sketch and photographs of section showing earthquake-induced intrusions in Holocene sediments and underlying

Wisconsinan braided-stream sands observed in ditch about 15 km northwest of Marked Tree, Ark. ................................. 32
Aerial photograph of mima mounds in the northern part of St. Francis Basin ..................................................................... 36
Sketches showing pseudonodules formed in a shaking experiment by Keunen ................................................................ 37
Schematic drawing of the development of load-casted ripples, caused by ripple crests sinking into soft mud ....................... 37

TABLE
Page

Estimated relative susceptibility of saturated cohesionless sands to liquefaction during strong seismic shaking ........................... 4

EARTHQUAKE-INDUCED LIQUEF ACTION FEATURES IN THE
COASTAL SETTING OF SOUTH CAROLINA AND IN THE FLUVIAL
SETTING OF THE NEW MADRID SEISMIC ZONE

By S.F. OBERMEIER,1 R.B.JACOBSON,1 J.P. SMOOT,l R.E. WEEMS,1 G.S. GOHN,1 ].E. MONROE,2 and
BS. POWARS1

ABSTRACT

Many types of liquefaction—related features (sand blows, ﬁssures,
lateral spreads, dikes, and sills) have been induced by earthquakes in
coastal South Carolina and in the New Madrid seismic zone in the
Central United States. In addition, abundant features of unknown and
nonseismic origin are present. Geologic criteria for interpreting an
earthquake origin in these areas are illustrated in practical applica-
tions; these criteria can be used to determine the origin of liquefaction
features in many other geographic and geologic settings.

In both coastal South Carolina and the New Madrid seismic zone, the
earthquake-induced liquefaction features generally originated in clean
sand deposits that contain no or few intercalated silt- or clay—rich
strata. The local geologic setting is a major influence on both develop-
ment and surface expression of sand blows. Major factors controlling
sand-blow formation include the thickness and physical properties of
the deposits above the source sands, and these relationships are
illustrated by comparing sand blows found in coastal South Carolina (in
marine deposits) with sand blows found in the New Madrid seismic zone
(in ﬂuvial deposits). In coastal South Carolina, the surface stratum is
typically a thin (about 1 m) soil that is weakly cemented with humate,
and the sand blows are expressed as craters surrounded by a thin sheet
of sand; in the New Madrid seismic zone the surface stratum generally
is a clay-rich deposit ranging in thickness from 2 to 10 m, in which case
sand blows characteristically are expressed as sand mounded above the
original ground surface.

Recognition of the various features described in this paper, and
identiﬁcation of the most probable origin for each, provides a set of
important tools for understanding paleoseismicity in areas such as the
Central and Eastern United States where faults are not exposed for
study and strong seismic activity is infrequent.

INTRODUCTION

The near—surface and surface expressions of
earthquake-induced liquefaction are highly dependent on
the local geologic setting. This paper discusses the types
of features, primarily sand blows, that formed on level or
nearly level ground in two very different geologic set-

Manuscript approved for publication April 20, 1989.
1U.S. Geological Survey, Reston, Va.
2U.S. Army Corps of Engineers, Memphis, Tenn.

 

tings, one in which the parent sediments have a marine
or near-marine origin (coastal South Carolina) and one in
which the parent sediments have a ﬂuvial origin (New
Madrid seismic zone). The diverse types of sand blows
and the equally diverse controls on their formation
became apparent during ﬁeld studies recently conducted
to locate regions subjected to strong earthquake shaking,
as indicated by the presence of liquefaction-induced
features. In coastal South Carolina, ﬁeld studies focused
on sand blows of very late Pleistocene to Holocene age,
as well as features caused by the 1886 Charleston earth-
quake. In the New Madrid seismic zone, the search was
largely restricted to sand blows caused by the 1811—12
earthquakes. In both areas, features of nonearthquake
origin occur that might be confused with earthquake-
induced features. Geologic criteria have been developed
by which the earthquake-induced liquefaction features
can usually be distinguished. (“Liquefaction” in this
paper is deﬁned as transformation of water-saturated
sediments from the solid to liquid state as a consequence
of increased pore pressure; this transformation indicates
that grains are suspended by pore water. “Fluidization”
is deﬁned as the process of both suspension and transport
of grains by pore water.)

The scope of the paper includes both intruded features
(for example, dikes and sills) and vented features (for
example, sand blows). The term “sand blows” is used to
indicate features formed where earthquake shaking
causes liquefaction at depth followed by the venting of
the liqueﬁed sand and water to the surface. Features
described as sand boils in this paper form in the absence
of earthquake shaking and involve transport of sediment
to the surface by artesian ﬂow (springs). This distinction
in terms is made because the processes within the ground
that lead to development of earthquake-induced features
differ from processes not related to earthquakes. Sand
blows are thought to often form in response to liquefac—
tion that is in turn followed by segregation of sand and

2 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

water at depth, upward development of a vent, and
ﬁnally the violent discharge of soil and water that soon
diminishes to an ebbing flow (Scott and Zuckerman,
1973). (A slightly different model for some sand deposits
is proposed in this paper.) In contrast, sand boils form by
the downward development of a vent that typically does
not enlarge violently and that continues to flow as long as
high artesian pressure remains, commonly for days or
years. This difference in development processes between
sand boils and sand blows can often be distinguished in
the ﬁeld by use of techniques discussed in this paper.

COASTAL SOUTH CAROLINA

The strongest historic earthquake in the Southeastern
United States took place in 1886 near Charleston, SC.
Throughout much of the epicentral region, an area about
35 km wide and 50 km long, the Modiﬁed Mercalli
intensity ranged from IX to X (Bollinger, 1977). The
estimated body-wave magnitude (mb) was between 6.6
and 7.1 (Nuttli, 1983a). The potential for a future earth-
quake having the strength of the 1886 earthquake is a
major concern in engineering design in the Southeast.
The concern is reinforced by a 300—year historical record
of continuing weak seismic activity near Charleston. The
source of the earthquakes in the Charleston area remains
unknown, and seismotectonic hypotheses are widely
disparate, despite many geologic, geophysical, and seis-
mic studies during the past decade. No faults or fault
systems have been identiﬁed that adequately explain the
large 1886 Charleston earthquake or the other smaller,
historic earthquakes that have occurred throughout
much of South Carolina (Hays and Gori, 1983; Dewey,
1985; Science News, 1986). Because direct evidence of
seismotectonic conditions is lacking and because the
historic earthquake record is too limited to provide a
dependable basis for estimating the frequency of moder—
ate to strong earthquakes, we undertook a search for
pre-1886 sand blows.

Results of the search are shown on ﬁgure 1. Figure 1
shows the approximate boundary of the 1886 Charleston
earthquake meizoseismal zone; the sites conspicuous in
1886 for development of many sand blows, described as
“craterlets” by Dutton (1889); and the sites of pre-1886
sand blows that we discovered. The unshaded portion of
ﬁgure 1 encompasses the area that was searched for sand
blows. Radiocarbon ages of sand-blow materials at site
HW show that at least three pre-1886 Holocene earth-
quakes have produced sand blows in that area (Weems
and others, 1986). It is not yet known how many earth-
quakes are represented by the sand blows at the widely
scattered sites on ﬁgure 1, nor are the seismic source
zones known for these prehistoric sand blows. What is
known is that the prehistoric sand blows extend far
beyond the limits of 1886 sand blows (see ﬁg. 1) in

 

34°

 

79°

  
  
 
 
  
 
 
 
  
  

— — — Approximate boundary
of 1886 meizoseismal
zone (Bollinger, 1977)

= Areas conspicuous in 1886
for craters (Dutton, 1889)

' Pre-1886 sand-blow site

designation (this study)

X Outer limit of historically
documented 1886 sand
blows (Georgetown)

 

 

 

 

LOCATION OF STUDY AREA

0 30 MILES

0 30 KILOMETEHS

 

 

FIGURE 1. —1886 and older sand-blow sites. Unshaded onshore region
is predominantly of marine sediments younger than about 240,000
years. Shading pattern denotes region of older marine sediments
that was not reconnoitered. Younger ﬂuvial sediments occur locally.
All sand-blow sites in the region with no shading are in marine-
related deposits, and all sites in region with shading are in ﬂuvial
deposits. Numerous sites discovered in and near the 1886 meizoseis-
mal zone are not shown because of lack of space. Abbreviations
adjacent to sand-blow sites are speciﬁc site designations.

sediments having approximately the same liquefaction
susceptibility (Obermeier and others, 1986; unpublished
data) and that the strongest Holocene shaking has prob-
ably been near Charleston (Obermeier and others, 1989).

All the pre-1886 sand blows that we found have no
expression on the ground surface that is discernible by
onsite examination or on aerial photographs. The sand—
blows are seen only where exposed in walls of excava-
tions at least 1.5 m deep, typically in drainage ditches

COASTAL SOUTH CAROLINA 3

and borrow pits. At most sites shown on ﬁgure 1, at least
three or four sand blows are exposed within a few
hundred meters of one another. The following section
focuses on the geologic setting in which these sand blows
originated.

REGIONAL SETTING—GEOLOGY AND LIQUEFACTION
SUSCEPTIBILITY

In South Carolina, the coastal region is known locally
as the “low country” because it has low local relief (1—3
m) and low elevation (0—30 m) and because vast expanses
of swamp and marshland are under water much of the
year. Most of the Carolina low country is covered by a 5-
to 10—m—thick blanket of unconsolidated Quaternary
marine and ﬂuvial deposits, which lies on semilithiﬁed
Tertiary sediments (McCartan and others, 1984). The
Quaternary sediments primarily occur as a series of six
well-deﬁned, temporally discrete, interglacial beaches
and associated back barrier and shelf deposits that form
belts subparallel to the present shoreline. The oldest
beach deposits are farthest inland and are at the highest
altitudes; younger beach deposits are progressively
closer to the ocean and are at successively lower alti-
tudes. Most beach deposits are 8 to 15 m thick.

Cutting across these marine and marginal-marine
deposits at nearly right angles are four major rivers, the
Savannah, Edisto, Santee, and Pee Dee (ﬁg. 1). Border-
ing these rivers are ﬂuvial terrace sediments of rather
limited extent that consist almost exclusively of clean
sand (that is, sand without clay, silt, or gravel).

Figure 1 shows the approximate areal extent of the
marine-related deposits (beach, shelf, and open-sound
back barrier) designated as Q1, Q2, and Q3 by McCartan
and others (1984). The part of ﬁgure 1 containing units
Q1, Q2, and Q3 is shown Without shading. Q3 deposits
are about 200,000 to 240,000 years old (Szabo, 1985) and
are present as far as 20 to 40 km inland from the modern
coast. The intervening Q2 deposits are about 80,000 to
130,000 years old (Szabo, 1985). Unit Q1 is closest to the
ocean and is made of younger deposits. The search for
sand blows was generally restricted to units Q2 and Q3.
Older units have such a low susceptibility to liquefaction
(due to effects of chemical weathering) that the likelihood
of forming sand blows has been extremely low during the
late Pleistocene and Holocene. Because unit Q1 generally
has such a high ground-water table, the possibility of

‘ ﬁnding exposed sand blows was quite limited.

Formation of sand blows in any geologic setting
depends primarily on the depth to the water table, the
properties and thickness of materials in the depth range
susceptible to liquefaction during shaking, and the thick-
ness and characteristics of sediments above the zone that
was liqueﬁed during shaking. Speciﬁc relations between
liquefaction susceptibility and subsequent formation of

 

sand blows in the South Carolina low country are as
follows:

(1) A water table very close to the ground surface
usually greatly increases the susceptibility to liquefac-
tion, even in comparison with a water table at depths of
only 3 to 4 m. The modern water-table depth is generally
0 to 1.5 m throughout the low country and is typically a
subdued mimic of the surface topography. The water
table is deepest under hilltops and may come to the
surface in swales. As surface elevations decrease toward
the ocean, the water table is generally nearer the sur-
face; within 15 km of the ocean, the water table is rarely
deeper than 1.5 to 2 m.

(2) Clean sands are generally the only materials
observed to have liqueﬁed in the Charleston region. At
one site (site SAN, ﬁg. 1) the source sand bed contains as
much as 3 to 5 percent clay; typically there is less than 1
percent silt and clay at all other sites. The liqueﬁed sands
are generally ﬁne-grained, well-sorted (that is, uni-
formly graded) beach sands.

Principal properties of sand that control liquefaction
susceptibility during shaking are degree of compaction or
state of compactness (known as “relative density” by
geotechnical engineers), sand-grain size and sorting, and
cementation of the sand at grain-to-grain contacts.

The state of compactness is commonly a reﬂection of
the energy (or mode) of deposition. To illustrate, beach
deposits laid down in a high-energy environment such as
a pounding surf zone are generally more dense than sand
deposited in a quiet zone away from the inﬂuence of the
surf. Higher density makes the surf-zone deposits more
difﬁcult to liquefy. Locally within the surf zone, though,
there are many regions of low-energy deposition where
the sand is not so compacted by wave pounding. In
addition, for a given relative density, the ﬁne-grained,
well-sorted sands of ancient and modern beaches
throughout the low country are much more susceptible to
liquefaction than standard sands used for engineering
analysis (Cullen, 1985), and this increased susceptibility
due to grain-size effects and lower local compactness
must account in part for the widespread liquefaction
during the 1886 earthquake.

Cementation in near-surface subaerial environments is
chieﬂy a reﬂection of the age of the deposit. Even slight
cementation by agents such as silica, calcite, or clay
dramatically reduces liquefaction susceptibility; thus liq-
uefaction potential is related to age of deposits in many
geologic settings (Youd and Perkins, 1978), and this
relation is particularly apparent in the low country.
Table 1 lists deposits in the low country shown in ﬁgure
1 in terms of depositional environment (type of deposit)
and age and provides an estimate of their relative
susceptibilities to liquefaction during strong seismic
shaking.

4 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

 

 

   
 
 
  
   
    
 

 
 
 
 
 

 

Near-shore deposit . . -
of shelly, silty, . . ~
and clayey sand

 

 

  
 
 
 
  
     
    
 

SE NW

I4 L]

I‘ 200—500 m r]

/Water table Filled sand-blow craters

Water table 2‘5 m
/atsurface V _—_'i
,_T _ _IT _ . _ . agoonal deposi
_- agoona .-_ _ - ~ '- g : -: . _, : ..,.—..—:of clay:~
:deposit_ —_ _ uBnaiiolrenri-lgarigeegcs’snd Black, Bh soil horizon' ' —

 

Semilithified
Tertiary marl

 

 

FIGURE 2. ——Schematic cross section of representative barrier showing sediment types, ground-water table locations, ﬁlled sand-blow craters, and
Bh (humate—rich) soil horizons. Modern shoreline is located southeastward. Lagoonal clay deposit at left is younger and lower in elevation than

the barrier-bar deposit.

Features large enough to be interpreted as possibly
having an earthquake origin in the low country were
found only in sand deposits having a total thickness
exceeding 2 to 3 m; within this 2— to 3-m—thick deposit,
the thinnest individual source stratum was 0.3 m.

(3) The local geologic setting has a major role in the
formation of earthquake-induced sand blows. The geo-
logic setting most frequently associated with sand blows
is the crest or ﬂank of Pleistocene beach ridges, where a
thin surﬁcial cover of a clay-bearing sand or humate-rich
sand overlies clean sand. According to ﬁrst-hand obser-
vations of effects of the 1886 earthquake by Earl Sloan,
“these craterlets are found in greatest abundance in belts
parallel with (beach) ridges and along their anticlines”
(Peters and Herrmann, 1986, p. 68). A schematic cross
section through a typical low-country beach ridge, such
as the ridges described by Sloan, is presented in ﬁgure 2.
To a much lesser extent, sand blows have been found in
back-barrier environments.

A thin clay—bearing stratum (or other stratum having
very low permeability) above the liqueﬁed zone is gen-
erally an important control on development of sand blows
(Scott and Zuckerman, 1973; Obermeier, 1988; Ishihara,
1985). A veneer of nonliqueﬁable sediment, 1 to 2 m
thick, aids greatly in the formation of and recognition of
sand blows in the low country. On the other hand, a
clay-bearing or other low-permeability stratum thicker
than 4 to 5 m prevents sand—blow formation at the great
majority of sites in the low country; a thickness greater
than 2 to 3 m seriously impedes such formation.

CHARACTERISTICS OF AND CRITERIA FOR SAND-BLOW
FORMATION

Two kinds of pre—1886 sand blows have been recog-
nized: (1) ﬁlled sand-blow craters (craterlets) and associ-

 

ated sedimentary structures and (2) sand volcanoes that
have vented to the surface, leaving relict sand mounds.
The crater-type sand blows generally occur only where a
surﬁcial soil having less than several percent clay has
formed on a parent material of clean sand. Where the
surﬁcial soil is richer in clay, vented-sand volcanoes are
much more likely to form.

Exact locations were not known for 1886 craterlet sand
blows when our study was initiated in 1983. Only one
1886 vented type of sand volcano had been discovered
and examined (Cox, 1984). Thus it has been necessary to
locate sedimentary features that display structures con-
sistent with an earthquake origin and to develop criteria
for interpreting whether or not these features have an
earthquake origin.

TABLE 1.— Estimated relative susceptibility of saturated cohesionless
sands to liquefaction during strong seismic shaking

[Sourcez Youd and Perkins (1978). Strong seismic shaking is deter-
mined by two parameters, peak horizontal acceleration and duration
of largest acceleration. For a mb=5 earthquake, strong seismic
shaking is deﬁned as an acceleration of about 0.2 g for at least several
seconds; for a stronger earthquake, the threshold acceleration is
about 0.15 g for a duration of 10 seconds; for a much stronger and
longer duration earthquake, the threshold acceleration can be less
than 0.1 g (T.L. Youd, Brigham Young University, oral commun.,
1985)]

 

Liquefaction susceptibility for deposits of various ages

 

 

Types of deposit <500 ﬁndizl; yr Pleistocene Pre-Pleistocene
Dunes .......... High Moderate Low Very low
Beach (low wave

energy ........ High Moderate Low Very low
Foreshore ....... High Moderate Low Very low
Beach (high wave

energy) ........ Moderate Low Very low Very low
River channel. . . . Very high High Low Very low
Flood plain ...... High Moderate Low Very low

 

COASTAL SOUTH CAROLINA 5

Geologic criteria that we have developed for interpret-
ing whether near-surface features are earthquake—
induced sand blows generally have four elements:

1. The features have sedimentary characteristics that
are consistent with an earthquake—induced liquefaction
origin; that is, there is evidence of an upward-directed,
strong hydraulic force that was suddenly applied and was
of short duration.

2. Characteristics such as shape, width, and depth are
consistent with historical observations of liquefaction
during the 1886 earthquake.

3. The features are in ground-water settings where a
suddenly applied, strong hydraulic force of short dura-
tion could not be reasonably expected except from
earthquake-induced liquefaction. In particular, these
settings are extremely unlikely sites for artesian springs.

4. Similar features occur at multiple locations, prefer-
ably within a few kilometers of one another, having
similar geologic and ground-water settings. Where evi-
dence of age is present, it should support the interpre-
tation that the features formed in one or more discrete,
short episodes that individually affected a large area and
the episodes were separated by long time periods during
which no such features formed.

As fewer of these criteria are satisﬁed, the conﬁdence
in an earthquake origin generally diminishes. Subse-
quent sections describe the application of the four crite-
ria.

Whether earthquakes or other mechanisms induce
small liquefaction-ﬂowage features is often impossible to
determine. Small synsedimentary liquefaction-ﬂowage
features such as dikes and sills as much as 0.25 to 0.5 m
long and a few centimeters thick are not unusual in
point—bar deposits and in beach surf zones. Very small
features similar to sand blows are also common in beach
surf zones. These small features result from a variety of
forces, including dynamic wave loadings and static
slumping. An earthquake origin was considered a possi-
ble mechanism in this study only if the ﬂowage features
were much larger than these small features and if, in
addition, the ﬂowage features cut the surface soil proﬁle.

FILLED SAND-BLOW CRATERS

All ﬁlled sand-blow craters belong to a single morpho-
logic group having many common features in the ﬁll
sediments, although systematically occurring variants
also occur. The normal type of ﬁlled sand-blow crater is
discussed ﬁrst. In a later section there is discussion of
two variants that may indicate association with
earthquake-induced landslides.

Almost all pre-1886 sand-blow sites shown on ﬁgure 1
have sand blows whose original morphology and size are
comparable to the 1886 craters described by Dutton

 

(1889), except that the craters are now ﬁlled with
sediment. Figure 3 shows craters that represent
moderate- to large-sized craters produced by the 1886
earthquake. A crater is a hole at the ground surface that
forms as liqueﬁed sand vents to the surface. In the
process, the forceful upward surging of sand and water
also scours and enlarges a hole and deposits sediment
beyond the crater. When ﬂowage stops, a surﬁcial sheet
of ejected sand and soil surrounds the rim of the crater,
and clasts of dark soil are commonly scattered along the
base of the steep wall (ﬁg. 3). Examination of photo-
graphs shows that, in 1886, the surﬁcial sheets commonly
appear to have had thicknesses near crater rims of about
15 to 20 cm and maximum diameters rarely exceeding
about 3 to 4 m. The maximum reported thickness of
vented sand was 1 m, and the maximum crater diameter
was about 6 m (Dutton, 1889).

About 75 percent of the sites on ﬁgure 1 are on barrier
and nearshore marine sands. The sands are almost
exclusively well-sorted, ﬁne— to medium-grained quartz,
with less than 5 percent heavy minerals and mica
(McCartan and others, 1984; Gelinas, 1986). Surface
weathering has imposed a soil proﬁle on these marine—
related sands. The soils are classiﬁed as spodosols,
alﬁsols, and ultisols, but most sand blows occur in poorly
drained areas where humic spodosols (humods) have
formed. Humods are characterized by a thin (<10- to
15-cm) surﬁcial A horizon (organic matter and several
percent sand) overlying a thin (<10- to 15-cm), very light
gray E horizon; the E horizon overlies a thick (0.5- to
1.5-m) black Bh horizon (humate-enriched sand contain-
ing a few percent clay) and a variably thick (0. 1- to 1-m),
gray to light-yellow B—C horizon (transition zone
between B and C horizons). The B—C grades down into
C-horizon sands (parent material), which are very light
gray in the upper 1 to 2 m and grade down to a greenish
hue. Beneath the upper 0.5 to 1 m of the C horizon, the
original bedding consists of thin, horizontal, black heavy—
mineral laminae about 0.25 to 0.5 cm apart. The sand-
blow features cut the solum and the C horizon. The
source sand beds that liqueﬁed during earthquake shak—
ing generally occur within the depth range of 3 to 10 m.

Figure 4 is a vertical section of a ﬁlled sand-blow
crater that is representative of the type observed at
almost all of our study sites (ﬁg. 1). The ﬁgure illustrates
characteristics that we consider to be compatible with an
earthquake-induced liquefaction origin. In ﬁgure 4, the
soil horizon is cut by an irregular crater, which is ﬁlled
with stratiﬁed to structureless (that is, nonstratiﬁed)
and graded sediments. The ﬁll materials are ﬁne- to
medium-grained sand and clasts from the Bh, B—C, and C
soil horizons and sand from depths much below the
exposed C horizon. Walls of the crater are generally

6 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

 

FIGURE 3. —Craters produced by the 1886 earthquake. A, Sketch from
a photograph of an 1886 crater (sand blow at Ten Mile Hill, near the
present Charleston airport). Note that the crater contains sand
sloughing toward the lowest parts and that there is a constructional
sand volcano located in the right part of the crater (at arrow). The

smooth and sharp when viewed up close, especially in the
lower part of the crater. Walls of some craters are very
jagged, however.

The Bh horizon generally is much thinner over the
central part of the ﬁlled crater than on the sides, and the
Bh horizon of the laterally adjoining undisturbed soil is
abruptly thicker. Clay content in the Bh horizon in the
crater is much less than in the undisturbed soil. Figure 5
is a sand-silt-clay ternary diagram of particle sizes of
pairs of samples taken at identical depths inside and
outside the liquefaction craters. The lines connecting
pairs of samples show that crater samples have consis-
tently less clay and silt than the adjacent undisturbed
soil. The Bh horizon on the ﬁlled crater typically is
thicker, is more clay rich, and has better developed soil
structure (that is, peds) with increasing age; craters
older than about 5,000 to 10,000 years have Bh horizons
that approach the thickness and development of those in
laterally adjacent undisturbed soils.

A sequence of ﬁve layers, each having speciﬁc charac-
teristics, occurs beneath the Bh horizon in the ﬁlled
crater. The sedimentary characteristics of these layers
are less distinct in older craters because of pedogenesis
in the crater deposits. Layer 5 (ﬁg. 4) is a structureless
(that is, massive), gray, humate-stained sand, which
overlies a thinly (2- to 3-mm) laminated sequence of
alternating light- and dark-colored sands (layer 4). The
lamina typically are discontinuous and irregular in thick-
ness, as illustrated in ﬁgure 4. The dark color is generally

 

crater is surrounded by a thin blanket of sand partly veneered with
cracked mud. B, Photograph of typical crater produced by the 1886
earthquake. Note the thin blanket of ejected sand around the crater
and sand and clasts of dark soil within the crater. (Photograph from the
archives of the Charleston museum.)

imparted by humate staining; the dark lamina may
contain appreciable amounts of silt and clay. The basal
bed of this sequence (layer 4) is clay-rich in perhaps 10
percent of the ﬁlled craters and is rarely thicker than 1
cm. The basal bed sharply overlies a medium-gray struc-
tureless sand (layer 3). Layer 3 contains many small
clasts (1- to 5—mm diameter) of Bh material, charcoal, and
wood. This clast-rich layer grades down into a structure-
less sand zone (layer 2) containing many intermediate-
sized clasts (5— to 20-mm diameter) of friable Bh material
and occasional extremely friable clasts of light-colored
sand. The clasts of Bh material and sand have, respec-
tively, the same color, mineralogy, consistency (that is,
resistance to being crushed between ﬁngers), and fria-
bility as the adjoining sands of the B—C and C horizons.
Layer 2 grades down into layer 1, containing densely
packed intermediate-sized (1—5 cm) and large-sized (>5
cm) clasts of Bh material in a structureless sand matrix;
the large clasts have diameters exceeding 25 cm in many
ﬁlled craters. Many of the clasts in layer 1 have their long
axes vertically oriented. At and below the thinly strati-
ﬁed sequence (layer 4), the sides of the bowl are sharply
deﬁned by a color boundary and by the presence of clasts
within the bowl.

Beneath the bowl are vents containing structureless
clean sand. Sides of the vents sharply cut bedding in the
C horizon.

The matrix sand in layers 1 through 3 contains an
extremely small percent of clay-sized material and clay

COASTAL SOUTH CAROLINA

 

    

Bh

 
     
 
 

/ Layer 4:

  
  

 
  

   

/ .

   
  
  

  
 

_ . Layers: structureless san'd ‘ Bedded sequence
Undlsturbed -.~' .' ,- -- .- “ : _- -l\/ ofalternatmg clean

soil ' ‘ ,'~_'.'Vand organic-matter-
/ -_ 1, rich sand'

Layer 3: Sand and I
many small clasts

         

 

 

  
  

   

'- .' .~"'-’\.I¢‘-'.1'\..;.’r .. . .4.“ 5
:T:: L‘ 7‘% B—C . .3 '. .' -' Layer 2: Sand and scattered-'2 T472“ -
:'_‘_—_—._'_ x ‘ E g clasts of clean C-horizon sand and . . D. B—C
—:—:~::—:—::—:{ rrr QAchh were, - 949mm
:—_'_—_—_*T—_—3 W8” ' - ' ‘ ‘ (7. :3 H; w
_ , _ C _____ _\ ~ ' ~

 

 

 

' ' ‘ rLayer 1: Sand and large clast57:::'_ __ .ﬁ
—:of Bh_horizon material _:__—,_—:__—_—__‘

 

 

 

ND VERTICAL EXAGGERATIDN

FIGURE 4. —-Schematic cross section of normal type of ﬁlled sand-blow crater. Letters correspond to soil horizon designations. The ﬁlled crater
in this ﬁgure much predates the 1886 earthquake, based on thickness of Bh horizon in the ﬁlled crater.

10%5‘“ minerals. Both the percent of clay minerals and percent

of clay- and silt-sized material in layers 1 through 3 are
much less than in the laterally adjoining undisturbed
B—C and C horizons, particularly for craters formed less
than a few thousand years ago (Gelinas, 1986). The
maximum grain size of the matrix sand in layers 1
through 3 is greater than in the laterally adjoining and
overlying undisturbed materials in some ﬁlled craters.

The following phases have been interpreted in the
formation of the ﬁlled sand-blow craters: (1) after
earthquake-induced liquefaction at depth, a large hole is
excavated at the surface by the violent upward discharge
of the liqueﬁed mixture of sand and water; (2) a sand rim
accumulates around the hole by continued expulsion of
liqueﬁed sand and water after the violent discharge; (3)
sand, soil clasts, and water are churned in the lower part
of the bowl, followed by settling of the larger clasts and
formation of the graded-ﬁll sequence of sediments; and

 

 

100% sand

0% silt

10% clay

o undisturbed soils

0 liquefaction features

FIGURE 5. —Ternary diagram showing grain sizes of paired samples at
identical depths inside and outside liquefaction craters.

 

(4) the crater is intermittently ﬁlled by adjacent surface
materials to form the thin stratiﬁed-ﬁll sequence, during
the weeks to years after the eruption. In the craters

8 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

predating the 1886 earthquake, the sand blanket ejected
from the crater is indistinguishable in the ﬁeld from the
surface and near-surface (A, E, and Bh) soil horizons,
because the blanket has been incorporated within these
soil horizons.

This interpretation of origin of the ﬁlled craters is
based in part on comparison with photographs and
descriptions of 1886 craters. The ﬁlled craters are similar
to the 1886 craters shown in ﬁgure 3 in that the dimen-
sions (diameters) at the surface are about the same and
the depths to the contact of the stratiﬁed ﬁll to the
graded ﬁll (see ﬁg. 4) are about the same. The lowermost
stratiﬁed layers of ﬁlled craters (layer 4) are composed of
humate-stained, slightly clayey or clean sand that gen—
erally contains a few clasts of Bh horizon material and, in
a few ﬁlled craters, clasts of C horizon sand. (See ﬁg. 4.)
Presence of the intact very weak, friable clasts of C
horizon sand suggests very rapid initial inﬁlling of the
crater, as illustrated in ﬁgure 3B. In the ﬁlled craters,
clasts of C—horizon material ﬁrst must have been ejected
onto the adjacent ground surface or walls of the crater
before falling back or washing back into the bottom of the
crater and being deposited within the beds of layer 4. We
know of no other reasonable explanation of the means by
which these clasts could have been transported to the
stratiﬁed zone.

The presence of friable, angular clasts of C- and
Bh-horizon material in the graded-ﬁll portion is consis-
tent with a short—lived, churning type of upwelling from
the vent. Water commonly ﬂows for a day or so from
earthquake-induced sand blows. The violent, boiling
phase is much shorter in duration. Hence, the presence
of friable clasts argues against a long-term artesian
spring origin for these features; a spring-induced long-
term churning phase, lasting days, that slowly dimin-
ishes in ﬂow would abrade, round, and (or) disintegrate
the clasts of Bh- and C—horizon material. Short-term,
nonearthquake-related springs have been eliminated as a
possible mechanism by Which craters are formed along
the crest of the beach ridges, because such springs
cannot form in this topographic-geologic setting. (This
point is discussed in detail later.)

Our interpretation of the origin of the craters is also
supported by the presence of sand-ﬁlled tabular frac-
tures, whose overall shape and dimensions strongly
suggest that they are “incipient craters.” These fractures
are rather common at some places in the meizoseismal
zone of the 1886 earthquake, where craters are plentiful.
Figure 6 shows V- and U-shaped fractures (ﬁg. 6A also
shows a connecting vent) that are ﬁlled with sand we
believe was transported upward from depth, on the basis
of the freshness of minerals in the fractures. The frac-
tures, which are tabular, generally widen with depth

 

until they connect to a single, near-vertical large sand-
ﬁlled ﬁssure (that is, a vent). The sand-ﬁlled fractures
probably represent the early phase of development of
craters; for the features in ﬁgure 6, however, the upward
forces were too weak to excavate the overlying material.

It is possible that liquefaction led to the production of
craterlets because of a fortuitous combination of sedi-
ment properties in and above the zone that liqueﬁed
during earthquake shaking. The source beds that lique-
ﬁed were exceptionally susceptible of liquefaction, in
that generally they were very loose (engineering sense),
ﬁne-grained, uniformly sized, and free of clay (Dickenson
and others, 1988); these properties would cause the
source beds to liquefy abruptly and, once liqueﬁed, to
ﬂow readily (Seed and others, 1983; Youd, 1973). We
suspect that the liqueﬁed sand strata suddenly applied a
large point force to the overlying sediment (through a
hole or weak zone such as that left by a decayed root),
causing a V— or U—shaped crack to form, through which
the liqueﬁed sand violently vented because of its excep-
tional ability to ﬂow. The V— and U-shaped cracks
occurred because overlying sediment is humate ce-
mented, has no pronounced planes of weakness, and is
very brittle; the process is similar to formation conchoi-
dal fracture in an isotropic, brittle medium, caused by
the application of a point load.

In summary, the ﬁlled craters that we found have a
morphology that is consistent with descriptions and
photographs of craters caused by the 1886 earthquake.
The general geologic setting (Pleistocene beach crests
and ﬂanks) and the locations of swarms of sand blows
reported by geologists immediately after the 1886 earth-
quake (Peters and Herrmann, 1986) are very near or
coincident With two sand-blow sites found during this
study (HW and ARP on ﬁg. 1). The morphology of the
walls of the craters, the stratigraphy of the ﬁll of the
craters, and the evidence that they formed in relatively
sudden, discrete episodes serve to demonstrate that the
ﬁlled craters are the result of earthquake-induced lique-
faction. Alternate origins, such as short- or long-term
artesian springs, ocean wave pounding, thrown trees, or
other mechanisms that have the potential to create
similar features are discussed in detail in the section
entitled “Other Possible Origins.”

VENTED-SAND VOLCANOES

Figure 7 is a schematic cross section of a pre-1886
vented-sand volcano. Where the zone that liqueﬁed at
depth is overlain by a thick clay-rich stratum, a sand
blow typically is expressed on the ground surface as a
mound of sand. The thickest part of the mound ranges
from a few centimeters to as much as 0.25 m. At a few

COASTAL SOUTH CAROLINA

 

I:-
Shovel

/ Pre-1886-filled
sand-blow crater

BANK
OVERHANG

fissur/ ~ / B—C

     

  

- §

 

 

/;C:'_—.'___
\A _ ./:—.ZZ£—Z—
0 20 INCHES
FLOOR OF DITCH 0 30 CENTIMETERS
A ND VERTICAL EXAGGERATION

 

 

BANK
OVERHANG

/Brownish—black,
massive Bh-horizon sand

Sand-filled fissure Sand-filled fissure
' bot : _:

.-——@ ﬂ ' —+»+ ——~—'_Ac|astsof

20 INCHES

   

 

B FLOOR 0F DITCH o 30 CENTIMETEHS
N0 VERTICAL EXAGGERATION

 

 

FIGURE 6. —Schematic cross sections showing sand-ﬁlled ﬁssures interpreted as resulting from liquefaction and flowage during the 1886
earthquake. Light-yellow, intruded sand in ﬁssures is determined to have been vented from a depth of 6 m, on the basis of grain—size
and mineralogical data. A, V-shaped sand-ﬁlled ﬁssure and vent. Fissure cuts soil horizon developed on pre-1886 ﬁlled sand-blow
crater. B, U-shaped sand-ﬁlled ﬁssure.

10 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

 

 

 

 

       

______1___#__

:Clasts of clay-bearing stratum _

 

     
  
  
  
    
   
    

:maximum depthE
_ to unweathered :
clean sand in vent:

mime—athe—reﬁlgn ;n_d—:::
"t —_—to highly weathered ::
' EOE-Leamgﬁni—E} _ _ _ _ 7 A

: — Eynweathe'eEZ—EEEEEEEi
—_—_—_ clean sand,_—_—_:—;—__—_—_—_:

7 :structureless:—_ :_ # _—_—_ __

 

 

 

 

FIGURE 7. —Schematic vertical section of representative vented-sand volcano.

localities, mounds have discernible bedding. The mounds
are generally thickest above their junctions With the
widest steeply dipping sand-ﬁlled dikes (that is, feeder
vents) that extend downward through the clay-bearing
stratum. Often the sand in the mound and in the dikes
contains clasts of the clay-bearing stratum that have
been torn from the walls of the vent. The sand in the
dikes is structureless and has no discernible bedding; the
sand in the central parts of the dikes may be better
sorted and coarser than that near the edges. Dike widths
range from about 0.25 to 1 In.

An 1886 sand blow of this type was described by Cox
(1984). Similar sand blows have been observed at many
places around the world, including the New Madrid
seismic zone (to be discussed later). Figure 8 shows a
variant of this type of sand blow, in which the vented
sand has coalesced into a continuous surface cover; some
of the white surface sand is pedogenic E horizon, which
is virtually indistinguishable from the older, vented
sand. The volume of ﬂuidized sand expelled to the
surface has been so large at some places that the
clay-bearing stratum has downdropped noticeably. At
the site shown in ﬁgure 8, for example, comparison of soil
structure and clay content of the undisturbed soil with
the clay—bearing stratum above the hand shovel showed
that the stratum above the hand shovel had been down-
dropped about 1 m.

 

Near-surface sediments in the depth range of 1.5 to 2
m are commonly so intensely weathered and discolored
that textural analysis is required to determine a possible
origin by venting. A useful (but not sufﬁcient) test for
venting is comparison of the coarsest sand fraction in the
suspected vent With the grain size of sand in the laterally
adjoining clay-bearing stratum or soil horizon. An earth-
quake origin for the sand is not considered likely unless
the coarsest sand in the dike is signiﬁcantly different in
size from that in the crosscut clay-bearing stratum. Such
a textural difference cannot be the result of soil—forming
processes. In addition, there must be no possibility that
the coarsest sand fraction has been introduced from
above the vent. At three vented-sand volcano sites
shown on ﬁgure 1 (sites CH, BR, and SAN) that have
been interpreted as earthquake induced, other evidence
of venting is also present. Examples include clay-bearing
clasts in the sand mound combined with sand—ﬁlled sills
and dikes that Widen downward and, at depth, extend
into slightly weathered and unweathered sediments. At
some places, the sand-ﬁlled dikes and sills cut through
ground that appears to have been shattered (that is,
irregularly, intensively fractured and intruded), which is
very suggestive of forceful intrusion.

Features similar to the vented-sand volcanoes, but for
which springs or other nonearthquake sources cannot be
easily eliminated as possible origins, are common in

COASTAL SOUTH CAROLINA 11

 

FIGURE 8. ——Vented-sand volcanoes that have coalesced to form a continuous sheet of sand on the ground surface. At least two
episodes of venting separated by long periods of time are represented. Note irregular pattern of fracturing. (Shovel is
60 cm long.)

many lowland areas. These sites have not been included
on ﬁgure 1.

FEATURES SHOWING EVIDENCE OF LATERAL SPREADS OR
GROUND OSCILLATIONS

Lateral spreads3 were commonplace features of the
1886 earthquake (Peters and Herrmann, 1986). Most of
these spreads developed in ﬂuvial terrace deposits, bor-
dering streams in low areas. Obvious surface evidence of
lateral spreads has not persisted to the present, and
ditches and pits are so rare in the wet locales where
lateral spreads formed that none has been found. How—
ever, along the ﬂanks of some Pleistocene beaches,
reverse shears are present that may indicate an incipient
formation of lateral spreads or ground oscillations. Shear
displacements commonly range from 1 to 4 cm. Slope of
the ground surface is typically less than 1 percent for
hundreds of meters downslope or upslope. These slopes
are so gentle and the possibility of high artesian pres-
sures is so remote that gravity-induced slumping is
virtually impossible. Some of these reverse shears
almost certainly have an origin in earthquake-induced
liquefaction. At one site, for example, reverse shears

3A lateral spread is a landslide formed by a laterally moving slab, commonly
of large areal extent; earthquake-induced lateral spreads commonly form where
surface slopes range between 0.5 and 5 percent (Youd, 1978). The presence of
lateral spreads generally indicates a large areal extent of liqueﬁed material.

 

dipping in opposite directions (toward one another)
formed about 10 m apart in the stratum that liqueﬁed
during shaking, and sand blows traceable to this liqueﬁed
stratum formed between the shears; the most likely
mechanism that could have formed the opposite facing
shears was alternating directions of ground movement,
caused by earthquake oscillations, in the liqueﬁeWa—
tum. (Such a mechanism is illustrated as ﬁg. 2—11‘in
“Liquefaction of Soils during Earthquakes,” published
by the National Academy Press, Washington, DC, in
1985.) v

Reverse shears can occur as isolated features but ar
generally found in association with sand blows. Figure 9
shows the typical relationship for ﬁlled craters. A
reverse shear is present on the downslope side of the
crater, and the upthrown block cuts the C, B—C, and Bh
horizons. The shear formed prior to venting of sand from
the source stratum at depth, because the vent is not cut
or distorted.

At some sites, the shears along crater edges could
have formed only in response to earthquake-induced
lateral spread movement because the shears are trace-
able into and along the bedding of the stratum that
liqueﬁed. Only rarely is an exposure deep enough to
allow determination of whether the shear goes into a
stratum that liqueﬁed, and thus an earthquake origin
cannot be conﬁdently assigned to all sites. However, we
are of the opinion that these reverse shears along crater
edges suggest an earthquake origin, even Where the

12 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

 

Ground surface

     
  
  
   
 
 

   
   

\ 7:?
\\m\

 
    

Bh or anic-matter-rich sand /,3_'

Undisturbed

Layer 3: Sand and
many small clasts

I '. j scattered

 

 

N0 VERTICAL EXAGGERATION

Liam Si .Strliétureleé$ sand 55

‘ :IL'ayéEz'; Saha' aha
_'clean sand and 5 I;

3'. Bil-horizon materialf j'

. :Ej/entsonaim 6%}

4—— Downslope (0.5—2 percent)

   

   
 

  

    

     
  
  
  
 
  

Undisturbed

/////°/

1 Layer 4: Bedded sequence
1/ of alternating clean and
_ _"__ __ _'_‘__, ; _/\c<)rgan|c-matter-r|ch sand
ragga W/ﬂﬁ/j
'” 5/! z

  
      

crater
wall

clasts of .3

5Q

@Q

-__ @: :iaﬁ 1—2 SEid'and IaTgeFaEtE:

    
  
  

.' {#:of Bh-horizon materialii—L

   
  

 

 

FIGURE 9. —Schematic cross section of ﬁlled sand—blow crater, illustrating aspects possibly associable with downslope movement. A reverse shear
soil is along the lower left side of the ﬁlled crater; an underturned edge forms the lower right part of the ﬁlled crater beneath the B—C and
C soil horizons; large clasts along the right side are much higher than in other soil parts of the ﬁlled crater.

shears have formed on gently sloping ground (less than 2
percent) as much as 5 m below the beach crest.

Another aspect of ﬁlled craters that possibly indicates
an incipient lateral spread is also illustrated by the
position of the clasts along the right side of the crater of
ﬁgure 9. Large clasts occupy the region from the vent to
the vicinity of the bedded sequence (layer 4). The lower
part of this clast zone turns beneath the B—C and C
horizons. These clast relations are found in perhaps 25 to
50 percent of craters having reverse shears. The consis-
tent positional relationship of underturned edge with
clasts along the crater edge suggests control by minor
downslope movement, which took place while the sedi-
ments in the crater were still ﬂuid enough to segregate
slightly.

REGIONAL DISTRIBUTION OF SAND BLOWS

The regional distribution of sand blows is systematic
and predictable in coastal South Carolina, in terms of size
and number of sand blows. This distribution conforms

 

with the pattern that would be anticipated from an
earthquake origin.

Numerous worldwide observations show that the num-
ber, size, and type of earthquake-induced liquefaction
features are a reﬂection of three principal factors. Of
primary importance are (1) severity of shaking (con-
trolled by earthquake accelerations and duration (Seed
and Idriss, 1982)), (2) liquefaction susceptibility (con-
trolled largely by state of compactness and cementation
(Youd and Perkins (1978)), and (3) local geologic controls
(affected primarily by depth to water table, thickness of
impermeable sediments above liqueﬁed stratum, defor—
mation properties, permeability of sediments above liq-
ueﬁed stratum, and permeability of liqueﬁed stratum
(Ishihara, 1985; Obermeier, 1989; Obermeier and others,
1986)). The local geologic controls on the Pleistocene
beaches in coastal South Carolina are constrained within
a very narrow range because (1) bedrock accelerations
should have been ampliﬁed about the same amount at
many places by previous earthquakes, (2) the liquefac-
tion susceptibility of the source sand beds has been about
the same, and (3) local geologic controls on morphology

COASTAL SOUTH CAROLINA 13

and size of sand blows on beach ridges has been about the
same, for a time period extending at least throughout the
Holocene (Obermeier and others, 1989; also discussed
later). Thus, in the vicinity of the epicentral region, the
sand blows should generally be largest (have largest
diameters) and be most abundant. This hypothesis is
conﬁrmed in the Charleston region by observations of
effects of the 1886 earthquake (Peters and Herrmann,
1986). Sand blows produced by the 1886 earthquake were
abundant Within the 1886 meizoseismal zone, which was
a region about 35 km Wide and 50 km long. Beyond this
zone, the sand blows were smaller and scattered. Only
rare, small sand blows formed more than 10 km beyond
the meizoseismal boundary. Because all the factors
involving sand-blow production are about the same in
many places throughout coastal South Carolina, this
same relation of sand blows to epicenter location should
hold true for pre—1886 Holocene earthquakes.

In particular, it has been found that for craters having
one apparent age (based on soil proﬁle) and whose
diameters (measured near the ground surface) are less
than about 1 m, there are at most two or three craters
exposed in a nearby l-km-long ditch cutting across beach
ridge deposits. Wherever maximum diameters of craters
having one apparent age are about 2 m, there are more
(as many as 5 to 10) craters exposed in a nearby
1-km—long ditch. Wherever maximum diameters of cra-
ters having one apparent age are 3 m or larger, there are
a greater number (as many as 20) exposed in a nearby
l-km-long ditch. This well-deﬁned relation between
number of sand blows and sand-blow diameters is con-
sistent with an earthquake origin in our opinion.

We emphasize that a single, isolated feature that
appears to be earthquake-induced would not be inter-
preted by us to be compelling evidence for prior earth-
quake shaking. Compelling evidence requires at least
several features, scattered over a region of at least
several square kilometers.

OTHER POSSIBLE ORIGINS

Origins other than earthquakes for ﬁlled sand-blow
craters and vented-sand volcanoes that have been con-
sidered include compaction-induced dewatering and soft-
sediment deformation; artesian springs; landslides; ﬁll-
ings in decayed stump and root holes; ground disruption
by fallen, root—wadded trees; and liquefaction caused by
storm—induced, ocean-wave pounding. Criteria for
assessing each of these potential sources are discussed
below.

Compaction—induced dewatering and defamation.—
Syndepositional and postdepositional dewatering by
compaction occurs in sediments during or shortly after
their deposition. This dewatering generally takes place

 

in response to rapid deposition of sediments (especially
coarser, denser sediments) above very soft, clay-rich
sediments, which causes buildup of pore-water pressure
and gravitational instability (Dzulynski and Walton,
1965; Allen, 1984; Lowe, 1975, 1976); deposition of silts or
ﬁne sands rapid enough to cause buildup of pore-water
pressure can also lead to instability (Dzulynski and
Smith, 1963; Sanders, 1960). The kinds of sedimentary
structures formed by these processes include load struc-
tures, dish structures, convolute lamination, sand dikes,
and faults (Pettijohn and Potter, 1964). In our study,
earthquake—interpreted features generally have soil hori—
zons that are much thinner and less well developed (and
thus younger) than soil horizons on laterally adjoining,
undisturbed parent sediments; in the relatively few
instances where both the parent sediments and crosscut-
ting features have essentially the same degree of soil
development, the earthquake—interpreted features con—
tain clasts of Bh material at depths generally about 0.5 m
below the laterally adjoining Bh material in the parent
sediment. The apparent difference in age between the
parent sediment and ﬁlled-crater or vented-sand volcano
determined on the basis of soil development, combined
with lack of reason to suspect sudden, nonseismic surface
stressing or long-term pore-water pressure buildup, is
sufﬁcient reason to eliminate postdepositional dewater-
ing as a source mechanism.

Artesiom springs. —The regional and local topographic
and ground—water setting of many sites rules out any
signiﬁcant likelihood for the occurrence of a sudden,
strong increase in the hydraulic force of a spring. The
beach crests are generally ﬂat lying for many kilometers
along their crests, and the crests are well above the
lagoonal deposits. Where ﬁlled craters are found on
ancient beaches, a short- or long-term spring origin is not
believed possible if the following conditions are met: (1)
the ﬁlled craters cut humate-rich sandy soils that are also
cut With numerous, highly permeable sand-ﬁlled root
holes and burrows that extend well into the C horizon, (2)
the ﬁlled craters are much above the lagoonal deposits
adjoining the beach (as illustrated in ﬁg. 2), and (3) the
ﬁlled craters are Within 1 to 2 m of the beach crest.
Short-term springs induced by a hurricane deluge (or any
other mechanism) have not been observed where these
criteria are met.

Short-term springs are suspected to be rather common
in lowland areas of lagoonal deposits, however. Features
interpreted as earthquake-induced sand volcanoes are
restricted to sites where artesian springs are thought to
have been unlikely, and in addition, there is other
evidence for an earthquake mechanism, such as frac-
tured ground (as illustrated in ﬁg. 8). Typical sites are
elevated locations near deeply entrenched rivers, where
artesian pressures would have been relieved by lateral

14 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

ﬂow to the rivers rather than by vertical ﬂow to higher
elevations, which would have been required to form the
volcanoes.

Landslides. —All sites on ﬁgure 1 are on level or
nearly level ground, hundreds of meters from any steep
slopes. Downslope movement on these nearly level sites
could be initiated only by seismic shaking, considering
relations between surface topography, ground-water
setting, and strength of materials.

Fillings in root holes—Holes caused by decayed
stumps and roots, and later ﬁlled with clean sand, are
common in the study area. Although the ﬁlled holes are
generally circular in plan view, many display poorly
deﬁned layers of clay mineral segregation (due to weath-
ering) and gradually taper downward; they do not have
the well-deﬁned laminar bedding of the ﬁlled craters and
underlying graded zone of sediments. Typically, maxi-
mum depths of root penetration are about 2 to 3 m.

Ground disruption by trees. —Tree throw in the South
Carolina low country is generally restricted to hardwood
species having wide, shallow root systems. These trees
may blow over in wind storms and rip up a wad of
sediment caught within their root mats, thus creating
pits and mounds. Taprooted pines and cypress trees very
rarely tend to throw; instead, these species break off
near the ground surface. Pits excavated by hardwood
tree throw tend to be shallow (usually less than 50—100
cm deep); when ﬁlled, they do not contain sediment
introduced from depth, and they almost always lack the
orderly internal stratigraphy of liquefaction craters. At
some few places, though, pits excavated by thrown trees
can be distinguished from craters excavated by
earthquake—induced liquefaction only by the presence of
feeder vents or by the presence in the crater of sediment
(generally sand or silt) that has been introduced from
strata beneath the base of the crater. Veriﬁcation that
sediment has been introduced from depth may require
mineralogical, weathering, and grain-size analysis (Geli-
nas, 1986).

Ocean-wave pounding. —The disruption of subaerially
formed soils shows that liquefaction occurred long after
deposition of the parent sediments. Storm-induced,
ocean-wave poundings are not a credible mechanism of
liquefaction at most of the widely scattered sites because
the elevations are well above modern sea level (up to 15
m) and are too far inland. In addition, no sedimentary
records indicative of inland surges of the ocean, such as
soil horizons buried by storm-deposited sediments, have
been found.

FEATURES OF WEATHERING ORIGIN

A wide variety of features produced by chemical
weathering mimic earthquake-induced liquefaction fea-
tures. Such weathering features include the white, E-

 

horizon sand that commonly blankets the surface; pedo—
genic tongues; and BE or fragipan horizons.
Distinguishing liquefaction features from weathering
features is much more difﬁcult where older liquefaction
features have been extensively weathered.

A loose, clean, white sand blanket covers large parts of
the South Carolina low country that is underlain by
sandy sediments of barrier beach and nearshore marine
origin. In undisturbed soil proﬁles, the white sand under-
lies a thin, 0- to 15-cm-thick, dark-gray A horizon.
Although some of this white sand has been periodically
remobilized by surﬁcial processes, it is a pedogenic E
horizon, formed by weathering and leaching of the under—
lying sandy sediments (Gamble, 1965). The pedogenic
origin of the E horizon is demonstrable by its eluvial-
illuvial relationship with the underlying B horizon. Clays
and labile minerals have been removed from the E
horizon, and weathering products have been deposited in
the B horizon; however, particle-size distribution and
resistant heavy-mineral percentages are nearly identical
in both horizons. In addition, the boundary between the
E and B horizons is commonly abrupt and irregular and
is characterized by narrow, near-vertical tongues of
white E horizon penetrating downward into the B hori-
zon. Laterally continuous, interpenetrating boundaries
of this type are more likely to be indicative of geochem-
ical leaching rather than clastic deposition. At some sites
resistant sedimentary features, such as thin pebble beds,
pass through the horizon boundary, showing that it is not
a sedimentary contact.

In summary, although a blanket of ejected white sand
often exists at earthquake liquefaction sites, all white
sand blankets are not necessarily formed by earthquake
liquefaction. In younger liquefaction features, bedding
within ejected sand blankets helps to demonstrate lique-
faction origin if other ﬂuvial and eolian origins can be
rejected. With age, the usefulness of this criterion
decreases as soil mixing by ﬂora and fauna destroys
bedding, making ejected sand blankets appear superﬁ-
cially like massive pedogenic E horizons.

The pedogenic boundary between E and B horizons
can be gradational or quite sharp and can also be highly
convoluted. Gamble (1965) describes boundaries between
loose, white E-horizon sand that grades over 1 to 2 mm
of depth to brown, clayey sand. The boundary also
commonly has 2- to 3-cm-wide, 5- to 10-cm-long tongues
of E horizon descending vertically into the underlying B
horizon. Locally, we have observed tongues of E-horizon
sand that extend more than a meter into thick, red to
brown (7.5YR 5/6—5YR 5/8), clayey B horizon (ﬁg. 10).
Tongues of this size and shape can give the impression of
fractured and brecciated ground and might be mistaken
for liquefaction features unless examined carefully for
sedimentary characteristics.

 

COASTAL SOUTH CAROLINA 15

 

 

Masswe. light-gray to
right-yeiIMishmbw :
‘ senate *9:va n

 

Ground surface
A horizon

 

E horizon

 

f BE horizon

 

B horizon

 

C horizon

 

 

 

A Samples from B horizon

0 Samples from tongues

 

 

 

 

100 | i l 6
E 90 — é _
t 80 _ A samples fromBhorizon _
3 0 samples from tongues
8 70 -
2’ so — A _
g 50 — g —
a;
a. 40 — _
‘3
'5 30 — —
2
3 2 _ _
E 0
5 10 — —
0 A Q | J |
-1 0 l 2 3 4 5
B Particle size, phi units

The single ﬁeld criterion that is most useful for distin-
guishing pedogenic tongues from liquefaction features is
the downward closing of the narrow, nearly vertical sand
bodies. Pedogenic tongues almost always close down-
ward. Alternatively, if they connect at depth with a
source bed, an earthquake-induced liquefaction origin is
possible. Unfortunately, many exposures are too shallow
to allow use of this criterion.

[1 20 INCHES

O 50 CENTIMETERS
N0 VERTICAL EXAGBERATION

FIGURE 10.—Schematic sketch and grain-sized data for pedogenic
tonguing. A, Pedogenic tonguing of BE- and E-horizon sand into
underlying B horizon of an exposure in a drainage ditch near
Rincon, Ga. (near Savannah). B, Cumulative particle-size distri-
butions for sand fraction samples from Rincon, Ga., shown in A.
B-horizon samples (triangles) are virtually indistinguishable from
tongue samples (circles), indicating that the tongues are probably
not vented sand.

Pedogenic tongues can range in morphology from
tubular (Gamble, 1965) to planar (deﬁning B horizon
polygons; Nettleton and others, 1968a). Hence, tongue
morphology is tenuous evidence for pedogenic or earth-
quake origin.

An example of pedogenic tonguing is shown in ﬁgure
10A, which is a schematic drawing from an exposure in a
drainage ditch near Rincon, Ga. The tongues extend

16 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

  
    
        
 
 

Bhir horizon

0 20 INCHES

Approximate level of seasonal
ground-water table, pre-ditch excavation—>l

0 so CENTIMETERS s'umped “dime“

N0 VERTICAL EXAGGERATIDN

 

7 le, masszve, we —sorte san , ,
' '.very loose and friable ‘ , . -.

  

. Ehorizon '

Grayish-black, massive Bh-horizon sand,

I ‘— -‘~ 4 '
J @5991 144/

Dark»reddish-brown,
massive Bhir-horizon
sand, very firm

  

 

 

 

FIGURE 11. —Cross section through a white, pedogenic sand tongue from an exposure in a drainage ditch at site HW, near Hollywood, 8.0. Prime
designations on E and Bh horizons indicate the lower horizons of a bisequal weathering proﬁle. Bhir is iron—enriched Bh material.

downward from a massive, pale-brown to gray E hori-
zon and narrow with depth. Tongues may form prefer-
entially along soil structure polygon boundaries (ped
boundaries) or where tap roots have penetrated the B
horizon. Once formed, tongues become the loci of con-
centrated weathering because the higher permeability of
the leached E-horizon sand guides inﬁltrating solutions
through the clayey B horizon. Increased weathering in
the tongues may alter the particle-size distribution of the
original sediment either by producing or destroying
minerals in the ﬁner particle-size fractions. Therefore, a
comparison of the particle-size gradation exclusive of the
clay and silt fractions (less than 63 micron) may show
whether the tongue sand has been injected from a
separate source. Sand size and cumulative particle-size
distributions of six samples taken from areas marked in
ﬁgure 10A are shown in ﬁgure 103. The extreme simi-
larity between the tongue and the nontongue portions
strongly suggests that the tongues were weathered in
place, although there is a remote possibility that injected
sand could have the same sand-particle-size distribution.
Alternatively, a liquefaction origin would be more
strongly suggested if the tongue sand-particle-size dis-
tribution was distinctly different from that of the adja-
cent soil.

Other evidence for pedogenic versus liquefaction ori-
gin of tonguelike features is the mineralogy of the sand
fraction. Because weathering is concentrated in the
E-horizon tongues, the presence of abundant fresh, labile
minerals, relative to adjacent soil mineralogy, would be

 

good evidence for injection. Conversely, if the tongue
and adjacent soil have similar mineralogies, or if the
tongue has more highly weathered minerals than the
adjacent soil, a pedogenic origin is more probable. By
using mineralogic criteria, Gelinas (1986) assigned a
liquefaction origin to sand vents at the McL, SAN, and
CH sites (ﬁg. 1) by demonstrating that suspected vent
sands have greater abundances of easily weathered
feldspars and hornblendes than the adjacent undisturbed
soil.

Another type of pedogenic tonguing is sketched in
ﬁgure 11 from an exposure in a drainage ditch near site
HW (ﬁg. 1), where the ditch crosses the crest of a
relatively well drained beach ridge. At least 10 features
of this size and orientation were exposed in a 60-m
section of the ditch. In vertical exposures, 10— to 30-
cm-thick tongues of White sand dip gently downslope.
The tongues are surrounded by black to dark-orange,
humate and iron—oxide-stained sand. Some of the near—
horizontal parts of the tongues can be traced for dis-
tances up to 7 m. At the upslope end of each tongue, the
feature abruptly turns up, breaks through the overlying
Bh soil horizon, and ﬂares toward the surface. At the
surface, the tongue is continuous with white sand of the
E horizon. The downslope end of the tongue is commonly
terminated by black-and-orange-stained rind that is con-
tinuous with the Bh horizon; in other cases the tongue-
rind contact becomes diffuse downslope.

The white sand tongue features were originally inter-
preted by us as intrusions of liquiﬁed sand (Gohn and

COASTAL SOUTH CAROLINA 17

others, 1984, p. 11); however, subsequent observations

have led to the conclusion that the tongues are weather—

ing features unrelated to liquefaction events. Essen—

tially, they are extreme convolutions in the pedogenic E

horizon caused by concentrated inﬁltration of soil solu-

tions at places where taproots of trees have broken
through brittle, relatively impermeable Bh horizons.

This interpretation relies on several ﬁeld observations:

0 The tongues are continuous with white E-horizon
sand overlying the Bh horizon. Pedogenic origin of
the E horizon sand is established by satisfaction of
the sedimentologic and mineralogic criteria discussed
previously.

0 The rind surrounding the white tongues is zoned into
black Bh and dark-orange, iron-enriched Bh material
(Bhir) stripes identical to the sequence found at the
top of undisturbed B horizons. The vertical sequence
of horizons in an undisturbed soil proﬁle shows that
dark organic compounds are deposited closer to the
surface than iron oxyhydroxides after solutions inﬁl—
trate through the overlying E horizon. A similar
zonation in the rind surrounding the tongues suggests
that soil solutions are migrating laterally along the
tongues and outward toward their margins.

O The sharp to diffuse contacts between the tongues
and surrounding sands show no appreciable sand-
grain—size variation across them. The only difference
between the tongues and the rinds is that the tongues
contain black and orange colloidal material that
bridges and coats sand grains.

0 No bedding is present in the tongues.

o The ﬂared portion of each tongue (the portion closest
to the surface) is characterized by evidence of a
taproot, which is either a rotted root in-place or a
narrow V-shaped zone of loose, organic-stained sand
subjacent to the White sand tongue. E-horizon
tongues were found in all stages of development,
from incipient leaching and bleaching surrounding
areas adjacent to relatively fresh taproots to wider
and more deeply bleached zones that formed as the
taproots became more decayed.

O The tongues all dip downhill in the direction of the
low-gradient, shallow ground-water ﬂow system.
Mottles around rootlets in the C horizon are also
elongated in downhill directions. Coincidence of the
tongue orientation with ground-water ﬂow direction
suggests that after the inﬁltrating solutions break
through the Bh horizon where taproots have
decayed, the inﬁltrating solutions are entrained in
the shallow ground-water ﬂow system. Dissolved
organics and iron oxyhydroxides in the ground water
are then deposited downgradient as a rind adjacent to
the E-horizon tongues.

 

Another category of pedogenic feature that might be
confused with earthquake-induced liquefaction is the BE
or BE’ horizon. (The prime designation is used to denote
the lower sequence of eluvial and illuvial horizons in a
bisequal proﬁle, in the general order A, E, Bh, E’, BE’,
B (t,x), C.) Bisequal proﬁles with BE’ horizons form at
the transition between poorly drained and moderately
well drained landscape positions (Nettleton and others,
1968a,b). BE or BE’ horizons are transitional between
eluviated E horizons and illuvial B, Bt (argillic), or Bx
(fragipan) horizons and are characterized by irregular to
prismatic bodies of clay-rich sand surrounded by leached,
clean sand. Studies of this type of weathering proﬁle in
the North Carolina coastal plain indicate that the clean
sand forms from the progressive destruction of a clay-
rich Bt horizon (Daniels and others, 1966; Nettleton and
others, 1968a, b; Steele and others, 1969). In this pro-
cess, a Bt horizon is leached at the depth of a ﬂuctuating
water table, which removes clay and labile minerals and
leaves patches of white to gray quartz sand that often
coalesce to deﬁne a polygonal pattern in plan View. This
horizon, composed of patches of leached and unleached
material, is termed an E’ or BE’ horizon, depending on
the extent of leaching. With continued leaching of clay
and labile minerals, the soil volume decreases, resulting
in collapse and the formation of a dense Bx horizon.

In the ditch at the HW site, bisequal soils occur
between red, oxidized spodosols on beach ridge crests
and black, organic-rich spodosols in interridge depres-
sions. For the example shown in ﬁgure 12, the BE’
horizon is a layer of pale-brown, clayey sand and White
sand. The top of the layer is 1 to 2 m beneath the ground
surface, and the layer is 0.3 to 0.5 m thick. Both above
and below the layer is light-colored, well-sorted, ﬁne- to
medium-grained quartz sand. This sand cuts vertically
across the clay layer at many places and creates irregular
masses of the brown clayey sand. In three dimensions it
can be seen that the sand forms irregular vertical walls.
In vertical section, the sand walls are typically about 5 to
10 cm wide but can be as much as 30 cm wide. They are
spaced at both irregular and at regular intervals. The
contacts of the sand and the brown clayey sand are
typically sharp with regard to both color and texture.
The sand-grain sizes in both the leached sand and clayey
sand masses are essentially the same.

Sharp contacts between the leached quartz sand areas
and the pale—gray to pale-brown clayey material origi-
nally suggested that this feature was formed by some
type of ground disruption, possibly liquefaction. How-
ever, the similarity with features attributed to pedogen-
esis elsewhere and the consistent occurrence in horizons
underlying undisturbed pedogenic horizons led us to
interpret these features as pedogenic.

18 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

 

Bank overhang I 2'0 INCHES

l

o——o

l
50 CENTIMETERS
N0 VERTICAL EXAGGEHATION

    

  

Massive Bh horizon sand;

 

    
  
    
 

 

Not exposed

grayish-black grading down c
to dark-reddish-brown / .3
5
.C
.C
a:
White, massive; well-sortéd sand - 8
. B
.C
A
C
O
.u
‘6
.5
Eu
m

33 White, wellr-sortedisand‘,
-' locally stained palebrown

 

  
  

   

 
 

Not scraped

 

 

 

FIGURE 12.—Vertical section through BE’ horizon at site HW, near Hollywood, S.C.

OVERVIEW OF THE SOUTH CAROLINA LIQUEFACTION
STUDY

Pre-1886 sand blows were ﬁrst discovered in 1983
(Gohn and others, 1984; Obermeier and others, 1985)
near the town of Hollywood (site HW, ﬁg 1). There, the
sand blows are exposed in an unusually deep ditch (up to
3 m) for a distance of 5 km. Over a hundred sand blows
have been discovered in this ditch, and the depth of the
ditch has permitted detailed study of entire features to
depths that include source vents. For these reasons, site
HW is generally used for the descriptions of ﬁlled craters
in this paper. Since the discovery at Hollywood, features
we interpret as sand blows have been found at sites
throughout much of coastal South Carolina.

Features at all sites shown on ﬁgure 1 are interpreted
to be of earthquake origin, although the conﬁdence level
differs for various sites. Sites where we have greatest
conﬁdence of an earthquake origin are those where the
following features occur: (1) craters have formed on
topographically high beach crests, (2) numerous craters

 

are present near where craters were reported to have
been abundant during the 1886 earthquake, (3) ground
oscillation shears have formed in opposite directions, (4)
lateral spreads that could not be gravity-induced have
formed and have shears traceable into a liqueﬁed stra-
tum, or (5) shattered ground is cut by numerous sand-
ﬁlled dikes in settings Where high artesian pressures
could not have been involved. Sites for which we have
the highest conﬁdence of an earthquake origin are BR,
AR, HW, ARP, RRR, CH, FM, WV, SAN, 0L, and
SOPO.

All other sites on ﬁgure 1 are ﬁlled craters that are
more than several meters below the beach crest, causing
the conﬁdence level to be lower. Some craters at site
MYRB have reverse shears, however, an occurrence that
makes an earthquake origin seem likely. '

Elimination of all sites from ﬁgure 1 except those in
which we have the highest conﬁdence does not affect our
interpretation of Holocene seismic activity (discussed
subsequently). Our interpretation about earthquake

COASTAL SOUTH CAROLINA 19

ages and relative severity of shaking is based only on
data from the sites for which we are most conﬁdent of an
earthquake origin.

EARTHQUAKE AGES

Craters are typically the only features for which
radiocarbon ages related to earthquake ages can be
generated, because other liquefaction-related features
are not found in association with preserved organic
matter. Three methods have been used to bracket the
times of crater formation (Weems and others, 1986): (1)
radiocarbon ages of woody material (tree limbs or pine
bark) that fell into the open crater soon after crater
formation, (2) dating of roots sheared off at the edge of
the crater (predating crater formation) and dating of
roots that grew into the stratiﬁed ﬁll portion of the
crater (postdating crater formation), and (3) dating of
clasts of Bh material that fell into the graded ﬁll zone of
the crater. The ﬁrst method yields a highly accurate age
for the time of earthquake occurrence, Whereas the other
two yield a broad range of possible ages. Sufﬁcient data
have been collected at site HW (near Charleston; ﬁg. 1)
to show that at least three pre-l886 earthquakes pro-
duced sand blows within the past 7,200 years. Radiocar-
bon dating of a clast of pine bark in a crater at site ARP
(also near Charleston) independently veriﬁes the middle
of these three events. The only deﬁnitive statement
about earthquake recurrence that can presently be made
is that near Charleston there have been at least four
sand—blow—producing (mb probably >5.5, discussed later)
earthquakes within the past 7,200 years (including the
1886 event). Accurate ages of crater formation have been
obtained from some sites more than 100 km from
Charleston. These dates differ from ages near Charles-
ton, thereby suggesting that the craters far from
Charleston originated from epicentral regions also far
from Charleston. At many sites far from Charleston,
there are at least two generations of craters that are long
separated in time of formation.

SHAKING SEVERITY ESTIMATION

Insufﬁcient radiocarbon ages have been collected from
liquefaction features throughout the Carolina coastal
region to deﬁne epicentral regions of separate earth-
quakes. Adequate data have been collected, however, to
estimate the relative shaking severity throughout the
coastal region during the Holocene. This estimate is
provided by measurement of the number and size of
craters at the sites shown on ﬁgure 1.

The methodology for estimating shaking intensity is
based on the premise that the number and size of
liquefaction features are greatest where earthquake shak-
ing has been strongest for a ﬁxed geologic setting and

 

liquefaction susceptibility. The condition of a ﬁxed geo-
logic setting is met almost ideally on many of the
Pleistocene beaches, as discussed in an earlier section.

The condition of a ﬁxed liquefaction susceptibility is
also almost certainly satisﬁed at the widely scattered
sites on ﬁgure 1. Source sands typically are loose (based
on limited Standard Penetration Test data and numerous
observations of ease of augering) and have about the
same thickness. Moreover, the thickness and properties
of nonliqueﬁable sediments overlying the source stratum
lie within a narrow range. It is also a certainty that
recurrences of liquefaction do not greatly diminish the
potential for formation of more large craters in loose
sands. This statement is veriﬁed by the observation that
at site HW there are many large craters that formed in
each of at least three generations of Holocene earth-
quakes, with each generation widely spaced in time.
Thus, at sites in beach deposits on ﬁgure 1, liquefaction
susceptibility is generally high and has not been greatly
reduced by previous occurrences of liquefaction.

The other major variable, depth to the water table, is
about the same from site to site (very shallow depth) and
probably has been shallow throughout the Holocene
(Obermeier and others, 1987). Evidence for location of
the water table throughout the Holocene is provided by
location of the base of the Bh horizon. (See ﬁg. 4 for
location of this horizon at a typical ﬁlled crater.) The
maximum depth of the seasonal water table during the
Holocene is marked very nearly by the base of the Bh
horizon. (The Bh is deﬁned as the subsoil zone of
accumulation of organic matter and is formed in these
soils at the lower limit of vertical inﬁltration of water.)
Throughout the coastal region, the base of the Bh
(generally 0.6 to 1 m below land surface) is nearly
coincident with the present—day water table. Radiocar-
bon ages from the basal Bh horizon are 5,000 to 10,000
years at site HW (Weems and others, 1986, p. 7).
Because these ages are mean residence times of organic
matter in a dynamic system characterized by continuing
vertical inﬁltration of younger organic matter, some of
the organic matter has been there even longer. Thus, it
can be concluded that the water table has been very
shallow throughout the Holocene over wide areas of the
South Carolina Coastal Plain.

HOLOCENE EARTHQUAKE SHAKING

Both the abundance and diameters of pre-1886
Holocene craters are greatest within the 1886 meizoseis-
mal zone for a given age of craters. On the basis of these
criteria, we know that shaking has been much weaker
north of the Santee River (Obermeier and others, 1989).
Intermediate shaking has taken place between Charles-

20 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

ton and the Santee River and also between Beaufort and
the Savannah River.

Conﬁdence in this interpretation is high for the area
between Charleston and Wilmington, because of the
hundreds of kilometers of ditches we searched. Our
conﬁdence is also high for the 1886 meizoseismal zone and
for the area between Beaufort and the Savannah River.
Our conﬁdence is not nearly as high for the area between
the Beaufort and the Edisto River nor for the area south
and southeast of the 1886 meizoseismal zone; this lower
conﬁdence is caused by the limited number of ditches and
pits available for inspection.

Whether or not the pre—1886 Holocene shaking in the
1886 meizoseismal zone is associable with earthquakes

stronger than the 1886 event can be determined only by ,

additional radiocarbon ages for craters at sites far
beyond the 1886 meizoseismal zone.

Based on worldwide observations in the ﬁeld, the
minimum earthquake strength required for liquefaction-
induced features is mb equal to approximately 5; such
features are rare for mb less than 5.25 to 5.5 (Carter and
Seed, 1988). It is likely that an earthquake slightly
stronger than about 5.5 is adequate to produce numerous
small liquefaction features in coastal South Carolina
because of the exceptional liquefaction susceptibility of
many marine sand deposits there (discussed previously
and discussed in Dickenson and others, 1988).

Many of the pre-1886 craters near Charleston that we
have observed probably were not produced by earth-
quakes as small as mb 5.5, however. Many pre—1886
craters are large (diameters of as much as 3—4 m are not
uncommon) in comparison to historical descriptions of
1886 craters (Dutton, 1889) and our observations of the
sizes of 1886 craters. As a result, we suspect that some of
the pre—1886 craters were formed by earthquake shaking
that was stronger at those sites than in 1886. Direct
comparison of 1886 and pre-1886 crater sizes cannot be
used for estimating earthquake magnitudes, however,
because of the lack of knowledge about distances
between craters and their associated epicentral regions
and lack of knowledge about changes in liquefaction
susceptibility caused by a previous occurrence of lique—
faction. About the only comment that can be made with
much assurance is that the prehistoric craters were very
likely the result of earthquakes stronger than mb 5.5;
some of the earthquakes probably were much stronger
because of the widespread distribution of large craters.

NEW MADRID SEISMIC ZONE

The succession of great earthquakes collectively des-
ignated as the New Madrid earthquakes of 1811—12
caused severe and widespread ground failure throughout

 

a large area near the Mississippi River in southeastern
Missouri, northeastern Arkansas, western Kentucky,
and western Tennessee. The earthquakes caused multi-
tudes of ﬁssures and sand blows (Fuller, 1912), set off
numerous landslides (Fuller, 1912; Jibson, 1985), and
caused localized doming and submergence of the ground
surface (Russ, 1982). Earthquake effects were particu-
larly severe in the St. Francis Basin (ﬁg. 13). According
to the earliest extensive documentary account (Fuller,
1912), three major shocks occurred: December 16, 1811;
January 23, 1812; and February 7, 1812. The surface-
wave magnitudes (Ms) of these earthquakes are esti-
mated to have been between 8.3 and 8.8 (Nuttli, 1983a).
Estimated Modiﬁed Mercalli (MM) intensities are XI to
XII throughout large areas near the epicenters (Nuttli,
1973); these intensities are regional values based largely
on historical accounts of liquefaction-induced ground
failure. No faults associated with the 1811—12 earth-
quakes have been found that cut through alluvium, loess,
or older strata to the surface. The epicenters are esti-
mated to lie within the large area of sand blows near the
Mississippi River, shown on ﬁgure 14. This interpreta-
tion of epicenters is based on the study by Fuller; a study
of MM intensities by Nuttli (1973); recent geologic,
geophysical and modern seismicity studies (McKeown
and Pakiser, 1982; Crone and others, 1985); and
engineering—geologic studies of factors controlling the
distribution of sand blows (Obermeier, 1989). (The epi-
central regions during the 1811—12 earthquake and the
region of epicenters on ﬁgure 14 deﬁne the “New Madrid
seismic zone.”)

Figure 14 is a map by Obermeier (1988) that shows the
percent of the ground surface covered by sand vented to
the surface in the St. Francis Basin. The map is the latest
of a series of maps (Fuller, 1912; Saucier, 1977; Heyl and
McKeown, 1978) showing sand blows throughout the
alluvial lowlands of the basin. Signiﬁcant differences
appear between each of the maps, particularly north of
the town of New Madrid. These differences occur for a
number of reasons, the most likely being that (1) some
surface soils are so sandy that sand blows can be
observed only on the older aerial photographs used by
Obermeier, (2) the modern farming practice of land-
leveling has destroyed many sand blows since the 1950’s,
(3) some nonearthquake features cannot be distinguished
from sand blows except by ﬁeld excavations, and (4) the
map by Obermeier extends more to the limits of ground
covered by vented sand, whereas the earlier maps
emphasize areas of abundant sand blows.

Sand-blow deposits and other manifestations of
liquefaction-induced ﬂowage occurred far beyond the
limits of sand-blow deposits shown on ﬁgure 14, but
beyond these limits the deposits were generally

NEW MADRID SEISMIC ZONE 21

 

>2

0 20 MILES

0 20 KILOMETERS

 

    
  

H Mississippi

EXPLANATION

 

 

 

 

Alluvium on small streams

 

 

 

 

 

Mississippi River meander belts

 

Late Wisconsinan glacial outwash

 

 

 

 

 

River
Early Wisconsinan glacial outwash

 

 

 

 

 

Undifferentiated pre-Wisconsinan
depositional terraces

 

 

 

 

 

 

 

 

FIGURE 13. —Late Quaternary alluvial deposits of St. Francis and Western Lowlands Basins (from Saucier, 1974).

restricted to modern ﬂood-plain sediments adjoining
rivers. These very young (late Holocene) sediments
generally have higher liquefaction susceptibility than the
late Wisconsinan and early Holocene ﬂuvial deposits,
which underlie the area of sand-blow deposits shown on
ﬁgure 14. Liquefaction features were especially common-
place in alluvium along some of the small streams of the
Western Lowlands Basin (ﬁg. 13), particularly those
nearest Crowleys Ridge. Sand blows were reported
(McDermott, 1949) in alluvial deposits as far north as
Cahokia, 111. (which is very near to St. Louis, Mo.), and
as far northeast as the Wabash River valley (Street and

 

Nuttli, 1984). Some possible sand blows were formed in
some of the river valleys beyond the limits of sand blows
on ﬁgure 14. For example, Dickey (1985) reports fea-
tures on aerial photographs in the Western Lowlands
that appear to be sand-blow deposits. At many places,
though, sand-blow deposits on modern ﬂood plains have
been covered by a veneer of alluvium deposited since the
1811—12 earthquakes.

In the New Madrid area, visible sand-blow deposits
are still plentiful on the ground surface, and the epicen-
tral regions are relatively well established. These char-
acteristics permit a more detailed evaluation of geologic

22 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

91° 90°

 

37°

36° ‘

 

 

EXPLANATION

m

Upland areas

U

Alluvium with few or
no observed sand blows

 

Alluvium with recognizable
sand-blow deposits (>1 per—
cent ground coverage), but
not major ground coverage

 

Alluvium with more than
25 percent of ground surface
covered by sand-blow deposits

+

Epicenters for three largest
1811—1812 earthquakes
accordingto Nuttli (1979)

/ Fault

/
/ Fault zone
/ //
/

0 20 MILES

0 20 KILOMETERS

 

 

FIGURE 14. —Map showing area covered by vented sand ( Obermeier, 1988), estimated epicenters 0f strongest 1811—12 earthquakes (from Nuttli,
1979), and faults and fault zones (from Hamilton and Zoback, 1982).

controls on the formation of liquefaction features than is
possible in coastal South Carolina.

REGIONAL GEOLOGIC AND SEISMOTECTONIC SETTING

The St. Francis Basin is a topographic lowland con—
taining 30 to 60 m of late Quaternary alluvium, which
originated from deposition associated with the Missis-
sippi and Ohio Rivers and their tributaries. The lowlands
are typically very ﬂat, having about 0.3 to 2 m of local
relief. Numerous small scarps, 2 to 6 m high, result from
depositional or erosional processes. The water table is

 

very high and is within 1 to 2 m of the ground surface at
many places. Elevations in the lowlands exceed 100 m
near Cairo, 111., and decrease southward to less than 60
m near Memphis, Tenn.

Widespread ﬂoods were common prior to the construc-
tion in this century of manmade levees along the Missis-
sippi River. Standing water occupied the lower parts of
the lowlands throughout much of the year before the
levees were built and large drainage ditches were exca-
vated. Until about 50 years ago, dense forests and
cypress swamps covered most of the basin; today the
forests have been replaced by large ﬁelds of soybeans,
rice, and other row crops.

NEW MADRID SEISMIC ZONE

23

 

 

sediments

‘substraturnr'.

Dominantly clay-rich

Braided stream deposits—>k— Meander belt deposits

 

-','_grading downward to“. '. ,_

'.. QOarSe Sand and grave] til-.3.- '..: '..‘~ :- '.. ‘..... .'-- '..: . ;. ‘.. ‘..” _ ~. . . . . _

 

Flood

Natural stage

   
 

Normal E
river
stage

   

 

Crowleys Ridge

 

 

»9 . artesian seepage area E
g 3 Generally clay-rich 0_3 km I . 04,500 m _ g) 8
:- 2 Tertiary deposrts .2 .2
‘5 D: a D:
(B E
l< 40—60 km >
[A 50—80 km —

V =Water level

 

 

 

FIGURE 15. —Schematic east-west cross section showing geologic and ground-water setting of St. Francis Basin.

CHARACTERISTICS OF QUATERNARY ALLUVIUM

Cyclic Pleistocene glaciations directly and indirectly
controlled the origin, character, and distribution of vir-
tually all the Pleistocene deposits in the basin (Saucier,
1974). Although continental glaciers did not extend into
the St. Francis Basin, they supplied large volumes of
glacial meltwater and outwash to southward—ﬂowing
river systems. Up to 60 m of valley ﬁll composed of very
coarse grained, well-graded (engineering sense) glacial
outwash (gravel and sand) were ﬁrst deposited by
aggrading braided streams. After maximum aggrada-
tion, numerous braided-stream terraces were deposited.
The ancestral Mississippi River changed from a braided
to a meandering regimen in early Holocene time, and
since that time, slow aggradation has taken place in most
parts of the valley. Most meander-belt deposits are
medium-grained, well-graded sand; locally, silts and
clays have been laid down.

Sikeston Ridge (ﬁg. 13), the highest terrace of St.
Francis Basin, is early Wisconsinan in age, while the
other braided-stream terraces are late Wisconsinan. The
earlier braided-stream deposits are topographically
higher, have greater relief, and have sandier surfaces
than younger deposits. The older braided-stream depos-
its occur in several terrace sublevels separated by 2 to 6
m. These deposits generally are capped with 3 to 7 m of
silty or clayey sediments (the topstratum) containing
some thin sand strata. The topstratum abruptly overlies
the clean, outwash sands and gravels (the substratum)
over large areas except at the highest sublevels and
interﬂuves. The topstratum was deposited either by
relatively slack streams as individual braided-stream
levels were successively abandoned or by overbank
deposition during widespread ﬂooding of the Mississippi
River and local streams that now occupy the topographic
lows. Substratum sands and gravels generally are found

 

in intercalated strata and lenses ranging from a few
centimeters to a meter thick and tend to become coarser
with depth.

The Mississippi River meander belts represent succes-
sive courses formed by lateral migration of the river.
Most of the meander belt consists of point-bar “accre—
tion” topography of parallel arcuate ridges and swales,
abandoned channels in various stages of ﬁlling, and
natural levees. Point-bar deposits are generally clean,
well-graded sands4 that are capped by the topstratum
except locally at the highest elevations. Point-bar depos-
its of the Mississippi River typically have 2 to 3 m of local
relief. Many abandoned channels are ﬁlled with 30 m or
more of soft clays and silts and are swampy and densely
forested.

Large backswamp areas are found along the margins
of some Mississippi River meander belts (Saucier, 1974).
Backswamp deposits were formed by seasonal ﬂood-
Waters depositing silts and clays in low areas of the ﬂood
plains during the Holocene. The deposits average 12 m in
thickness but are as much as 18 m thick.

The ground-water setting is generally simple, as
shown in ﬁgure 15. Major rivers such as the Mississippi
and St. Francis both supply and drain large volumes of
water in the substratum. Locally, especially within 0.5 to
1 km adjacent to levees, artesian conditions occur many
years during ﬂooding. Beyond this area of ﬂood—related
springs near levees, the streams have cut through the
topstratum, thus preventing artesian conditions. In addi-
tion, throughout the St. Francis Basin many topograph-
ically elevated areas make artesian conditions impossi-
ble. Present ground-water conditions are shown in a
fairly detailed manner in a ground-water report by

4In this paper, the clean sands beneath the ﬁne-grained cap are referred to
as the “substratum,” and the ﬁne-grained cap as the “topstratum,” irrespective of
origin as meander-belt or braided-stream deposits.

24 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

 

   
  
  

I ll

Surface projection
of buried rift complex \(

11/0 \

 

  

+

+ ~ 4 + + 4 + 4

owercrust+ + o + +
+

 

 

 

FIGURE 16.—Conﬁguration of the buried New Madrid Rift Complex
(from Graile and others, 1984). The structurally controlled rivers,
Paleozoic rocks in the cratonic sedimentary basins, and the Missis-
sippi Embayment, all associated with the buried rift complex, are
also shown. Dark areas indicate intrusions near the edge of the
buried rift.

Krinitzsky and Wire (1964); detailed geologic maps by
Saucier (1964) and Smith and Saucier (1971) can be used
to help determine present and Holocene ground-water
conditions. No evidence indicates .that ground-water
conditions (especially the piezometric surface) were sig—
niﬁcantly different during the past few thousand years in
the St. Francis Basin, except locally near manmade
levees during times of ﬂooding.

SEISMOTECTONIC SETTING

The St. Francis Basin lies in the northern Mississippi
Embayment, which is a broad, southward-plunging syn-
cline (ﬁg. 16). The northern embayment is underlain by a
late Precambrian intraplate rift that has been active
periodically since its formation (Braile and others, 1984).
Within this rift lies the New Madrid seismic zone. This
zone is the most seismically active area in the United
States east of the Rocky Mountains. Continuing fault
movement, in response to regional compressive and
perhaps thermal stresses, is the most likely cause of
modern seismicity (Braile and others, 1984).

Since 1812, at least 20 earthquakes having estimated
body-wave magnitudes between 3.8 and 6.2 have
occurred in the region (Nuttli, 1982). Most estimates for
the return period of great earthquakes are between 500
and 700 years (Hopper and others, 1983).

Figure 17 shows locations of instrumentally recorded
earthquakes and principal structural elements in the
New Madrid seismic zone. The trend of modern epicen-
ters extending from Marked Tree, Ark., to Ridgley,
Tenn., coincides at least approximately with the epicen—
ters interpreted by Nuttli (1979) of the three principal

 

shocks of the 1811-12 earthquakes. Subsurface faults are
inferred by seismic-reﬂection data to lie beneath the

most seismically active areas (O’Connell and others,
1982; Crone and others, 1985).

CHARACTERISTICS OF AND CRITERIA FOR EARTHQUAKE-
INDUCED LIQUEFACTION FEATURES

Reports of level—ground and near-level ground failure
features made shortly after the 1811—12 earthquakes
noted great multitudes of vented-sand volcanoes, linear
ﬁssures up to 6 m deep and hundreds of meters long,
craters many meters in diameter, and lateral spreads
hundreds of meters long (for example, see Penick (1976)).
Fuller (1912) described vicinities where these features
were especially abundant. Individual vented—sand volca-
noes and some long linear ﬁssures through which sand
vented are the only features that are still readily visible
on the ground surface. Intruded dikes and sills are
common in walls of deep (>3 to 4 m) drainage ditches in
the clay-rich topstratum. We have not found other types
of level-ground failure features such as deep craters or
lateral spreads having downdropped blocks at the head
such as those described by Fuller (1912, p. 48), primarily
because we have not searched extensively in areas where
these features would have formed (that is, locations in
sandy topstratum for craters and locations near streams
or scarps for lateral spreads). Limited observations in
the New Madrid seismic zone by Wesnousky and others
(1987) and by us of the characteristics of sand blows
formed in a sandy topstratum indicate that open, deep
craters are atypical forms. Instead, intrusions that
extensively shattered the ground near the surface locally
erupted to form shallow craters (up to 2 m deep).

The geologic criteria (previously discussed) for inter-
preting an earthquake origin with regard to speciﬁc
features are identical in both the South Carolina and
New Madrid earthquake areas.

VENTED-SAND VOLCANOES

The vented-sand volcanoes in the New Madrid area are
similar to, though much larger than, those discovered in
South Carolina. (See ﬁg. 7.) Individual vented-sand
volcanoes induced by the 1811—12 earthquakes are dome-
like accumulations of clean sand on the ground surface.
Fuller (1912, p. 79) noted that “the normal blow is a patch
of sand nearly circular in shape, from 8 to 15 feet across,
and 3 to 6 inches high.” Such small sand blows as Fuller
described can rarely be found at present. The sand blows
that are now obvious range from about 0.3 to 1.3 m in
height at the center and thin to a feather-edge at a
diameter of 20 to 60 m.

Most vented-sand volcanoes have a well—deﬁned inter-
nal stratigraphy. Along the base of the vented-sand

NEW MADRID SEISMIC ZONE 25

 

92°00 91°00 00°00 39°00 00°00
I ( I I I (“I
> \ ILLINOIS ﬂp/
H \ \ l/ I“.
38°00' — \ \ .. ,’,, T
37°00 —

 

 

 

 

 

36°00' _
35°00 - ’
I ,3, ,: :, 3;... I L 1‘ :1; L |
EXPLANATION
N Plutons Faults
0 Earthquake epicenters Rift boundaries

FIGURE 17. —Northern Mississippi Embayment (shaded area) showing earthquake epicenters, plutons, rift boundaries, and faults (from Zoback
and others, 1980).

volcano deposits, Where the vented sand spread over the
ground surface, many irregular 1- to 3—cm-long clasts of
topstratum clay (generally a blue-gray, highly plastic
clay) are scattered throughout the sand. The clasts are
largest and most plentiful closest to the vent, which is
beneath the central part of the dome. The basal part of
the sand-blow deposit contains numerous pieces (1—3 cm
in diameter) of rounded charcoal fragments and other
very low density materials vented to the surface from the
substratum. This basal part contains and is overlain by a

very clean, generally medium- to coarse—grained sand.
This sand grades upward to a much ﬁner sand, which is
capped by a clayey, organic silt stratum, 0.5 to 4 cm thick
(ﬁg. 18); this cap may also contain multiple very thin
(l—mm-thick) clay-rich layers (Saucier 1989). The organic
matter in the silt is made up of small pieces of coal,
lignite, and wood. The ﬁning-upward sequence, from the
basal clay clast-bearing sand to the organic silt stratum,
represents the transition from the initially rapid, turbu—
lent eruption to the ﬁnal ebbing ﬂow out of the vent.

26 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

 

FIGURE 18.—Stratigraphy of vented-sand volcano showing organic-
rich silt between two ﬁning-upward sequences of sand. Presence of
two ﬁning-upward sequences is inferred to represent two separate
liquefaction events or reactivation during a single earthquake.
Height of shovel is 40 cm.

Locally, near the vent, the beds may dip steeply toward
the vent if ﬂowage has been sufﬁciently slow. (Dipping
beds in an excavation through vented-sand volcano
deposits are shown in ﬁg. 5 in an article by Newmann-
Mahlkau (1976)). If a violent eruption has taken place,
the lower and central portion of the vented—sand volcano
may be sheared and disrupted. Figure 19 shows
earthquake-induced “eruptive vent” in sand-blow depos-
its of the 1811—12 earthquakes.

A basal bed, several centimeters thick, of pea-sized
clean gravel is the coarsest basal material that we have
found in the sand blows. Within a sand matrix, isolated
gravels up to 3 cm in diameter have been vented; these
coarsest materials were vented only near the earthquake
epicenters. Sand blows far from the epicenters, near the
boundary of sand-blow occurrences (ﬁg. 14), are rather
small and are made up of ﬁner grained material. Here the
basal sands are often ﬁne-grained and ﬁne upward to a
silty sand.

At least two ﬁning-upward sequences characteristi—
cally occur in vented-sand volcanoes near epicentral
regions; each sequence represents a separate occurrence
of venting induced by the 1811—12 earthquake.5 How—
ever, generally only one sequence occurs near the
regional boundary of sand blows, which is shown in
ﬁgure 13.

Not only are the largest sand blows near the epicentral
regions, but in some places the vented sand coalesces
into continuous sheets. The sheets are up to 2 m thick

5Interpretation of 1811—12 earthquakes is based on the absence of a soil
proﬁle in any of the ﬁning-upward sequences.

 

and at several places blanket tens of square kilometers
(for example, just south of the Missouri-Arkansas bound-
ary and near the December 16, 1811, epicenter shown on
ﬁg. 14).

Vented-sand—blow deposits are obvious on aerial pho-
tographs at many places. Figure 20 shows some exam-
ples. The light-colored spots show vented sand blows,
and the light-colored linear features show ﬁssures
through which sand has vented. (Some of the linear
features near streams probably mark edges of lateral
spreads, but we have not conﬁrmed this suspicion in our
study, and therefore in this paper such features are
simply called ﬁssures). Figures 20A and 208 illustrate
how sand blows typically occur at irregular, almost
erratically spaced intervals, and the crazing pattern of
vented sand suggests that the ground is fractured.
Figure 200 illustrates that where less sand is vented to
the surface, the venting is more erratic and is restricted
to localized areas and that relatively small differences in
site characteristics become important controls on vent
locations. Comparison of ﬁgures 20A and 203 illustrates
that sand-blow deposits may be more apparent in the
older photographs than the more recent photographs.

Some surﬁcial features unrelated to earthquakes have
a similar appearance to earthquake-related features
shown in ﬁgure 200. An earthquake origin was attrib-
uted to questionable features only if the features had a
concentration of irregular, large (2—3 cm) clay clasts at
the base (implying a strong hydraulic force); had a
ﬁning-upward clean sand (implying diminishing ﬂow
within a short time period); were irregularly spaced
(suggesting ground fracturing); and (4) were located
where spring ﬂow from uplands or from beneath levees
could not have occurred. (Sand boils formed by flow
beneath levees are discussed later.)

Only limited data have been collected to characterize
vent shapes and sorting and flow structures in the vent
ﬁlling. The vents are frequently steeply to vertically
oriented and are ﬁssure shaped. Vents generally appear
to be ﬁlled with a structureless mixture of silt, sand, and
clay clasts ranging from sand sized to having a length of
up to 20 cm. The clasts have been transported up the
vent and are derived from sidewalls and beds at depth.

FISSURES

Earthquake-induced ﬁssures presently appear on the
ground surface as lines of continuously or discontinuously
vented sand. The ﬁssures are common on level ground
far from any (erosional) scarps but are especially com-
monplace on those parts of terraces immediately above
scarps and on river banks. Vented sand locally forms
small ridges, but the ridges are generally not as high as
nearby vented-sand-volcano deposits and sometimes are

NEW MADRID SEISMIC ZONE

barely discernible. Widths of ﬁssure openings tend to
range between 0.3 m to a feather edge, most being small;

 

 

Plow zone

: Quartzose sand

     
 
  
 

   

. Eruptive
000005300 \
: wo \

 

 
 
  
 

Charcoal

Buried B horizon

25 INCHES
|

 

l
50 CENTIMETERS

 

 

 

 

27

the wider openings almost certainly indicate association
with lateral spreads.

Fissure patterns can be highly sensitive to the local
geologic setting. Fissures have a strong tendency to
follow the crests of point-bar deposits; on braided-stream
deposits having essentially a uniformly thick topstratum,
the ﬁssures are commonly oriented randomly (Ober—
meier, 1989). Topstratum thickness appears to be an
important control on development of ﬁssures; more ﬁs-
sures develop at locations having a thin topstratum.

Long irregular ﬁssures through which sand has
vented, such as those shown in ﬁgures 20A and ZOB,
have not been observed to have originated by springs in
the alluvial lowlands of the St. Francis Basin. Slumps
and lateral spreads along rivers (generally caused by
rapid hydraulic drawdown after ﬂoods) can generate
ﬁssures, but we have not observed sand ﬂowing to the
surface through the ﬁssures except for small amounts
associated with water oozing through the slide mass. In
addition, aseismic lateral spreads and slumps that we
have observed do not extend back from banks along large
rivers more than 50 m. Thus we believe that ﬁssures
associated with vented sand that is located more than 50
m from present or former banks of rivers can generally
be attributed to seismically induced slope instability.
Conﬁdence in this interpretation increases with increas-
ing width of the sand-ﬁlled ﬁssure and with increasing
distance from the banks. In order to conﬁdently attribute
an earthquake mechanism to a speciﬁc vented-sand ﬁs-
sure, however, the location of the river at the time of
ﬁssure formation must be established, an undertaking
that may be quite difﬁcult.

Fissures formed on terraces far from streams can be
attributed to earthquakes more conﬁdently. In inter-
preting an origin for any feature, it is still necessary to
assess ground—water conditions at the site and ascertain
that the ﬁssures could not have originated by some
process related to syndepositional deformations.

INTRUDED FEATURES

Intruded features are sand-ﬁlled cracks that do not
reach the surface, and therefore the features are visible

4 FIGURE 19. —“Eruptive vent” that cut stratiﬁed vented deposits of
a vented-sand volcano (from Haller and Crone, 1986). A, Number
102 is above a severely disrupted zone. Distances between vertical
edges of photograph is approximately 1 m. B, Number 61 is at
contact between vented sediments and the original ground surface.
Distance between vertical edges of photograph is approximately 1
m. C, Line drawing of “eruptive vent” in B. Heavy lines are
stratigraphic contacts; ﬁne lines show laminations within the
“eruptive vent,” dashed where discontinuous. Open circles are
aligned clay fragments.

28 EARTHQUAKE—INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

 

0
I
l
0

FIGURE 20. ——Aerial photographs showing vented sand, interpreted as
the product of liquefaction and flowage during the 1811—12 earth-
quakes. Note pattern of crazing in ﬁgures 20A and B, which is
especially indicative of ground breakage caused by earthquake—
induced liquefaction. A, Taken in 1941, showing more than 25
percent of ground surface covered by vented sand (small white

I
1 KILOMETER

spots and thin linear zones) near Marked Tree, Ark. B, Taken in 1959,
showing more than 25 percent of ground surface covered by vented
sand (small white spots and thin linear zones) near Portageville, Mo. C,
Taken in 1940, showing localized venting of sand near Reelfoot Lake,
Tenn. Regions with vented sand are outlined.

NEW MADRID SEISMIC ZONE 29

is
J. 1

 

O——O

FIGURE 20. —

only in exposures cut into the subsurface. A simple, very
common type of intrusion having an earthquake origin is
a near—vertical, wedge-shaped, sand-ﬁlled dike. At many
places in the region of vented sand, the clay—rich topstra-
tum is cut by vertical dikes that are spaced tens to
hundreds of meters apart. The dikes extend upward from

 

1 Kl LOMETER

Continued.

the substratum. Dike widths are several centimeters
near the substratum; widths commonly narrow upward
to an apex near the ground surface. The sand in the dike
is generally structureless and can contain clasts of top-
stratum carried upward; the larger clasts are often
oriented vertically.

30 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

 

1 KILOMETER

FIGURE 20. —Continued.

Combinations of dikes and sills are also common in the only slightly deformed clay beds, but locally thin clay
topstratum. The sills generally follow bedding planes or beds can be strongly warped. Figure 21A shows a typical
other horizontal planes of weakness in the topstratum. relationship of near-vertical dikes connected to a later-
Sills tended to form in thin beds of silt or sand beneath ally extensive sill. (The sill extends beyond the photo-

 

NEW MADRID SEISMIC ZONE 31

graph and is at least 25 m long.) The internal layering of
this sill is also typical of many of those we have studied.
Individual laminae are composed of small pieces of char-
coal or of much ﬁner grained sand and silt. (See ﬁgs. 213
and 210.) A single laminae of silt or very ﬁne sand can
have a length as long as 10 cm. Lamina made up of pieces
of charcoal can be much longer, and lengths of 300 cm are
not unusual. In some places, the sills are structureless,
having no traces of bedding, and contain many clay clasts
in a sand matrix; in other places, sills have graded
bedding with clay clasts concentrated along the base.
The absence of bedding or the presence of graded
bedding seems quite understandable; however, the pres-
ence of planar strata of ﬁne sand and silt within a much
coarser sand cannot be explained by us with any degree
of conﬁdence. (Possibly the planar stratiﬁcation develops
by repeated intrusions of ﬂuidized sand into an open
crack.) Most certainly, however, sills having stratiﬁed
bedding are common earthquake-related features that
resulted from intrusion of liqueﬁed substratum sands.
(Sills that have stratiﬁed and graded bedding originating
from earthquake-induced liquefaction of much deeper
source beds have also been observed in the South Caro-
lina study at sites HW and BLUF on ﬁg. 1.)

The clay clasts in the sills are generally angular,
suggesting a brittle or shattering mode of breakage of
the clay stratum from which the clasts were derived.
This brittle mode of breakage also suggests that when
the stratum was “shattered” the clay stratum was de-
watered signiﬁcantly from the initial depositional condi-
tion. Such shattering indicates forceful intrusion and
thus is not equivalent to the process causing formation of
syndepositional soft-sediment deformation features (dis—
cussed later). We have observed that the shattered clay
strata can have a consistency that is so soft that a person
can force his thumb a few inches into a stratum.

Sills in the topstratum at depths of about a meter or
less below the ground surface may be very wavy in
vertical section and can thin and thicken dramatically
within a meter, producing bulges at the surface. Sills can
be as much as 0.5 m thick near the surface. Such large
thicknesses seem less common at greater depths.

Earthquake-induced sand dikes that cut the upper
substratum sands and the lower topstratum are shown in
ﬁgure 22A, which is a schematic depiction of a vertical
ditch bank exposed about 15 km northwest of Marked
Tree. The exposure also illustrates many of the other
common forms of intrusions in the topstratum. In the

FIGURE 21. — Sand dikes and sills, interpreted as having originated b
by liquefaction and ﬂowage during the 1811—12 earthquakes. A,
Overview and line drawing of typical dikes and sills. Rectangle
shows area of ﬁgure 213. B, View showing detailed layering in sill.

0, Very close View showing detailed layering in sill.

 

  

O———D

Injected
sand

 

0 50 INCHES

H—~—H

0 100 CENTIMETERS

 

 

2 INCHES

l
|
5 CENTIMETERS

32 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

GROUND SURFACE

BED

 

——‘/‘

—\’/' i_.gl

—_l.
y d—
__4_

\

——+

l

 

t

 

Clay, medium gray,

 

 

blocky; altered by

 

44>
+ a
i -—l

+

soil processes

 

 

 

 

'l'l'l'l'l'l'l

 

 

’5-_clay clastslx'f '-

Sand, ﬁne- to medium-
grained, clean, orangish-
brown, massive

 

 

 

Clay, pale-blue, blocky,

 

_ —1——_ _ f —‘;_ — T —L'_ G micro-fractured, well
1 I laminated at base

 

 

 

 

 

 

 

 

 

 

 

 

Sand, medium- to coarse-grained,
clean, prominent foreset beds

 

E Sand, fine-grained, medium brown,
silty, contains planar crossbeds

 

————-————————TOPSTRATUM ———_————————————————————-»

Clay, highly plastic, grading up
to sand, ﬁne-grained, silty,
bluish-gray, massive

 

 

 

 

(

Clay, sandy, coarsening upward to
sand, ﬁne-grained, bluish-gray,
weakly laminated

Unconformity

 

 

 

Sand, very ﬁne grained, silty,
bluish-gray, well laminated at base

 

with carbonaceous partings

 

 

composed of charcoal grains

 

 

 

<-———-—-SUBSTRATUM-————+f——-—_———-———-

Médiﬁrﬁééﬁd.

 

 

Sand, medium-grained, clean
buff-brown with yellow stains,
crossbedded, contains wood
(dated at 12,000 years
before present)

0 3 FEET

F—#1

0 l METER

N0 VERTICAL EXAGGERATION

 

 

 

FIGURE 22.—Section showing Holocene sediments (topstratum) and
underlying Wisconsinan braided-stream sands (substratum) in ditch
about 15 km northwest of Marked Tree, Ark. Earthquake-induced
intrusions cut section at many places. A, Schematic diagram of
stratigraphic relationships and earthquake—induced liquefaction fea-

tures (numbered 1 through 5). Feature 1—Dike of medium sand that cut:
substratum and topstratum. Features 2 and 3—Intruded dikes and sills o
massive, clean, medium sand. Feature 4 —Dike of medium sand and largi
clasts from bed G. Feature 5—Dikes of medium sand, truncated b:
unconformity. Feature 6—Pseudonodules collapsed into bed D.

NEW MADRID SEISMIC ZONE 33

Bed C intruded sand
(feature 3)

W“; my 9%

vents
(feature 1)

small dike
(feature 5)

 

FIGURE 22.—Continued. B, Photo graph of beds A—D, showing sand oriented vertically in ﬁgure 22B. Knife is 12 cm long. E, Photograph of
intrusion (feature 3). Knife is 12 cm long. C, Photograph of vents part of feature 4, showing clay clasts in sand matrix intruded into bed H.
(feature 1) cutting substratum sand (bed A) and bed B. Knife is 12 cm Knife is 12 cm long. F, Photograph of dike of unknown origin (feature 5),
long. D, Photograph of plan view of bed B, showing intrusion cutting through bed C and truncated by bed D. Coin diameter is about
structures caused by feature 1; the plan view is in the area of the knife 1 cm.

34 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

upper part of the substratum (bed A, ﬁg. 22A), small
dikes branch out from a large dike (feature 1), cut
through the basal topstratum bed (bed B) at horizontal
intervals of 0.5 to 1 m, and extend upward about 0.5 to 1
m. However, a few dikes and sills have intruded at much
higher levels (features 2, 3, and 4). At places nearby (not
shown) dikes extend to the surface, and sand has been
vented to produce sand blows.

The dikes and sills shown on ﬁgure 22A as features 1
through 4 contain clean, structureless, medium-grained
sand. The edges of the intrusions are very sharp in
clay-rich beds (such as beds C, D, and G). Edges are
generally less distinct in beds of clean permeable sand.
Dikes commonly terminate in a permeable sand bed (bed
H); near the terminus there are often numerous clasts of
clay-rich material from the subjacent bed.

Features labeled 1 through 4 in ﬁgure 22 are probably
earthquake—induced because (1) they are widely distrib-
uted (common for tens of kilometers), (2) they contain
evidence of intrusion by large volumes of water-
saturated sediment (dikes and sills are commonly up to
15 cm wide), (3) there is evidence of forceful intrusion
(clean, medium-sized sand containing large clay clasts),
and (4) they are found in sites where artesian conditions
are unlikely. Because of the lack of weathering of sand in
the intrusions near the ground surface, we believe that
features 1 through 4 probably formed during the 1811—12
earthquake.

The dike shown as feature 5 has an uncertain origin.
Three small dikes that are truncated at the contact of
beds C and D were exposed in a 25—m section exposed
along the ditch but were not found in other nearby
exposures of beds C and D nor in the lowermost portion
of the topstratum at exposures further away (where beds
C and D cannot be traced). The dikes contain a large
amount of silty ﬁne sand and cannot be traced far down
into substratum sands. From this information, we inter-
pret that the dikes quite possibly originated due to
springs formed near the base of a streambank or, less
likely, as slump-related features.

Feature 6 (pseudonodules) is discussed in the next
section of the paper.

FEATURES OF UNKNOWN OR NONEARTHQUAKE ORIGIN

The principal features in this category include sand
boils, mima mounds, load structures, sand dunes, “deer
licks,” and sand ridges on point-bar deposits. Some of
these features can be distinguished from sand blows on
aerial photographs. Others may remain uncertain in their
origin even after ﬁeld studies.

 

SAND BOILS

During ﬂoods that occur every few years, sand boils
originate due to artesian ﬂowage beneath levees. Sand
boils are common in lowlands near levees along the large
rivers, particularly the Mississippi River (Kolb, 1976),
and are generally restricted to the areas within 0.5 to 1
km of the levees. (See ﬁg. 15.)

The sedimentary structures and morphologies of sand
boils are very similar to and may not be distinguishable
from those of (earthquake—induced) sand blows. For
example, both processes vent sand to the surface from
source beds at depth. The stratigraphic position of the
bed that provided the sand can be used to indicate origin
at some places. In a sand boil, the sand on the surface
commonly can be traced to a sand bed that immediately
underlies the cohesive (and thus relatively impermeable)
topstratum. In contrast, the source bed for a sand blow
may be much deeper. By itself, this criterion is of limited
value and generally is best used in a context that includes
tests to determine the earthquake-induced liquefaction
susceptibility of the various beds (discussed later).

The shape of the vent in plan view may help distin—
guish between sand boils and sand blows. Both sand boils
and sand blows can have circular vents; however, many
sand blows have linear vents that follow large ﬁssures.
Where large sand-ﬁlled ﬁssures widen downward and
extend for tens to hundreds of meters laterally, they are
probably vents associable with earthquake-induced liq-
uefaction (providing the possibility of nonseismic slump-
ing can be eliminated as the source of the ﬁssures).

Sand boils have vents whose diameters range from 3
cm to greater than 1 m. Numerous small sand-ﬁlled
ﬁssures or tubes can be connected to the main vent of a
sand boil; clay clasts up to 7 cm in diameter, derived from
sidewalls, have been observed within the vent. The
upper parts of the vents (as deep as 1 m or more below
the original ground surface) of sand boils may contain
stratiﬁed deposits of sand layered with other materials
that fell into an open hole. A nonearthquake origin can be
proven at some places because of the age of the materials
in the vents. For example, the presence of agricultural
produce such as rice or soybeans far down in the vent
virtually eliminates an association with historic earth-
quake origin in the New Madrid region, especially if the
vent is within 1 km of a levee. (Soybeans and rice have
been cultivated only within the past 50 years or so, much
postdating the last liquefaction-producing earthquake,
the 1895 Charleston, Mo., earthquake.)

MIMA MOUNDS

Mima mounds (sometimes called prairie mounds) are
domes of sand—rich material, usually less than 30 m in
diameter and 1 m high, that were not formed by venting;

NEW MADRID SEISMIC ZONE 35

in exceptional cases the domes are as much as 1.7 m high
but have diameters of only 10 to 12 m. Mirna mounds are
present in alluvial lowlands (especially in the northern
parts of the St. Francis and Western Lowlands Basins,
north of the town of New Madrid) and are also in upland
areas west of the Western Lowlands (ﬁg. 13). The origins
of mima mounds are not understood. In many upland
areas, the mima mounds are formed on nonliqueﬁable
deposits and therefore are of nonearthquake origin.
Mirna mounds on alluvial lowlands can be identiﬁed as
not resulting from earthquakes if excavation shows an
absence of vents to connect the mounds to source beds
below (Fuller, 1912, p. 80).

On aerial photographs, mima mounds can often be
distinguished from sand blows by criteria illustrated in
ﬁgure 23, which is an aerial photograph of mima mounds
taken in the northern part of the St. Francis Basin.
These criteria are regularity of spacing, alignment, and
size. In contrast, sand blows are typically irregularly
spaced, have widely differing sizes (heights and diame-
ters) at locations nearby, and are not aligned along such
precisely deﬁned curves. Comparison of sand blows
formed on point-bar deposits (ﬁg. 20C) with ﬁgure 23
illustrates the contrast.

Locally it can be difﬁcult to distinguish between sand
blows and mima mounds on aerial photographs, and
excavation is required to examine for feeder vents or
other evidence of origin. Mima mounds in the vicinity of
the St. Francis Basin characteristically have well-
developed soil horizons that obviously predate the
1811—12 earthquakes and in addition are slightly
cemented, making them difﬁcult to auger using hand
tools.

LOAD STRUCTURES

Many Holocene ﬂuvial deposits in the Central United
States have abundant syndepositional, soft-sediment
deformation features known as “load-ﬂow” or simply
“load” structures. Load structures are bulbous down-
ward intrusions of sandy or silty material into underlying
weaker, ﬁner grained muddy sediment. Two common
types of load structures occur in the topstratum of the St.
Francis Basin—pseudonodules and load-casted ripples.
Pseudonodules occur when overlying sandy or silty sed-
iments become detached and sink to become isolated
subspheroidal bodies in the underlying clayey bed.
Pseudonodules are usually found laterally adjacent to
other undetached load structures (Allen, 1984, p.
359—360). Pseudonodules were experimentally produced
in the laboratory by Kuenen (1958) by hammering a
container in which sand overlies water-saturated, very
soft clay. (See ﬁg. 24.)

Sims (1975) correlated subspheroidal pseudonodules
found in modern lake sediments, similar to those pro-

 

duced by Kuenen (and some possibly morphologically the
same as those produced by Kuenen), with known earth-
quake events. Sims argued that structures such as those
produced by Kuenen could be interpreted as earthquake-
induced if (1) they occur in a seismically active region, (2)
they are restricted to speciﬁc stratigraphic horizons, (3)
they are correlative over large areas within a sedimen-
tary basin, and (4) there is no detectable inﬂuence of
slope movement or failure. However, pseudonodules
similar to Kuenen’s can form without shaking by the
gravity-induced instability of denser sediment overlying
less dense sediment (Allen, 1984, p. 363). The degree of
soft-sediment deformation is controlled by the difference
in densities between the two adjacent layers and the
strength of the underlying layer. Sand deposited rapidly
over water-saturated, muddy or extremely soft clay is
ideal for deformation.

In the case of load-casted ripples, sandy or silty
intrusions form because of the unequal loading of migrat-
ing ripples of sand on a clayey mud substratum. Load-
casted ripples show progressively deformed radial inter-
nal lamination caused by the rotation of the ripple
cross-laminations as the ripples sink (Dzulynski and
Walton, 1965, p. 146—149). The asymmetry reﬂects cur-
rent direction, as illustrated in ﬁgure 25. The develop-
ment of load-casted ripples requires local deformation
synchronous with deposition of overlying sand, and
therefore load-casted ripples cannot likely be related to
earthquake events.

The distinction between possible earthquake-induced
load structures and those produced by rapid sedimenta-
tion is illustrated in a ditch exposure near Marked Tree,
Ark. The section in plate 1A is composed of bluish clays
interbedded with layers of sand, apparently related to
intermittent crevasse delta sand deposition in a water-
ﬁlled swale as illustrated in plate 13. Small, deltalike
sand bodies formed convex-upward sand lenses contain-
ing climbing ripple cross-laminations. The lenses are
coarser upward, ﬁne toward the edges, and are laterally
adjacent to layers of pseudonodules (25—50 cm long and
5—25 cm Wide) made up of similar grain sizes. The base of
each sand lens has load structures, including load-casted
ripples (pl. 13 ). The load structures make up a greater
percentage of each sand lens as it thins toward the edge
(see north-south View of pl. 18); downstream, each sand
lens apparently grades into the pseudonodules. (See
east-west view of pl. 1A.) The cross-lamination in the
sand lenses has a gradational contact to the deformed
ripple cross-laminae in the underlying load structures,
whereas ripple cross-laminae above and laterally adja-
cent to the load structures are undeformed. (See north-
south View of pl. 13.) The pseudonodules are internally
ripple cross—laminated, and some have load-casted
ripples at the base. Isolated pseudonodules with the

36 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

 

0 1 MILE
I

I,

0

1 KILOMETER

FIGURE 23. —Aerial photograph of mima mounds in northern part of St. Francis Basin. Mima mounds often appear as trails of light dots and are
commonly formed on slightly elevated ridges of point-bar deposits.

internal fabric of load-casted ripples have an asymmetry

the sand lenses. The clay surrounding the pseudonodules
reﬂecting a current direction about the same as that for

has streaks of highly contorted silt, indicating ﬂowage

NEW MADRID SEISMIC ZONE 37

 

 

Water or an

 

 

 

 

 

 

 

 

 

 

 

 

0 5 INCHES

a

0 10 CENTIMETERS

FIGURE 24.—Pseudonodules formed by shaking, from Keunen (1958).
A layer of sand (unit 1) overlies very soft clay (unit 2) in A. With
shaking, pseudonodules developed (B and C) and became completely
enclosed in clay (D and E). Note the destruction of layering from C
to D. This deformation is due to sinking of the entire sand layer into
the very soft clay and also to localized sinking caused by uneven
loading.

around the pseudonodules. The lower portion of the
sequence of sediments in the trench is brecciated and
faulted; the breccia, including pseudonodules, is sur-

 

 

 

FIGURE 25.—Development of load-casted ripples, caused by ripple
crests sinking into soft mud. Note the progressive tilt of the internal
cross-lamination, which causes the downﬂow portion to be more
steeply inclined than the upﬂow portion (modiﬁed from Dzulynski
and Kotlarczyk, 1962). Load—casted ripples can also have an opposite
sense of rotation (upﬂow more steeply inclined than downflow), as
evidenced in ﬁeld observations in the St. Francis Basin.

rounded by a coarse-grained sand that has been injected
from beneath. These larger scale deformation features
are truncated by the uppermost sand lens (sand unit X),
which is also laterally equivalent to pseudonodules.

The load structures in plate 1A are interpreted as
synsedimentary in origin and not as earthquake-induced
on the basis of the following criteria: (1) The pseudonod-
ules are gradational into the sand lenses, including
load-casted ripples, with the same sense of ﬂow direc-
tion, (2) there are several layers of pseudonodules later-
ally equivalent to largely undeformed sand lenses, and
(3) the recurring conditions of sedimentation (rapid pro-
gradational deposition of sand over soft clay) were con-

38 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

ducive to this type of soft-sediment deformation.
Although load structures such as these may be formed by
earthquakes, their repeated occurrence in this sequence
(as also shown by Allen, 1984, ﬁgs. 9, 10; or in Coleman
and Prior, 1980, ﬁgs. 15, 21, 22, and 32) and their
apparent localized occurrence argue for a depositional
origin. (The injection features, brecciation of clay, and
faulting in this exposure that crosscut the syndeposi-
tional features are probably of earthquake origin.)

Pseudonodules formed by earthquake shaking would
be similar in shape to those formed by depositional
processes, because the mechanism of sinking is similar.
The environmental conditions conducive to pseudonodule
development are also ideal for the formation of convo-
luted bedding, ﬂuidization dikes, and slumping. Any
depositional environment in which coarse-grained sedi-
ment is introduced rapidly in otherwise quiet, standing-
water conditions, such as ﬂuvial overbanks, lacustrine or
marine deltas, and submarine fans, may include these
features as part of its regular sedimentary record. We
suggest that the difference between syndepositional
deformation and earthquake-induced deformation would
be in the relationship of the pseudonodules to the sur-
rounding sediments. Intense earthquake shaking might
cause collapse of an entire sand layer into clay, rather
than only the thin edge, and a small portion of the base of
a sand lens; in addition, the structures should be wide-
spread in deposits having the same age. A widespread
distribution of load structures formed at the same time is
difﬁcult to prove in most ﬂuvial depositional environ-
ments.

OTHER FEATURES

Field examination is often required to distinguish
between sand blows and small aeolian sand dunes, par-
ticularly on high terraces having very sandy surface
soils. Another feature requiring ﬁeld examination is a
“deer lick,” which is a very light—colored area, up to 15 m
in diameter, caused by a high sodium concentration in
poorly drained silty soil on level ground (Brown and
others, 1973). They are relatively common in western
Tennessee.

Aerial photographs often show arcuate ridges of sand
on modern and Holocene point-bar deposits, as illus-
trated in ﬁgure 200. These sand ridges form where the
ﬁne-grained topstratum is especially thin, and thus more
quickly eroded, or was never present over the point-bar
deposits. Field excavations are often required to deter-
mine if the sands have been vented versus exposed by
erosion.

LOCAL GEOLOGIC CONTROLS ON PRODUCTION OF
VENTED SAND

Some of the geologic controls on production of vented
sand during the 1811-12 earthquakes were ﬁrst noted by

 

Saucier (1977) and were later described in much more
detail by Obermeier (1988, 1989). Principal controls are
topstratum thickness and lithology, substratum sand
grain size (and, indirectly, permeability), susceptibility
to liquefaction of source sand (as related to age of the
sediment and measured by the Standard Penetration
Test (SPT)), and earthquake accelerations. The inﬂuence
of these parameters was determined by hundreds of test
borings, which were used to formulate a general notion of
the geologic setting where sand was vented and to
back-calculate 1811—12 earthquake accelerations. Along
the border where sand blows developed (ﬁg. 14), the
Simpliﬁed Procedure of Seed and Idriss (1982) was used
to estimate peak earthquake accelerations. Accelerations
along the border probably did not exceed 0.25 g (Ober-
meier, 1988, 1989).

Substratum sands in the St. Francis Basin are typi-
cally at least moderately dense (engineering sense). N1
values (SPT blow counts adjusted to an overburden
stress of 1 ton/ftz) in the depth range most likely to
liquefy (5—15 m) are rarely less than 10, and median
values are generally about 25.

Other basic data that should be kept in mind during
this discussion of geologic controls for liquefaction fea-
tures of the 1811—12 earthquakes are that the substratum
sands generally have a narrow range of thickness
(between 25 to 40 m); the epicenters of the three major
shocks in 1811—12 were almost certainly somewhere
within the areas where large volumes of sand were
vented; and the MS values of the three major shocks were
probably about 8.5 to 8.7.

TOPSTRATUM THICKNESS

In general, where nonliqueﬁable topstratum thickness
exceeded a critical threshold, there was increasing
resistance to development of sand blows and venting of
large volumes of sand (Obermeier, 1989). A topstratum
thicker than 10 m presented a major barrier to more than
scattered venting, even in close proximity to the epicen-
ter. No sand blows formed where the topstratum was
thicker than about 15 m. Within 20 to 25 km of the
epicentral regions, a 5- to 6-m-thick topstratum was
insufﬁcient to prevent severe, extensive venting. At
more distant localities, near the border of sand blows, a
3- to 4-m thickness of topstratum noticeably reduced
sand-blow development. Where a thick topstratum
restricted venting to the surface, intrusions within the
lower part of the topstratum were still plentiful.

TOPSTRATUM LITHOLOGY

Even though the great majority of our ﬁeld observa-
tions were at sites where the topstratum was dominantly
clay, we observed enough areas where the surﬁcial

OVERVIEW OF STUDIES 39

deposits were sand rich to make some general observa-
tions about the formation of craters and vented-sand
volcanoes. The largest craters occurred Where the top-
stratum was most sand rich at the ground surface. The
stratigraphy in the craters is generally not well deﬁned
in comparison to craters in South Carolina. The largest
clay clasts are at the base, and clast size decreases
upward in a structureless sand matrix. The transition
from graded zone to the overlying laminated bedding is
difﬁcult to ﬁnd in some craters; the laminated bedding is
generally deﬁned only by thin, wispy layers of slightly
different sizes of sand and silt. There is no evidence of
organic material accumulating in an open hole, as in
South Carolina.

We did not observe V— or U-shaped fractures in
regions of craters, as in South Carolina. Partly for that
reason, and partly because of the difﬁculty of making
moderately dense, well—graded sand ﬂow extensively
after being liqueﬁed, we suspect that the mechanism for
formation of craters in the New Madrid region differs
from the mechanism in South Carolina, where the erup—
tive craterlets were abundant. In the New Madrid
region, the craters may have formed primarily because of
abrasion and widening of the sidewalls as ﬂuidized sand
and water vented for a prolonged time. Similarly, large
craters did not tend to form extensively in the clay-rich
topstratum in the New Madrid region because of the
difﬁculty of abrading the sidewalls.

SUBSTRATUM GRAIN SIZE

Medium-grained and coarse-grained, clean, well-
graded sands greatly predominate in the substratum, in
the depth ranges that were liqueﬁed in 1811—12. Other
factors being equal, the most extensive development of
sand blows corresponded to areas underlain by medium-
grained sand. Much smaller and more scattered sand
blows developed over coarse-grained substratum sands
than over medium-grained sands. Coarse-grained sub-
stratum sands have much higher permeabilities than the
medium-grained sands; possibly the higher permeability
permitted pore pressures of the liqueﬁed sands t0 dissi-
pate rapidly enough to curtail sand—blow development.

RESISTANCE OF SOURCE SAND TO LIQUEFACTION

Only small regional variations in the relative density of
the late Quaternary substratum sands are found in the
St. Francis Basin; these sands are typically at least
moderately dense and thus moderately difﬁcult to liq-
uefy. However, historical accounts (Street and Nuttli,
1984) and data by Fuller (1912) show that 1811—12 sand
blows were common on modern ﬂood plains far from the
St. Francis Basin. The modern ﬂood plains are underlain
by younger and generally looser and more liqueﬁable

 

sand beds than those in the St. Francis Basin (Obermeier
and Wingard, 1985), thus explaining the development of
sand blows much further from the epicenters.

Liquefaction features in the New Madrid seismic zone
have not been observed to be induced by numerous
historic earthquakes having mb of 5.3 to 5.5, or by a
single earthquake having mb of 6.0; in 1895, many sand
blows (generally small) were produced by an mb 6.2
earthquake Whose epicenter was near Charleston, Mo.
(Obermeier, 1988). These sand blows developed in allu-
vium that is slightly less susceptible to liquefaction than
that at most other places in the New Madrid seismic
zone, excluding the modern (<500 yr) point—bar and
sand-bar deposits of the rivers in the region. Thus a
reasonable earthquake threshold is approximately mb 6.0
for small liquefaction features in braided stream and
meander belt deposits, excluding very young alluvial
deposits.

OVERVIEW OF STUDIES

Geologic criteria have been developed and geologic
controls have been evaluated for earthquake-induced,
level-ground liquefaction features in alluvium of coastal
South Carolina and the New Madrid seismic zone. Many
different types of earthquake features occur in these
areas, in addition to features of unknown or nonseismic
origin that might be interpreted as having an
earthquake-induced origin.

GEOLOGIC CRITERIA

Assigning an earthquake origin to possible sand blows
generally requires that four criteria be satisﬁed:

1. The features must have sedimentary characteristics
that are consistent with earthquake-induced liquefaction
origin; that is, there is evidence of an upward-directed,
strong hydraulic force that was suddenly applied and was
of short duration.

2. The features must have sedimentary characteristics
that are consistent With historically documented obser-
vations of earthquake-induced liquefaction processes.

3. The features must occur in ground-water settings
where suddenly applied, strong hydraulic forces of short
duration could not be reasonably expected except from
earthquake-induced liquefaction. In particular, such set—
tings are extremely unlikely sites for artesian springs.

4. Similar features must occur at multiple locations,
preferably at least within a few kilometers of one
another, having similar geologic and ground-water set-
tings. Where evidence of age is present, it should sup-
port the interpretation that the features formed in one or
more discrete, short episodes that individually affected a

40 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

large area and that the episodes were separated by long
time periods during which no such features formed.

As fewer of these elements are satisﬁed, conﬁdence in
interpretation of an earthquake origin generally dimin-
ishes. Considerable reliance has been placed on the
second element for interpreting an earthquake origin in
coastal South Carolina and the New Madrid seismic zone.
If, however, all the other criteria are unequivocally
satisﬁed outside these geographic areas, an earthquake
origin can still be ascertained in areas that have not had
historic earthquakes.

Criteria based on engineering analysis may be useful
for diagnosis of the origin of features. To illustrate, if the
source sand bed can be identiﬁed at a site of possible
liquefaction, and the sand bed proves to be nonliqueﬁable
on the basis of an engineering analysis, then an
earthquake—induced origin can be eliminated. As another
example, if mineralogical analysis associates vented sed—
iment with a deep stratum of loose sand rather than with
a shallower sand bed of much higher compactness, and an
engineering analysis of nonearthquake-related ground-
water ﬂow forces shows that the shallower, more dense
sand should have been transported up in preference to
the deeper, loose sand, then an earthquake origin may be
ascertained.

TYPES OF EARTHQUAKE FEATURES

In both coastal South Carolina and the New Madrid
seismic zone, earthquake-induced liquefaction features
originated in sand deposits that are relatively thick (3—50
m) and contain no or few intercalated silt- or clay-rich
strata. At the ground surface there is a cap that is much
less permeable than either the subjacent or deeper
source sand beds. Properties of the cap have a major
effect on the surface expression of the sand blows. In
coastal South Carolina, Where the cap is generally a
l-m-thick soil that is weakly cemented with humate, the
sand blows are expressed as craters surrounded by thin
sand sheets; in the New Madrid seismic zone, the cap is
generally a clay—rich deposit about 2 to 10 m thick, and
sand blows are expressed as sand volcanoes.

Both vented and intruded features can be related to
earthquake-induced liquefaction. Features characterized
by sand vented to the surface include more or less
circular, individual sand blows and also long, irregular
sand-ﬁlled ﬁssures, hundreds of meters long. Where a
clay-rich cap exceeds a critical thickness, sand is not
vented to the surface; instead, earthquake-induced liq—
uefaction is represented only by sand intrusions.
Intruded dikes usually have a massive internal structure,
whereas sills can contain both graded and laminar bed-
ding in coastal South Carolina and the New Madrid
seismic zone.

 

Lateral spreads produced by the 1811—12 earthquakes
were common. (See Fuller, 1912, p. 48, for a schematic
drawing.) We made no effort to locate lateral spreads in
the New Madrid seismic zone, but evidence of lateral
spreads is described in coastal South Carolina.

GEOLOGIC CONTROLS

The thickness of clay—bearing or impermeable surﬁcial
deposits can be an important control on development of
surface venting. In the St. Francis Basin, our data show
relations between the distance from the epicenter, thick-
ness of topstratum, and development of surface venting
produced by the 1811—12 New Madrid earthquakes; near
the epicenters, a topstratum thickness more than about
12 m prevented venting, whereas near the farthest limits
of sand-blow development, a topstratum more than 4 to
5 m thick greatly restricted development. In coastal
South Carolina, a cap exceeding 3 to 5 m thick generally
prevented development of sand blows for both 1886 and
pre-1886 earthquakes. These differences in critical thick—
nesses between coastal South Carolina and the New
Madrid seismic zone most likely reﬂect the inﬂuence of
factors such as thickness of source sand bed, susceptibil-
ity of source sand bed to liquefaction, and earthquake
magnitude. Similar relations between critical thickness
of surﬁcial cap and development of surface manifesta-
tions of liquefaction have been noted by Ishihara (1985).

The rate of expulsion of water to the surface while the
sand blow was forming possibly controls whether sand
blows formed as craters or as sand mounds, but this
speculation has not been tested. Certainly, in coastal
South Carolina, the sands are generally much looser than
sands in the New Madrid seismic zone, and this looser
state may have helped create a very thick, water-rich
zone during shaking, which was probably the primary
source zone during initial, very rapid explusion. Sands in
the New Madrid seismic zone characteristically are much
more permeable than the marine sands in coastal South
Carolina, yet craters are not common features in the
New Madrid seismic zone. Thus, the craters cannot be
necessarily associated with higher permeability alone of
source sands.

Vertical cracks formed by desiccation or other pro-
cesses may predispose clay-rich sediments to form
vented-sand volcanoes rather than craters. Vertical
cracks are common in both coastal South Carolina and
the New Madrid seismic zone.

Source beds for sand blows in the New Madrid seismic
zone are predominantly medium- and coarse—grained
sands; in the 1811—12 earthquakes, the coarse-grained
and thus highly permeable source sands produced fewer
and smaller sand blows than medium- to ﬁne—grained
source sands. Insufﬁcient data are available to form

SUGGESTIONS FOR FUTURE RESEARCH 41

conclusions about inﬂuence of grain size on sand-blow
size in coastal South Carolina, because almost all sands
are ﬁne grained.

Geologic age is an important control on liquefaction
susceptibility in both the New Madrid seismic zone and
coastal South Carolina. Youngest sediments are gener—
ally most susceptible to liquefaction, for a given mode of
deposition.

SUGGESTIONS FOR FUTURE RESEARCH

In recent years it has become increasingly clear that
detection of liquefaction features is valuable as a tool for
understanding earthquake activity. Liquefaction fea-
tures can be used to identify strong earthquake shaking
and to date the strong shaking events. We have
described our interpretations of earthquake- and
nonearthquake-related features from ﬁeld studies in the
New Madrid seismic zone and in coastal South Carolina.
Still, more work is needed to better identify potential
earthquake hazards in these geographic areas.

In the study of New Madrid 1811—12 liquefaction
features, much was learned that is relevant to any search
for paleoseismic liquefaction in moderately thick ﬂuvial
sediments of the Central United States. Aerial photo-
graphs are useful in a search for features formed in
1811—12, and earthquake deposits predating 1811—12 laid
down on terraces having clay-rich surface soils may also
be visible on such photographs. However, deposits laid
down on terraces having sandy surﬁcial soils have prob—
ably been so intensely reworked by wind that the depos-
its would be difﬁcult, if not impossible, to identify. Aerial
photographs should range from the earliest possible date
(generally the late 1930’s and early 1940’s) to about the
mid-1950’s. The earlier photographs may show many
features destroyed by modern farming, whereas more
modern photographs generally have better overall clar-
ity.

Field checking is generally required to determine the
origin of features that on aerial photographs are sus-
pected to have an earthquake origin. In particular, in
point-bar deposits, ﬁeld checking is often required to
verify the origin of sand that appears to have been
vented along slightly elevated, arcuate ridges. On ter-
races having sandy surface soils, sand blows are often
indistinguishable from sand dunes, and distinctive signs
of previous earthquakes may be restricted to long ﬁs-
sures that formed near and parallel to scarps along
terrace sublevels.

Searches in drainage ditches and sand pits are proba—
bly the only means for identifying liquefaction features
that much predate the 1811—12 earthquakes. The best
areas to search can be predicted by evaluating the

 

inﬂuence of geologic controls such as sediment age,
topstratum thickness, grain size of source beds, and
proximity to suspected epicenters.

Any search for paleoseismic features in coastal South
Carolina and in the Coastal Plain of the Eastern United
States should also make extensive use of exposures in
ditches; aerial photographs, for example, are useless for
locating 1886 and pre-1886 sand blows in South Carolina.
Determination of an earthquake or nonearthquake origin
for disrupted ground features in the Coastal Plain gen-
erally requires a team of people having many skills. In
addition to having an understanding of sedimentation
processes, soil science, and engineering mechanics, there
may also be a need for geochemistry studies because
weathering in these humid environments produces
bizarre features not previously described in the litera-
ture. We have described some weathered and disrupted
ground features that could cause confusion, but our
catalog is far from complete.

Research is also needed to provide criteria for distin-
guishing between features having an earthquake-
induced liquefaction origin and features having a short—
term spring origin. Locations where springs can be easily
eliminated as a cause, such as along the crests of beach
ridges of coastal South Carolina, are relatively sparse in
the Eastern United States. Fluvial terraces in lowlands
contain the only potentially liqueﬁable deposits in most
places of interest to paleoseismicity. Possible indicators
of origin include the ground failure mode; the nature of
the ﬁlling in vents, dikes, and sills; geologic characteris-
tics of the source beds; and engineering characteristics of
the source beds. The relevance of each indicator is
discussed below.

Ground failure mode.—Ground failure mode includes
features such as lateral spreads, single long fractures,
and shattered ground at the ground surface. All of these
features generally indicate an earthquake origin.
Earthquake-induced liquefaction does not always pro-
duce these types of features, however, and in many cases
they are not easy to recognize in ditch exposures even
where they are known to have been relatively common,
such as the meizoseismal region of the 1886 Charleston
earthquake. Use of techniques that can rapidly develop a
three-dimensional View, such as ground-penetrating
radar, may prove to be of great value.

Nature of filling—Vents associated with an
earthquake-induced liquefaction origin generally appear
to be ﬁlled with a structureless mixture of sand and silt
grains and clasts derived from sidewalls and beds at
depth. This structureless mixture apparently represents
transport through the vent as a slurry. In some places,
particularly near the top of the vent, the ﬁner grained
material has been winnowed out, thus indicating trans-
port on a grain-by-grain basis. For either the slurry or

42 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

the grain-by-grain mode of transport, it should be possi-
ble to place limits on the hydraulic forces that caused the
transport and, by means of a ﬂow-potential diagram,
calculate whether nonearthquake ﬂow forces could pos-
sibly have been the driving mechanism.

Numerous observations in the New Madrid seismic
zone show a wide variation in the nature of sill ﬁllings
caused by earthquake-induced liquefaction. The ﬁllings
range from layered bedding to graded bedding to a
structureless mixture of sand, silt, and clay clasts; for
paleoseismicity interpretations, we suspect that the
structureless mixture is associated only with the very
forceful injection caused by earthquake—induced liquefac-
tion, whereas layered and graded bedding may be asso—
ciated with either a spring or earthquake-induced lique-
faction origin.

In the New Madrid seismic zone, some of the larger
sills within 2 m of the ground surface have domed the
overlying beds as much as 0.3 m vertically over a
horizontal distance of 5 m; such large doming may be
unique to earthquake-induced liquefaction. The nature of
the ﬁllings in dikes may also be related to an earthquake-
induced liquefaction or spring origin. Relevance of the
ﬁllings in vents, dikes, and sills can be veriﬁed only by
research in the ﬁeld.

Geologic characteristics of source beds—Intuition
suggests that earthquake-induced liquefaction should be
much more effective than springs as a means to destroy
original bedding over a large area. This suggestion has
not been examined, however, and may be difﬁcult to
evaluate in practice. Earthquake-induced liquefaction
would also probably produce many more widespread sills
and dikes than springs.

Engineering characteristics of source beds—The
source beds containing the material vented by springs
should be located selectively with respect to the ground-
water setting and ﬂow forces. For springs, the source
beds must be connected to the source of ﬂowage in such
a manner that ﬂowage goes toward the vent, and the
ﬂowage forces must be large enough to carry material
from the source bed. This source bed would be, in
general, the uppermost stratum of wide lateral extent
with respect to hydraulic connection and would also be
the uppermost ﬁne-grained, noncohesive sand stratum.
The source bed with regard to earthquake-induced liq-
uefaction commonly lies much below (several meters) the
uppermost sands of wide lateral hydraulic connection,
according to limited observations in deep ditches in the
New Madrid seismic zone. Thus, permeability relations
may be quite useful for determining the causative mech-
anism responsible for venting materials.

Standard Penetration Test data that show relative
ability to liquefy during earthquake shaking should also
prove useful for determining origin at some places.

 

Geologic investigation may also be required to determine
origin in areas where Standard Penetration Test data
indicate possible source beds that are thin (<1 m) and lie
immediately beneath an impermeable cap, because
(injected) sills beneath impermeable caps have been
observed by us to approach a meter in thickness.

RELEVANCE OF LIQUEFACTION FEATURES

Recognition of the various features described in this
paper, and identiﬁcation of the most probable origin for
each, provides a set of important tools for understanding
the paleoseismicity in areas where faults are not obvious
at the surface and where historic seismic activity is
infrequent. Even where faults are available for study and
offsets can be documented, there may be doubt that the
offsets were associated with earthquakes; the presence
of liquefaction-induced features is a means of verifying
strong shaking.

The criteria we describe in this paper can be applied to
many worldwide geologic settings. We caution, however,
that veriﬁcation of paleoseismic events becomes gener-
ally more difﬁcult with increasing age of the event. The
distinction between earthquake and nonearthquake liq-
uefaction features is often impossible without knowledge
of ground-water conditions. Application of the criteria to
pre-Quaternary deposits may be extremely difﬁcult, if
not impossible, at many places.

REFERENCES CITED

Allen, J.R.L., 1984, Sedimentary structures—Their character and
physical basis: Developments in Sedimentology, Elsevier, Amster-
dam, New York, v. 30B, 663 p.

Bollinger, GA., 1977, Reinterpretation of the intensity data for the
1886 Charleston, South Carolina, earthquake, in Rankin, D.W.,
ed., Studies related to the Charleston, South Carolina, earthquake
of 1886—A preliminary report: US. Geological Survey Professional
Paper 1028, p. 17—32.

Braile, LW., Hinze, W.J., Sexton, J.L., Keller, G.R., and Lidiak,
E.G., 1984, Tectonic development of the New Madrid seismic zone,
in Gori, P.L., and Hays, W.W., eds, Proceedings of the sympo—
sium on the New Madrid seismic zone: US. Geological Survey
Open-File Report 84—770, p. 204—233.

Brown, W.T., Jackson, W.C., Keathley, G.L., and Moore, G.L., 1973,
Soil survey of Obion County, Tennessee: US. Department of
Agriculture, Soil Conservation Service, 58 p.

Carter, DR, and Seed, HE, 1989, Liquefaction potential of sand
deposits under low levels of excitation: Earthquake Engineering
Research Center, University of California at Berkeley, Report
No. 88/11, 282 p.

Coleman, J .M., and Prior, DB, 1980, Deltaic sand bodies: A 1980
short course, education course note series no. 15, American
Association of Petroleum Geologists, 171 p.

Cox, J .H.M., 1984, Paleoseismicity studies in South Carolina: Unpub-
lished M.S. thesis, University of South Carolina, Columbia, 75 p.

REFERENCES CITED 43

Crone, A.J., McKeown, F.A., Harding, S.T., Hamilton, R.M., Russ,
DR, and Zoback, MD, 1985, Structure of the New Madrid
seismic source zone in southeastern Missouri and northeastern
Arkansas: Geology, v. 13, no. 8, p. 547—550.

Cullen, C.J., 1985, Engineering tests on sands associated with Charles-
ton, South Carolina, seismic events: Unpublished M.S. thesis,
Virginia Polytechnic Institute and State University, Blacksburg,
61 p.

Daniels, R.B., Nettleton, W.D., McCracken, R.J., and Gamble, E.E.,
1966, Morphology of soils with fragipans in parts of Wilson County,
North Carolina: Soil Science Society of America Proceedings, v.
30, p. 376—380.

Dewey, J .W., 1985, A review of recent research on the seismotectonics
of the Southeastern seaboard and an evaluation of hypotheses on
the source of the 1886 Charleston, South Carolina, earthquake:
U.S. Nuclear Regulatory Commission, Washington, D.C.,
NUREG/CR—4339, 44 p.

Dickenson, S.E., Clough, G.W., and Martin, J .R., 1988, Evaluation of
the engineering properties of sand deposits associated with lique-
faction sites in the Charleston, South Carolina, area: Blacksburg,
Va., unpublished M.S. thesis, Virginia Polytechnic Institute and
State University, 115 p.

Dickey, D.D., 1985, Aerial reconnaissance study of areas that may
have sand blows caused by the New Madrid earthquakes, in
Hopper, M.G., ed., Estimation of earthquake effects associated
with large earthquakes in the New Madrid seismic zone: U.S.
Geological Survey Open-File Report 85—457, p. 142—149.

Dutton, C.E., 1889, The Charleston earthquake of August 31, 1886:
U.S. Geological Survey Ninth Annual Report 1887—88, p. 203—528.

Dzulynski, S., and Kotlarczyk, J ., 1962, On load-casted ripples:
Annales de la Societe Geologique de Pologne, v. 32, p. 148—159.

Dzulynski, S., and Smith, A.J., 1963, Convolute lamination, its origin,
preservation, and directional signiﬁcance: Journal of Sedimentary
Petrology, v. 33, p. 616—627.

Dzulynski, S., and Walton, E.K., 1965, Sedimentary features of ﬂysch
and greywackes: Developments in Sedimentology, v. 7, Elsevier,
Amsterdam, New York, 274 p.

Fuller, M.L., 1912, The New Madrid earthquake: U.S. Geological
Survey Bulletin 494, 119 p.

Gamble, E.E., 1965, Origin and morphogenetic relations of sandy
surﬁcial horizons of upper Coastal Plain soils of North Carolina:
Unpublished Ph.D. dissertation, North Carolina State University,
Raleigh, 254 p.

Gelinas, R.L., 1986, Mineral alterations as a guide to the age of
sediments vented by prehistoric earthquakes in the vicinity of
Charleston, South Carolina: Unpublished M.S. thesis, University
of North Carolina, Chapel Hill, 304 p.

Gohn, G.S., Weems, R.E., Obermeier, SR, and Gelinas, R.L., 1984,
Field studies of earthquake-induced liquefaction-ﬂowage features
in the Charleston, South Carolina, area: preliminary report: U.S.
Geological Survey Open-File Report 84-670, 26 p.

Haller, KM, and Crone, A.J., 1986, Log of an exploratory trench in
the New Madrid seismic zone near Blytheville, Arkansas: U.S.
Geological Survey Miscellaneous Field Studies Map MF—1858.

Hamilton, RM, and Zoback, MD, 1982, Tectonic features of the New
Madrid seismic zone from seismic reﬂection proﬁles, in McKeown,
F.A., and Pakiser, L.C., eds., Investigations of the New Madrid,
Missouri, earthquake region: U.S. Geological Survey Professional
Paper 1236, p. 55-82.

Hays, W.W., and Gori, P.L., eds., 1983, Proceedings of Conference
XX, A workshop on “The 1886 Charleston, South Carolina, earth-
quake and its implications for today”: U.S. Geological Survey
Open—File Report 83—843, 502 p.

 

Heyl, A.V., and McKeown, F.A., 1978, Preliminary seismotectonic
map of the central Mississippi Valley and environs: U.S. Geological
Survey Miscellaneous Field Studies Map MF—1011, scale
1:500,000.

Hopper, M.G., Algermissen, ST, and Dobrovolny, E.E., 1983, Esti-
mation of earthquake effects associated with a great earthquake in
the New Madrid seismic zone: U.S. Geological Survey Open-File
Report 83—179, 94 p.

Ishihara, K., 1985, Stability of natural deposits during earthquakes:
Proceedings of the 11th International Conference on Soil Mechan-
ics and Foundation Engineering, A. A. Balkema, The Netherlands,
p. 321—376.

Jibson, R.W., 1985, Landslides caused by the 1811—12 New Madrid
earthquakes: Unpublished Ph.D. thesis, Stanford University,
Stanford, Calif, 232 p.

Kolb, C.R., 1976, Geologic control of sand boils along Mississippi River
levees, in Coates, D.R., ed., Geomorphology and Engineering:
Stroudsburg, Pa., Dowden, Hutchinson & Sons, p. 99—113.

Krinitzsky, EL, and Wire, J .C., 1964, Groundwater in alluvium of the
lower Mississippi valley (upper and central areas): U.S. Army
Corps of Engineers, Waterways Experiment Station, Technical
Report no. 3—658, v. I and II.

Kuenen, RH, 1958, Experiments in geology: Transactions of the
Geological Society of Glasgow, v. 23, p. 1—28.

Lowe, DR, 1975, Water escape structures in coarse—grained sedi-
ments: Sedimentology, v. 22, no. 2, p. 157—204.

1976, Subaqueous liqueﬁed and ﬂuidized sediment ﬂows and
their deposits: Sedimentology, v. 23, no. 3, p. 285—308.

McCartan, L., Lemon, E.M., Jr., and Weems, RE, 1984, Geologic
map of the area between Charleston and Orangeburg, South
Carolina: U.S. Geological Survey Miscellaneous Investigations
Map I—1472, scale 1:250,000.

McDermott, J.F., 1949, Old Cahokia—A narrative and documents
illustrating the first century of its history: The Saint Louis
Historical Documents Foundation, Saint Louis, Mo.

McKeown, F.A., and Pakiser, L.C., eds., 1982, Investigations of the
New Madrid, Missouri, earthquake region: U.S. Geological Survey
Professional Paper 1236, 201 p.

Mosaic, 1976, Quakes in search of a theory: National Science Founda-
tion, Washington, D.C., v. 7, no. 4, p. 2—11.

Nettleton, W.D., McCracken, R.J., and Daniels, R.B., 1968a, Two
North Carolina Coastal Plain catenas—I. Morphology and fragipan
development: Soil Science Society of America Proceedings, v. 32,
p. 577—582.

Nettleton, W.D., Daniels, R.B., and McCracken, R.J., 1968b, Two
North Carolina Coastal Plain catenas—II. Micromorphology, com-
position, and fragipan genesis: Soil Science Society of America
Proceedings, v. 32, p. 582—587.

Neumann—Mahlkau, P., 1976, Recent sand volcanoes in the sand of a
dike under construction: Sedimentology, v. 23, p. 421—425.

Nuttli, O.W., 1973, The Mississippi valley earthquakes of 1811 and
1812: intensities, ground motion and magnitudes: Bulletin of the
Seismological Society of America, v. 63, p. 227—248.

1979, Seismicity of the Central United States, in Hatheway,

A.W., and McClure, C.R., Jr., eds., Geology in the siting of

nuclear power plants: The Geological Society of America, Reviews

in Engineering Geology, v. 4, p. 67~93.

1982, Damaging earthquakes of the central Mississippi valley,

in McKeown, F.A., and Pakiser, L.C., eds., Investigations of the

New Madrid, Missouri, earthquake region: U.S. Geological Survey

Professional Paper 1236, p. 15—20.

1983a, 1886 Charleston, South Carolina, earthquake revisted,

in Hays, W.W., and Gori, P.L., Proceedings of Conference XX,

 

 

 

 

44 EARTHQUAKE-INDUCED LIQUEFACTION FEATURES IN SOUTH CAROLINA AND THE NEW MADRID SEISMIC ZONE

A workshop on “The 1886 Charleston, South Carolina, earthquake
and its implications for today”: US. Geological Survey Open-File
Report 83—843, p. 44—50.

1983b, Average seismic source-parameter relations for mid-
plate earthquakes: Bulletin of the Seismological Society of Amer-
ica, v. 73, p. 519—535.

Obermeier, SR, 1988, Liquefaction potential in the central Mississippi
Valley: US. Geological Survey Bulletin 1832, 21 p.

1989, The New Madrid earthquakes—An engineering-geologic
interpretation of relict liquefaction features: US. Geological Sur-
vey Professional Paper 1336—B, 114 p, 11 plates.

Obermeier, SE, and Wingard, NE, 1985, Potential for liquefaction
in areas of Modiﬁed Mercalli intensities IX and greater, in
Hopper, M.G., ed., Estimation of earthquake effects associated
with large earthquakes in the New Madrid seismic zone, US.
Geological Survey Open—File Report 85—457, p. 127—140.

Obermeier, S.F., Gohn, G.S., Weems, R.E., Gelinas, R.L., and Rubin,
M., 1985, Geologic evidence for recurrent moderate to large
earthquakes near Charleston, South Carolina: Science, v. 227, p:
408—411.

Obermeier, S.F., Jacobson, R.B., Powars, D.S., Weems, R.B., Hall-
bick, D.C., Gohn, G.S., and Markewich, H.W., 1986, Holocene and
late Pleistocene (‘1) earthquake-induced sand blows in coastal
South Carolina: Proceedings of the Third US. National Confer—
ence on Earthquake Engineering, Earthquake Engineering
Research Institute, v. 1, p. 197—208.

Obermeier, S.F., Weems, R.E., Jacobson, R.B., and Gohn, G.S., 1989,
Liquefaction evidence for repeated Holocene earthquakes in the
coastal region of South Carolina: Proceedings of Conference on
Earthquake Hazards and Design of Constructed Facilities in the
Eastern United States, Annals of the New York Academy of
Sciences, February 24—26, 1988, v. 558, p. 183—195.

O’Connell, D., Bufe, C.G., and Zoback, MD, 1982, Microearthquakes

l and faulting in the area of New Madrid, Missouri—Reelfoot Lake,

Tennessee, in McKeown, F.A., and Pakiser, L.C., eds., Investi-
gations of the New Madrid, Missouri, earthquake region: US.
Geological Survey Professional Paper 1236, p. 31—38.

Penick, James, Jr., 1976, The New Madrid earthquakes of 1811—12:
Columbia, University of Missouri Press, 181 p.

Peters, K.E., and Herrmann, R.B., compilers and eds., 1986, First-
hand observations of the Charleston earthquakes of August 31,
1886, and other earthquake materials: South Carolina Geological
Survey, Bulletin 41, 116 p.

Pettijohn- F.J., and Potter, P.E., 1964, Atlas and glossary of primary
sedimentary structures: New York, Springer-Verlag, 370 p.
Russ, D.P., 1982, Style and signiﬁcance of surface deformation in the
vicinity of New Madrid, Missouri, in McKeown, FA, and
Pakiser, L.C., eds., Investigations of the New Madrid, Missouri,
earthquake region: US. Geological Survey Professional Paper

1236, p. 94—114.

Sanders, J .E., 1960, Origin of convoluted laminae: Geology Magazine,
v. 97, p. 409—421.

Saucier, R.T., 1964, Geological investigation of the St. Francis Basin,
lower Mississippi Valley: US. Army Corps of Engineers, Water-
ways Experiment Station, Technical Report 3—659.

1974, Quaternary geology of the lower Mississippi Valley:

Arkansas Archeological Survey, Research Series No. 6, 26 p.

1977, Effects of the New Madrid earthquake series in the

Mississippi alluvial valley: US. Army Corps of Engineers, Water-

ways Experiment Station, Miscellaneous Paper S—77-5, 10 p.

1989, Evidence for episodic sand-blow activity during the

1811—12 New Madrid (Missouri) earthquake series: Geology, v. 17,

p. 103~106.

 

 

 

 

 

 

Science News, 1986, A century after the Charleston earthquake: v.
129, no. 17, April 26, 1986, p. 263.

Scott, R.A., and Zuckerman, R.A., 1973, Sandblows and liquefaction,
in Committee on the Alaska Earthquake of the Division of Earth
Sciences National Research Council, 1973, The Great Alaska
Earthquake of 1964: Washington, D.C., National Academy of
Sciences, p. 179—189.

Seeber, L., and Armbruster, J .G., 1981, The 1886 Charleston, South
Carolina, earthquake and the Appalachian detachment: Journal of
Geophysical Research, v. 86, p. 7874—7894.

Seed, H.B., and Idriss, I.M., 1982, Ground motions and soil liquefac-
tion during earthquakes: El Cerrito, Calif., Earthquake Engineer-
ing Research Institute, Monograph series entitled “Engineering
Monographs on Earthquake Criteria, Structural Design, and
Strong Motion Records,” 134 p.

Seed, H.B., Idriss, I.M., and Arango, I., 1983, Evaluation of liquefac-
tion potential using ﬁeld performance data: Proceedings of the
American Society of Civil Engineers, Journal of the Geotechnical

' Engineering Division, v. 109, no. GT3, p. 458482

Sims, J .D., 1975, Determining earthquake recurrence intervals from
deformational structures in young lacustrine sediments: Tectono-
physics, v. 29, p. 141—152.

Smith, F.L., and Saucier, R.T., 1971, Geological investigation of the
Western Lowlands area, lower Mississippi Valley: US. Army
Corps of Engineers Waterways Experiment Station, Vicksburg,
Miss., Technical Report No. S—71—5.

Steele, F., Daniels, R.B., Gamble, E.E., and Nelson, L.A., 1969,
Fragipan horizons and Be masses in the middle coastal plain of
north central «North Carolina: Soil Science Society of America
Proceedings, v, 33, p. 752—755.

Street, R.L., and Nuttli, O.W., 1984, The central Mississippi Valley
earthquakes of 1811-12, in Gori, P.L., and Hays, W.W., eds.,
Proceedings of the symposium on the New Madrid seismic zone:
US. Geological Survey Open-File Report 84—770, p. 33—63.

Szabo, B.J., 1985, Uranium-series dating of fossil corals from marine
sediments of southeastern United States Coastal Plain: Geological
Society of America Bulletin, v. 96, p. 398—406.

Weems, R.E., Obermeier, S.F., Pavich, M.J., Gohn, G.S., Rubin, M.,
Phipps, R.L., and Jacobson, R.B., 1986, Evidence for three
moderate to large prehistoric Holocene earthquakes near Charles-
ton, South Carolina: Proceedings of the Third US. National
Conference on Earthquake Engineering, Earthquake Engineering
Research Institute, v. 1, p. 3—13.

Wesnousky, S.G., Schwig, ES, and Pezzopane, SK, 1987, Observa-
tions of soil liquefaction in the New Madrid seismic zone: Proceed-
ings of symposium on*“Seismic hazards, ground motions, soil
liquefaction and engineering practice in eastern North America,”
Oct. 20—22, 1987, New York, National Center for Earthquake
Engineering Research.

Youd, T.L., 1973, Liquefaction, ﬂow, and associated ground failure:
US. Geological Survey Circular 688, 12 p.

1978, Major cause of earthquake damage is ground failure: Civil
Engineering, American Society of Civil Engineers, v. 48, no. 4, p.
47—51. ,

Youd, T.L., and Perkins, D.M., 1978, Mapping liquefaction—induced
ground failure potential: Proceedings, American Society of Civil
Engineers, Journal of the Geotechnical Engineering Division, v.
104, no. GT4, p. 433—446.

Zoback, M.D., Hamilton, R.M., Crone, A.J., Russ, D.P., McKeown,
F.A., and Brockman, SR, 1980, Recurrent intraplate tectonism
in the New Madrid seismic zone: Science, v. 209, no. 4660, p.
971—976.

 

U.
U.

Z X NOIIVHEIOEJVXEI TVOIlIlEI/I

S. DEPARTMENT OF THE INTERIOR
S. GEOLOGICAL SURVEY

000'08

 

 

000'08

     
       

‘lEl/G'l VHS

133:!

B p
b o

O
8 o

STOCKTON FAULT

310

Stanislaus River

HOHV NOJ.)I30J.S

Tuolumne River

__SE_CTLOE

Merced River

BEND IN
SECTION

San Joaquin River

310

San Joaquin River

San Joaquin River
BEND IN
SECTION

510

BEND IN SECTION

SECTION B , 8‘

Kings River

peg 9x91 BJB|n_L

SECTION
c 7 C‘

HOHV CITEIIdSHENVG

      

BEND IN
SECTION

WHITE WOLF FAULT

PLEITO FAULT

BEND lN
SECTION

.1
<

,aq

SAN ANDREAS FAULT

.8 .5
§ §

TEAS") VEIS

 

'IEIAEIT VHS —
SHELBW

 

.920 ,

 

 

 

 

 

 

 

 

ox.

 

\\

 

 

 

 

 

 

.- grammar

 

 

 

 

 

 

I?
’70’ 50b. o o
S” INTERIOR—GEOLOGICAL sunvev, HESTON,VA~1991 35" 3°
Base from US. Geological Survey SCALE 1 150° 00° Geology compiled by J.A. Bartow, 1986-87, from
- - 10

State base map, compiled 1968, revrsed 1981 1,°_, ,_, H ,_, ,_. 4 2° 3° 4,0 M' LES L05 Angeles 09113111119111 ofWater 811d Poyver (1975.
Lambert Conformal Conic ro'ection based ﬁg. 25. MM), Harding (1976, ﬁg. 3a), Jennlngs (1977),

p I 1 o 1 o 20 30 4o 50 Dibblee (1979a), Herd (1979a), Webb (1981, ﬁg. 3), and

on standard parallels 33° and 45°

 

l—Il—ll—Il—ll—l

CONTOUR INTERVAL 500 FEET
NATIONAL GEODETIC VERTICAL DATUM OF 1929

60 KILOMETERS
j

Bartow(1984, 1985)

FEET

SEA LEVEL

     

CU

WALTHAM CANYON
FAULT

OTs TJ 9V

FEET

SEA LEVEL

10,000

20,000

COALINGA ANTICLINE

Tng (7)

21>

ORTIGALITA
FAULT

f

   

 

 

 

 

 

 

 

 

 

 

 

: a We we,
’ 9% ‘18 \ mu
mar...
sassmwmaa

Rx_.“_r"~___7
1:71. .

11

m l

FAULT ZONE

 

SAN JOAQUIN

 

 

 

 

 

 

CORRELATION OF MAP UNllS
QUATERNARY
N eogene CENOZOIC
TERTIARY
Paleogene
CRETACEOUS
AND MESOZOIC
JURASSIC
MESOZOIC AND PAIEOZOIC
DESCRIPl ION OF MAP UNI l S

Alluvial and lacustrine sediments (Quaternary)

Alluvial deposits, sedimentary rocks, and minor volcanic
rocks, undivided (Quaternary and Tertiary)—Used on cross
sections only and consists of units Os and Ts. Dashed line shows
approximate boundary between Paleogene and Neogene

Volcanic rocks (Tertiary)

Sedimentary rocks (T ertiary)—Marine and nonmarine rocks.
Dashed line shows approximate boundary between Paleogene
and Neogene

Great Valley sequence (Tertiary to Jurassic)—Melanges of sedi-
mentary rocks

Franciscan Complex (Tertiary to Jurassic)—Melanges of sedi—
mentary rocks, serpentinite, and blueschist in a sheared matrix and
coherent sedimentary rocks

Ultramafic rocks (Cretaceous and Jurassic)

Crystalline rocks of the basement complex (Mesozoic and Paleozoic)

Approximate area of structural arch

Contact—Queried where uncertain in cross sections only

Fault—Dashed where approximately located; dotted where concealed;
queried where uncertain. Bar and ball on downthrown side; R on
upthrown side of reverse fault; arrows indicate direction of relative

movement. Relative lateral movement on cross sections only: A, away

from observer; T, toward observer

Thrust fault—Dotted where concealed, Sawteeth on upper plate

Fold—Dotted where concealed. Showing approximate trace of axial
plane; arrows indicate direction 04f plunge

Anticline
Syncline

Boundary of structural region (see text)—Dashed where approxi—
mately located

Location of stratigraphic column—Number refers to column number
on plate 2

SAN JOAQUIN
VALLEV SYNCLINE
San Joaquin River

 
   

  
 

 
 
    

PROFESSIONAL PAPER 1501
PLATE 1

491%
#5
PC,

 

  
   

  

scar;
No. 1601
I ( e)
F313 2 I) 1992 can 7»

a 5

~LI.) 3

§§ 5% ,

s; A

n:< m

ELL be 2 bC FEET METERS

’ SEA LEVEL SEA LEVEL

 

SAN JOAQUIN VALLEY
SYNCLINE

TURK ANTICLINE

VERTIICAL EXAGGERATION X2

OTs

l
|
I
I
I
|
l
l
I
I
|
|
|
I
|
I
I
l

   

 

 

10,000
5,000
20.000
3!
OT bC FEET METERS
s _
’ SEA LEVEL —— SEA LEVEL
10,000 —
— — 5.000
— 20,000 _
; 30,000 _
: -10,000
— 40,000 _
— 50000 ‘15-000

 

 

0

SAN ANDREAS

FEET

SEA LEVEL

10,000

20,000

FAU LT

RECRUIT PASS FAULT

GENERALIZED GEOLOGIC MAP AND CROSS SECTIONS OF THE SAN JOAQUIN VALLEY AREA, CALIFORNIA

  

VERTICAL EXAGGERATION X 2

GRE E LEY FAULT SYSTEM

 

POSO CREEK FAULT

 

VERTICAL EXAGGERATION X 2

SOURCES OF lNFORMATlON
A—A' Wentworth and others (1984, fig. 4), Bartow
(1985), Wentworth (written commun., 1986)
B— ' Hoots and others (1954, pl. 6), Wentworth (1985)
C—C’ Church and Krammes (1957), Vedder (1970). Bartow (1984)
D— ' Church and Krammes (1958), Davis (1986)

 

METERS

SEA LEVEL

5,000

 

 

 

DEPARTMENT OF THE INTERIOR
U.S. GEOLOGICAL SURVEY

PROFESSIONAL PAPER 1503

PLATE 1

82°

83°

84°

85°

86°

87°

89°

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

SCALE 1:1000 000

50 MILES

25

 

 

 

 

F’——r

25

ERS

50 KILOMET

25

25

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

a? M if” _-

v’?

 

 

W

,"

s

 

 

 

 

 

 

 

 

 

 

     

*4va

 

 

 

 

 

 

 

 

 

 

“1%

8%
&v%

 

 

 

 

 

 

 

 

 

 

 

 
     

ml

s

 

 

  

 

  

 

 

l
85" 84" 83° 82°

86°

87°

l‘Genlogic quadrangle name is HAMMACKSVILLE,

3Geologic quadrangle name is MODEL.

88°

2Genlogic quadrangle name is HAZEL.

89°

lGenlogic quadrangle name is NEW MADRID SET

 

 

39° —

_
o
7
3

38° —

MAP OF KENTUCKY SHOWING 71/2’ QUADRANGLES

 

DEPARTMENT OF THE INTERIOR
U. S. GEOLOGICAL SURVEY

  

ORESTIMBA CREEK

SERIES AND AREA

SUBSERIES1 PALEO—

BATHYMETRY 3
DM SM NM SW

STRATIGRAPHY

SW NE

SYSTEM AND
SUBSYSTEM

uvium

HOLOCENE
PLEISTOCENE

O
C
]>
_4

ulare ormation u are F

PLIOCENE

"Del-
mon-
tian“

\
\

Fanglomerate

Oro Loma
Formation

(Carbona unit of
Raymond, 1969)

Mehrten
Formation

Mohnian

NEOGENE

Luisian

MIOCENE

Reliziﬁ .

Saucesian

Valley Springs
Formation

OLlGOCENE
Zemorrian

TERTIARY

Poverty Flat
Sandstone

Kreyenhagen Kreyenhagen

PALEOGENE
Narizian

EOCENE

Domengine
Sandstone

Domengine
Sandstone

"Penutian"
Ulatisian :

Tesla

Laguna Seca

Formation

BUI't'an Formation

PALEOCENE

Cheneyan w

1 Series and subseries boundaries from Berggren and others (1985).
2 Stage boundaries as follows: top of "Delmontian," Obradovitch and Naeser (1981);

"Delmontian"—Mohnian, MohnianvLuisian, and Luisianr
Relizian, Barron (1986); Relizian-Saucesian, Miller (1987); Saucesian-Zemorrian, Turner (1970), adjusted for new constants according to
Dalrymple (1979), and Poore and others (1981); Zemorrian—Refugian and Refugian-Narizian, Poore (1980); Narizian-Ulatisian, Ulatisian—
Penutian, Penutian—Bulitian, and Bulitian—Ynezian, Almgren and others (1988); Ynezian—Cheneyan, Poore (1980).

3 DM, deep marine (below shelf depth); SM, shallow marine (shelf depth); NM, nonmarine. Short dashes (- — —), water depth inferred; queried where

uncertain

" Based on data from Atwater (1970), Atwater and Molnar (1973), Dickinson and Snyder (1979), Page and Engebretson (1984), and Engebretson

and others (1985).

5 From Huber (1981).
6 Based on data from Evemden and others (1964), Turner (1970), Prowell (1974), Stewart and Carlson (1976), Dickinson and Snyder (1979),

Moore and Dodge (1980), Huber (1981), and Drinkwater (1983).

7 Modified from Haq and others (1987), adjusted to chronology of Berggren and others (1985). TA1 through T133 represent supercycles in the

sequence chronostratigraphy of Haq and others (1987).

3 Informal name of local usage.
9 Of Miller and Bloom (1937).

LOS BANOS-
ORO LOMA AREA

STRATIGRAPHY

NORTHEAST
VALLEY MARGIN
PALEO—

BATHYMETRY
DM SM N

BATEQLMEST'RY 3 STRATIGRAPHY

NE DM SM NM SW NE

Mehrten

Table Mountai
Latite

Formation

prlngs

Fm. Valley Springs

Formation

Formation

Tes|a(?) Formation

Moreno
Formation

VALLECITOS
SYNCLINE

STRATIGRAPHY

Nonmarine
rocks

Temblor L 1
Formation

Tumey
Formation

Shale
member

Kreyenhagen

o
Arroyo
Hondo

hale Member

Cantua
Sandstone Member

Cerros Shale
Member

Moreno
Formation

Sandstone
member

w.
5”.

Lodo Formation

KETTLEMAN HILLS
NORTH DOME

PALEO—
BATHYMETRY 3

DM

BATmLhEST—Rys STRATIGRAPHY

DM SM NM NW SE

Tulare Formation

San Joaquin

Etchegoin
Formation

Santa Margarita
ormation
Ridge.

McLure Shale
Member of
Monterey Formation

Upper variegated
unit 8
Temblor Formation
(upper part)

ower variegated
unit 8

Burbank sand 8

Whepley
shale8

Formation
part)

"Vaqueros"
sand 8

"Salt Creek"
shale 8

"Leda" sand 8

Tumey Shale

Kreyenhagen

Lodo

Formation

Moreno
Formation

SM NM SW

HAN FOR D-TU LARE
EAST-SIDE AREA
PALEO-

BATHYMETRY 3
DM SM NM

STRATIGRAPHY STRATIGRAPHY

NE

Tulare Formation L

San Joaquin

Kern River
Etchegoin

Formation Formation

McLure Shale
Member of
Monterey Formation

McLure Shale

Member

Devilwater Sh. and
Gould Sh. Mbrs.,

Monterey Shale

Olcese Sand

deposits,
undivided

- Buttonbed
Sandstone Member
Media Shale

Freeman Silt Member

Santos e
Member
(upper part)

undivided

Agua 85. Bed

Santos Shale
Member
(lower part)

Temblor Formation

Vedder(?)

SS. Mbr'.

Cymric
Shale

Member

Wagonwheel
Formation

Welcome

Shale

Kre enha en
y g Member

ale

Point of Rocks
Sandstone
Member

Gredal
Shale

Kreyenhage

Member
Avenal

amoso sand8 Sandstone

Lodo

Formation

Moreno
Formation

EXPLANATION

[ED] Hiatus

 

maps (figs. 5-13)

 
 

. $2325: Subduction

llllllll Wrench faulting

Rhyolite

"""IHIIIIII Andesite or rocks of mixed composition
_ Basalt

 

 

LOST HILLS-
DEVILS DEN AREA

PALEO—
BATHYMETRY 3

DM

Approximate time represented by paleogeographic

——-—?—Contact—Dashed where indefinite; queried where un-

certain

r\/\/\?I\ Unconformity—Queried where uncertain

SM NM SW

ELK HILLS

AREA

STRATIGRAPHY
NE

m

Tulare Formation

Etchegoin

Formation

Stevens

sandstone 8

Sh. M

McDonald Sh. Mbr.8

Devilwater Sh. and
Gould Sh. Mbrs.,

Monterey Shale

Media Shale
Member

83.
M

#Sandstone
Santos

Shale

Member

Temblor Formation

Wygal Sandstone
Member

Cymric
Shale

Member

Wagonwheel
Formation

Kreyenhagen

Tejon(?) Fm

PALEO-
BATHYMETRY 3

DM

 

BAKERSFIELD ARCH 10

AREA

BATWJST'W STRATIGRAPHY

DM SM NM N 3

LI

STRATIGRAPHY

SM M SW NE

Tulare Formation
Kern River
Formation Etchegoin Fm. and
nonmarine deposits,

Etchegoin undivided

Formation

Margarita" Fm.

' 9
Fruitvale Shale 9 Fruttvale Shale

Round

Round Mountain Silt .
Mountain

Monterey Shale

Olcese Sand

Freeman

Tecuya
Formation

Volcanic rocks

Silt Temblor
Jewett

Sand Formation

Formation

Vedder

Tecuya

Pleito

Formation

?

San 2

Z,
Emigdio '

Walker Formation

Formation

KreYenhagen Shale

Famoso

_-

a
1’ .

Formation

r
a
’a

a
,L_____

z

”a

SAN EMlGDlO
MOUNTAINS AREA

PALEO—
BATHYMETRY 3

DM

SM NM

TANGENTIAL MOVEMENT
Pacific plate: North American plate

NORMAL CONVERGENCE
Farallon plate: North American plate

OBLIQUE CONVERGENCE
Farallon or Kula plate: North American plate

PLATE TECTONIC
EVENTS4

Mendocino triple junction

 

 

Slight convergence in

Pacific and North

American

migrating northwestward
III IIIIIIIIIIIIIIIIIIIIlllllIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIII

IIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIIII7 Proto-San Andreas fault

past San Joaquin basin

e motio

ﬂ'i‘l

San Andreas

_|

transform

Wrench faulting concurrent

OCOCDNO) thN—I

  

PROFESSIONAL PAPER 1501
PLATE 2

SIERRA
NEVADA
UPLIFT RATE 5
mm /yr
0.4 .3 .2 .1

AGE
(Ma)

'N EUSTATIC SEA LEVEL
ADJACENT

AREAS 6

250 200 150 100 50 O METERS

Slip concen—
trated on San
Andreas fault

TBS

1)

Carson Pass-Sonora Pass

0.)
U)
C
(U
I
2
.0
.9
Q

Opening of
Gulf of
California

 

Deadman Pass

on San Andreas
fault
TBZ

First movement(?)
m w m Western Nevada / eastern California

lllllllllllll<—San Emigdio Mountains/ northern Gabilan Range
rn II|IIIIIIIIIIIIlIIIllllllllllllllllllllllIlllllllllllllllllIIIIlIllIlIIIIIIl

Central and northern lIIIIIII|llll|IIIlIIIIIIIIIIIIIIIIIIIIIIII

Central / southern Sierra Nevada — - - -_

{Southe

Sierra Nevada
Santa Cruz Mountains -

 

 

 

Neenach / Pinnacles areas’nn

Wm
WWW

With oblique subduction orocllnal
bending of southern Sierra Nevada

SOURCES OF STRATIGRAPHIC INFORMATION
[Number refers to column number on plate 2 and corresponding area on plate 1]

Bartow (1985)

Briggs (1953), Lettis (1982), Bartow (1985)

Bartow (1985)

Phillips and others (1974), Dibblee (1979b), Nilsen (1979), Berggren and Aubert (1983) Rentschler (1985)

Woodring and others (1940), Church and Krammes (1959), Sullivan (1966), Graham and others (1982),
Kuespert (1983), Bartow (1990)

Dunwoody (1969), Bartow (1990)

Dibblee (1973b), Graham and others (1982) Berggren and Aubert (1983), Bartow (1990)

Church and Krammes (1957), Dibblee (1973b), Maher and others (1975)

Bartow and McDougaIl (1984)

Hluza (1960), Nilsen and others (1973), Davis (1983, 1986), DeCelIes (1986), Lagoe (1986)

CORRELATION OF CENOZOIC STRATIGRAPHIC UNITS OF THE SAN JOAQUIN BASIN WITH TECTONIC, VOLCANIC, AND SEA-LEVEL EVENTS

?5
[7(9

595121777 6::
N U! (I 901
2. g; )

(or! L