‘Application of Total-Count Aeroradiometric

Maps to the Exploration for Heavy-Mineral
Deposits in the Coastal Plain of Virginia

By ANDREW E. GROSZ

With a section on Field-Spectrometer—Data Reduction

By KENNETH L. KOSANKE

 

GEOLOGICAL SURVEY PROFESSIONAL PAPER 1263

A method of exploration for nearshore marine placer
heavy-mineral deposits by the use of gamma-ray
ground-radiometrzc and aeroraa'iometiic techniques

 

 

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON: 1983

UNITED STATES DEPARTMENT OF THE INTERIOR

WILLIAM P. CLARK, Secretary

GEOLOGICAL SURVEY

Dallas L. Peck, Director

 

Library of Congress Cataloging in Publication Data
Grosz, Andrew E.
Application of total-count aeroradiometric maps to the exploration for heavy-mineral deposits in the Coastal plain of Virginia.

(Geological Survey professional paper; 1263)
“With a section on field-spectrometer-data reduction by Kenneth L. Kosanke.”
Bibliography: p.
Supt. of Docs. no.: I 19.16:1263
1. Heavy minerals—Virginia. 2. Radioactive prospecting—Virginia.

I. Kosanke, Kenneth L. Field-spectrometer—data reduction. 1983. II. Title. III. Series.
TN24.V8G76 622’.18 81-607581

AACRZ Rev.

For sale by the Superintendent of Documents, U.S. Govemment Printing Office
Washington, D.C. 20402

 

CONTENTS

 

 

 

 

 

 

 

 

 

 

 

 

 

 

Page
Abstract 1 Field methods
Introduction 1 Effects of fertilizer use on aeroradiometric measurements
Previous work 2 Laboratory analyses
Present work 2 Techniques
Acknowledgments 3 Results
The study area 3 Spectral radiometric characterization of anomalies -------
Morphology 3 Summary and conclusions --------------------------
Geology 3 Field-spectrometer-data reduction, by Kenneth L. Kosanke
Evaluation of the aeroradiometn'c data ---------------- 4 Selected references
ILLUSTRATIONS
[Plates are in pocket]

FIGURE

TABLE

FPS-”5°?

2"

. Plot of sequence of spectrometer gain settings during the field-data collection period

Generalized geologic map of the Outer Coastal Plain of Virginia.

Total-count contoured gamma-ray aeroradiometric map of the Outer Coastal Plain of Virginia.

Total-count contoured gamma-ray aeroradiometric map of the Portsmouth area, Virginia.

Land-use and land-cover classification of aeroradiometric anomalies in the Outer Coastal Plain of Virginia.

Size distribution of heavy-mineral species in 15 samples that contain more than 1 percent total heavy minerals collected
from aeroradiometrically anomalous localities in the Coastal Plain of Virginia.

Index map showing the relation of the Coastal Plain of Virginia to the State. the Outer Coastal Plain (the study area). and

the Inner Coastal Plain
Histogram of heavy-mineral percentages in 80 samples from aeroradiometrically anomalous localities in the Coastal
Plain of Virginia
Gammaray spectrum of a bulk source of K, U, and Th, showing the location of the improperly positioned spectrometer
windows

 

 

 

 

 

TABLES

Ground-radiometric and aeroradiometric signatures of sample localities in the Coastal Plain of Virginia —————————————————

2. Sieve analyses and heavy-mineral contents of 65 samples that contain less than 1 percent total heavy minerals collected

 

from aeroradiometrically anomalous localities in the Coastal Plain of Virginia
Sieve and heavy-mineral analyses of 15 samples that contain more than 1 percent total heavy minerals collected from
aeroradiometrically anomalous localities in the Coastal Plain of Virginia
Averages of equivalent thorium, equivalent uranium, and percent potassium in different lithologies measured in the
field

 

 

. Field-spectrometer background data before and after a change in gain setting coinciding with a repair of the instrument
. Calibration matrices (A) for the spectrometer before and after change in performance

 

III

@QO‘GOI

13

17
18

Pass

17
18

Page

10

14

16

17
18

 

APPLICATION OF TOTAL-COUNT AERORADIOMETRIC MAPS
TO THE EXPLORATION FOR HEAVY-MINERAL DEPOSITS
IN THE COASTAL PLAIN OF VIRGINIA

By ANDREW E. GROSZ

ABSTRACT

Total-count contoured aeroradiometric maps for the Coastal Plain
of Virginia were used in an effort to locate economic heavy-mineral
placer deposits. The principle behind this approach is that heavy-
mineral suites commonly contain radioactive minerals that, if the
concentration of heavy minerals is exposed at or within inches of
the surface, enable the deposit to be located by use of airborne
instruments because of its radiometric contrast with the host
sediment.

Detailed and regional geologic maps, soil maps, land-use and land-
cover maps, information on fertilizer use, and ground-spectrometer
data were used to study aeroradiometric anomalies for efficient ex-
ploration. Aeroradiometric anomalies in the Coastal Plain of
Virginia have three general causes. First, the most intense
anomalies are associated with cultural features, such as roads made
of granitic material. Second, most anomalies of high to intermediate
intensity are associated with land used for agricultural purposes
and evidently are caused by applications of radioactive fertilizer.
Third, anomalies of intermediate to low intensity are associated
with heavy-mineral deposits.

Results of this study show that aeroradiometric anomalies asso
ciated with heavy-mineral accumulations in the Coastal Plain of
Virginia have ground radiometric spectra in which thorium is the
strongest component and uranium and potassium are lesser
components.

Heavy-mineral accumulations found in this study by use of the
aeroradiometric data are not considered to be of economic impor-
tance, mostly because of the low percentage of economic minerals in
the heavy-mineral suites and also because of other factors such as
the very fine grained nature of the host sediments and competing
land use.

INTRODUCTION

Economically valuable heavy-mineral placer deposits
in nearshore marine sediments on the Atlantic Coastal
Plain are sources for a significant fraction of much-
needed titanium dioxide minerals in the United States.
Currently, most of the demand for such minerals is sup-
plied by foreign imports, particularly from Africa,
Brazil, and Australia. This investigation is primarily
concerned with locating such deposits on the Outer
Coastal Plain of Virginia by the use of aeroradiometric

 

maps, but it also documents a method of approach to ef-
ficient exploration and to the interpretation and evalua-
tion of such maps for other purposes in coastal areas.
Valuable heavy-mineral accumulations in sand ore
bodies on the Atlantic Coast have been discovered by a
variety of techniques. Deposits have been discovered by
geologic reasoning and shallow augering (Spencer, 1948;
Markewicz and others, 1958). Aeroradiometric surveys
have played, or in hindsight could have played, a part in
the discovery of several deposits, such as those at
Folkston, Ga. (Moxham, 1954), Green Cove Springs,

. Fla. (James Hetherington, oral communication, Decem-

ber 18, 1974), and Brunswick, Ga. (Stockman and
others, 1976). Surface sampling and shallow drilling
have been relied on in exploration efforts; the high costs
associated with such programs, however, warrant new
exploration techniques, such as the application of aero-
radiometric data, so that large areas can be scanned at
lower cost.

Extensive aeroradiometric surveys for the southeast-
ern Atlantic Coastal Plain States have been contracted
for by the United States Geological Survey (USGS).
Funding for the surveys and subsequent field investiga-
tions was supplied by the Coastal Plains Regional Com-
mission. The objective behind these surveys was to de-
tect concentrations of radioactive minerals exposed at
the surface.

Exploration for heavy minerals by the use of aero-
radiometric surveys is based on the assumption that
radioactive heavy minerals such as monazite, sphene,
and zircon are concentrated with nonradioactive heavy
minerals such as ilmenite, rutile, and sillimanite to form
placer deposits. Wave, tidal, and wind actions are the
mechanisms by which the heavy minerals are concen-
trated on present-day shorelines in beach sand. Concen-
trations are also present in former shoreline deposits
now found inland, commonly parallel to present shore-
lines. Placer deposits thus formed will exhibit radio-
metric contrast to their host sediment if radioactive
minerals are present in the heavy-mineral suite. Such

1

2 AERORADIOMETRIC MAPS IN EXPLORATION FOR HEAVY-MINERAL DEPOSITS

radiometric contrast is, in principle, detectable by air-
borne scintillation counters.

The study area is Virginia’s Outer Coastal Plain and
includes approximately 2,500 square miles (6,475
square kilometers), including the Delmarva Peninsula
(fig. 1). Parts of this area have been settled since colonial
days and are focal points of commercial, industrial, and
military activities on the Atlantic coast. With the ex-
ception of two broad clusters of urbanized areas (the tip
of the J ames-York Peninsula and the Norfolk-
Portsmouth-Chesapeake—Virginia Beach complex), the
bulk of the study area is used for agricultural purposes
and is sparsely populated.

Inasmuch as the James River drains an ilmenite-rich
terrane (Minard and others, 1976), the probability of
heavy-mineral concentrations on the Coastal Plain is
good. Because shoreline sands elsewhere are commercial
sources of heavy minerals, Force and Geraci (1975)
undertook a study to evaluate the possibility of heavy-
mineral deposits in the Pleistocene(?) shoreline sands in
eastern Virginia. A suite of 53 samples (pl. 1) was col-
lected and processed for heavy minerals in methylene
iodide (s.g. (specific gravity) 3.3), but the heavy-mineral
contents were much lower than those of deposits pres-
ently being mined. Of particular interest to this study is
that the radioactive-mineral content of the heavy-
mineral suites is very low to absent. The implications
‘ behind this are that, if the most promising lithology for
heavy-mineral deposits in the area is low in radioactive
heavy minerals, other less favorable lithologies are not
likely to contain much monazite and zircon; as a conse-
quence, aeroradiometric contrasts of heavy-mineral ac-
cumulations with respect to barren host sediments were
expected to be low.

PREVIOUS WORK

A large amount of literature on the theory and use of
airborne radiometric surveys exists (for example,
Moxham, 1954, 1960; Gregory, 1960; Schmidt, 1962;
Mahdavi, 1964; Pitkin and others, 1964; Neuschel,
1970; Neuschel and others, 1971; Neiheisel, 1976; Perl-
man and others, 1976; Stockman and others, 1976;
Force and Bose, 1977). Until recently, however, little
has been published on the application of such data to
the exploration for placer heavy minerals in coastal
areas. A study of the applicability of aeroradiometric
maps to placer prospecting in South Carolina’s Coastal
Plain (Force and others, 1978) was recently published; it
presents a method of classifying aeroradiometric
anomalies into different types. That study uses airborne
total-count and airborne spectral radiometric data in
conjunction with county soils maps, regional miner-

 

EXPLANATION n‘
lnner Coastal Plain F
a I“
Outer Coastal Plain
IV

 

0 50 100 MILES

  
 

6 5'0 100 KILOMETERS /

v/«\/\,‘r”

VlRGlNl'AV

.J
‘4 .-——L:—_l____(___l.__l_
83 82 81° 800 79° 78°

77" 76°

FIGURE l.—Index map showing the relation of the Coastal Plain of
Virginia to the State, the Outer Coastal Plain (the study area), and
the Inner Coastal Plain. Modified from Wentworth (1930, p. 2).

alogic trends, and regional geologic information. Re-
sults of that study show that aeroradiometric surveys
can be used .to find detrital heavy-mineral accumula-
tions containing radioactive heavy minerals. The study,
however, also shows that most aeroradiometric
anomalies are caused by deposits other than heavy
minerals and that the intensity of aeroradiometric sig-
nature of a heavy-mineral (deposit does not reflect its
economic value. Another recent study (Robson and
Sampath, 1977) tested the effect of heavy-mineral sand
deposits at Jerusalem Creek, New South Wales,
Australia, on a variety of geophysical instruments.
That study shows aeroradiometry to be among the more
promising geophysical techniques for locating heavy-
mineral deposits that crop out. Furthermore, the study
shows that, although the ore is highly radioactive (pre-
sumably because of monazite), 10 feet (3 meters) of over-
burden completely attenuated the response. However,
in general, the thickness of overburden needed to mask
gamma radiation is about an order of magnitude less
(Beck, 1975).

PRESENT WORK

Field investigations of aeroradiometric anomalies in
the study area Were conducted in the summer of 1978.
During this time, sediment and ground-radiometric
samples were collected for analyses, and several
anomalies west of the study area were sampled. All the
samples were analyzed in USGS laboratories at Reston,
Va. Ground-radiometric data were analyzed and cor-
rected by the Bendix Field Engineering Corporation,
Grand Junction, Colo. (see the section on “Field-
Spectrometer-Data Reduction”).

The method of approach in this study was partially

‘4

 

 

 

‘L

THE STUDY AREA 3

based on methods and results from previous studies on
the uses of aeroradiometric maps. The use of geologic,
land-use and land-cover, and fertilizer-use maps, in con-
junction with ground-spectrometer data, as “filters”
over the aeroradiometric maps is presented. This
method is part of an effort to identify anomalies caused
by materials other than heavy-mineral accumulations.

ACKNOWLEDGMENTS

It is a pleasure to acknowledge the aid of Isidore Zietz
in acquiring the aeroradiometric data, for suggesting
this study, and for arranging financial support. Funding
for the surveys and field investigations was supplied by
the Coastal Plains Regional Commission. The comple-
tion of this report was greatly facilitated by the friendly
cooperation, encouragement, and suggestions of many
interested persons. Discussions of ideas and helpful
criticisms were generously made at various stages in
this study by Eric R. Force, Frank E. Senftle, Robert B.
Mixon, David Gottfried, Zalman S. Altschuler, and
James P. Owens of the USGS.

Special appreciation is extended to James T. Lindow
of GeoMetrics, Inc., in Sunnyvale, Calif, who offered
many constructive criticisms and guidelines on explora-
tion involving spectral scintillometers.

Technical assistance by specialists has included (1)
X-ray mineralogical analyses by Patricia J. Loferski,
(2) computer program assistance by David R. McQueen,
and (3) size analyses by Dwight E. Wallace.

Property owners throughout the study area were
both friendly and helpful. I am grateful to all of these
people and many others for their generous help and
consideration. ‘

THE STUDY AREA
MORPHOLOGY

The study area has little relief. The most prominent
geomorphologic feature in the coastal area is the east-
facing Suffolk scarp and correlative scarps, which trend
north through Virginia’s Coastal Plain. This scarp is an
old shore face that forms a natural stratigraphic bound-
ary separating post-Pliocene units into an older group
to the west (the Inner Coastal Plain) and a younger
group to the east (the Outer Coastal Plain) (Oaks and
Coch, 1973, p. 4). Altitudes on the Inner Coastal Plain
range from 20 to 175 feet (6 to 53 meters) above sea
level, and relief is 20 to 50 feet (6 to 15 meters) near ma-
jor streams. The Outer Coastal Plain, in contrast, is char-

 

acterized by altitudes generally below 25 feet (8 meters)
above sea level, and relief is as much as 20 feet (6 meters)
locally along the James and Nansemond Rivers only.
Wide and flat undissected tracts form poorly drained
areas between major streams in both parts of the Coastal
Plain.

Depositional morphology predominates east of the
Suffolk scarp and correlative scarps, but fluvial erosion
has altered large parts of the original depositional sur-
faces west of the scarps. Linear morphologic features
have a north-south trend that is closely related to the
depositional morphology of ancient 'barrier-and-lagoon
environments (Oaks and others, 1974, p. 62). -

GEOLOGY

Published literature on the geology of Virginia’s
Coastal Plain spans more than 350 years and ranges
from the first known published reference to the gener-
alized geology of the Coastal Plain (John Smith, 1624,
cited in Wentworth, 1930) to more detailed studies in re-
cent years (Coch, 1968, 1971; Oaks and Coch, 1973;
Oaks and others, 1974; Johnson, 1972, 1976; Onuschak,
1973; Force and Geraci, 1975). Current USGS work on
the geology of the Delmarva Peninsula south of the
Maryland-Virginia State line (Robert B. Mixon, written
communication, 1979) provided additional information
on the character and composition of surficial lithologies
in the coastal area.

Recent workers have studied post-Pliocene sediments
east of the Suffolk scarp and correlative scarps to deter-
mine their origin, relation to sea level, geologic history,
and mineral composition. On the basis of these studies,
lithostratigraphic nomenclature for most surficial units
and pre-Pleistocene units has been suggested; major

groupings of sediments are thought to have been

deposited during submergent phases of emergent-
submergent cycles attributed fot the most part to gla-
cial eustatic effects. A generalized geologic map (pl. 1)
shows surface distributions of the units mapped in
southeastern Virginia and on the Delmarva Peninsula.
The oldest sediment in the study area is the Yorktown -
Formation of early Pliocene age, exposures of which, in
southeastern Virginia, are generally limited to river
bluffs and artificial excavations. The Yorktown Forma-
tion consists of marine clay, sand, silt, and coquinite.
Postdepositional fluvial erosion has created severe relief
locally in the top of this unit (Oaks and others, 1974,
p.62). '
Widespread, flat-lying units above the Yorktown
record a complex history of relative sea-level changes
through time. Oaks and Coch (1973, p. 108) recognized
six important periods of sea-level fluctuation, during

4 AERORADIOMETRIC MAPS IN EXPLORATION FOR HEAVY-MINERAL DEPOSITS

the submergent phases of which the following forma-

tions‘ were deposited, from oldest to youngest: the
Sedley Formation, the Bacons Castle Formation, the
“Moorings” unit, the Windsor Formation, the Great
Bridge Formation, the Norfolk Formation, the Kemps-
ville Formation, and Holocene units. Some of these for-
mations (for example, the Windsor Formation) are
found only west of the Suffolk scarp and correlative
scarps.

The post-Yorktown formations are predominantly
marginal marine clastic sediments, ranging in deposi-
tional environment from beach barrier to offshore de
posits and in grain size from clay to coarse sand. The
materials are predominantly quartz sand, silt, and feld-
spar; clay, mica, chert, and rock fragments are locally
important but generally are minor constituents. Oxida-
tion of exposed facies is generally 1 to 5 feet (0.3 to
1.5 meters) deep, whereas in older facies, oxidation
penetrates more than 10 feet (3 meters).

With the exception of Holocene facies, which are
virtually unweathered, most units show limonitic or
mottled hematitic (Coch’s Windsor Formation, upper
member) color in the weathered zone. Alteration of
quartz-grain surface features and of soluble heavy min-
erals is very slight to moderate with the exception of the
“Moorings” unit, where alteration is heavy. Soil is well
developed to a depth of approximately 3 to 4 feet (0.9 to
1.2 meters) on older units and, is virtually absent on
Holocene units. Clay enrichment is significant (0.25 to
8 feet; 0.08 to 2.4 meters) in lithologies having well-
developed soil zones and minor in lithologies without
soil zones (Oaks and others, 1974, table 2, p. 66).

Facies exposed at present on the Delmarva Peninsula
(units A through F, pl. 1) are of late Pleistocene and
Holocene age (Robert B. Mixon, written communica-
tion, 1979) and range in composition from the medium-
to coarse-grained sand, gravelly in part, interbedded
with lesser amounts of fine-grained sand, of banier
island or barrier spit complexes to fine-grained sand de-
posits on the bottom of an ancestral Chesapeake Bay.

EVALUATION OF THE AERORADIOMETRIC DATA

The total-count gamma-ray-intensity survey for
Virginia was flown and compiled for the USGS by Geo-
data International, Incorporated, in late 1976 and early
1977. Part of the survey covering southeastern Virginia
was flown and compiled in 197 5—76 by LKB Resources,
Incorporated (Force and Bose, 1977). The mismatch of
aeroradiometric contour lines at the border between the

"I'hese formations have not been adopted for US. Geological Survey usage; their usage
herein is based on the mapping of Oaks and Coch (1963. 1973) and Coch (1965, 1968).

 

two surveys is a phenomenon common to overlapping
airborne surveys. The instruments for such surveys are
generally not calibrated over pads of known radioele-
ment concentrations, and as a result, different instru-
ment packages and, more importantly, different detector
sizes will yield different results (in count-rate magnitude)
over the same sediments; these results can be correlated
only with difficulty.

The fundamental principle behind airborne radio-
metric surveys is that radioactive materials exposed at
the surface emit a spectrum of radioactivity that can be
measured by airborne scintillometers. Instruments of
this type commonly consist of a large-volume (400 to
500 cubic inches, 6,600 to 8,200 cubic centimeters) NaI
(sodium iodide) crystal coupled to an electronic system
that records the activity detected by the crystal. An air-
borne system registers radioactivity from three basic
sources, terrestrial, atmospheric, and cosmic. The
strongest component is the terrestrial source. Atmos-
pheric and cosmic sources generally account for a small
part of the total count rate observed but are highly
variable with time, altitude, and prevailing atmospheric
conditions; as a result, some airborne surveys register
substantial and variable radiometric‘count rates over
open bodies of water where values should be low and
constant. Airborne surveys are generally flown at an
altitude of 500 feet (150 meters), and flight lines are
oriented to cross the strike of geologic contacts. Spacing
between flight lines is commonly 1 to 1 1/2 miles (1.6 to
2.4 kilometers). Total-count contoured aeroradiometric
maps are the end products of such surveys; they are
commonly hand contoured, but recently computer pro-
grams have been written to do the contouring.
Aeroradiometric surveys, although more limited in
resolution and accuracy than surface methods, can best
show regional variations in radiation, from which some
estimates may be made of terrestrial radioelement con-
tent and of surface radiation intensities.

The aeroradiometric map (pl. 2) outlines major water
bodies in the study area, indicating that location ac-
curacy and instrument calibration are good. Large parts
of the survey, however, particularly west of the study
area, show major contrasts in radiometric signatures
that are inconsistent with other parts of the survey.
This inconsistency is due to switching from one type of
navigational system over the Outer Coastal Plain to
another over the Inner Coastal Plain and Piedmont
regions (Robert S. Foote, President, Geodata Intema-
tional, Incorporated, oral communication, 1978) and is
esentially a geographic location problem. Because the
areas of high contrast were well west of the study area,
the part of the survey covering the Outer Coastal Plain
is considered acceptable with the following stipulations.

Large parts of the survey, particularly near the

4‘

 

 

EVALUATION OF THE AERORADIOMETRIC DATA 5

Virginia-Maryland border and near the mouth of Chesa-
peake Bay, registered significant count rates over open
bodies of water where signatures should be very close to
zero. Because count rates of 30 to 60 counts per second
were registered over open water, radiometric values
over land are subject to an error equaling the count rate
observed over water for those flight lines that registered
greater than zero counts per second over water.

Further complications arise from the computer con-
touring technique. The total-count aeroradiometric map
of the Portsmouth area (pl. 3) is used as an example of
the computer contouring technique and its results. By
use of a matrix of measured values from adjacent east-
west oriented flight lines, “hybrid” numbers were first
extrapolated and then contoured. Such “hybrid”
numbers appear between flight lines and represent aver-
aged values that do not correspond to actual measured
values, either in their numerical value or in geographic
location. Plate 3 also shows how a point source of radia-
tion is treated by the plotter. The “dipole-like” anomaly
southeast of the Elizabeth River consists of a radio-
metric high (362 counts per second) and a proximal
radiometric low (48 counts per second). The high value
corresponds to radioactive source(s) within the main
building of the Portsmouth Naval Hospital—a point
source for all practical consideration—exaggerated by
contouring to more than a mile (1.6 kilometers) long and
to 1/3 of a mile (0.5 kilometer wide); the low is associated
with a major drainage that is truncated at the north by
the distended high. Such distortion is increased by
flight-line spacing greater than 1 mile (1.6 kilometers)
and is typical of areas where the count rate exceeds 200
counts per second.

In a preliminary reconnaissance study of natural
radiometric contrasts between surficially exposed
lithologic units in the study area, the lower member of
Coch’s Windsor Formation (where exposed near the Suf-
folk scarp) was found to be anomalously radioactive
with respect to adjoining lithologic units. This radio-
metric contrast cannot be found on the aeroradiometric
map, indicating that the resolution of the map with
respect to spatial extent is marginal in places.

Because of complicated cultural patterns in parts of
the study area, and because previous work has shown
that most anomalies in coastal plain areas are caused by
materials other than heavy minerals (Force and others,
1978), anomalies were classified prior to field investiga-
tion. For this purpose, land-use and land-cover maps
(USGS, 1977 a—c) provided an excellent first approxima-
tion of which anomalies would likely be caused by
heavy-mineral deposits and which would be caused by
cultural activities. The results of this approach are illus-
trated in plate 4. Land-use and land-cover maps are con-
structed from remote-sensor data, mostly U-2 photo-

 

graphs and satellite imagery. The resolution of these
maps is such that 10-acre (4-hectare) plots of land are
the smallest mappable units. Plate 4 represents the
intersection of the sets of radiometric signatures and
land-use and land-cover data for anomalous areas.
Analyses of numerical codes within anomalous areas
show that clusters of anomalies can be related to one or
more of eight types of land-use and land-cover settings.
The types are residential, commercial, industrial, agri-
cultural, forest, wetland, barren. and beach.

The area south of the James River can be character-
ized by two broad zones of different types of anomalies;
zone 1 includes residential, commercial, and industrial
types for the most part near urbanized areas, and zone 2
is agricultural, forest, wetland, and beach types for the
most part near the North Carolina border. Similarly, the
tip of the J ames-York Peninsula is characterized by pre-
dominantly residential and commercial types. A cluster
of anomalies near the southern bank of the Rappa-
hannock River is characterized as predominantly agri-
cultural and forested land. Anomalies on the Delmarva
Peninsula are almost exclusively of the agricultural
type; beach-type anomalies (not accessible in this study)
are on the Atlantic shore. The land-use map is more
than 6 years old, and, in some areas, major changes were
noted during field investigation.

Field investigation of the anomalous areas just de-
scribed proved the land-use and land-cover categoriza-
tion scheme applicable. Anomalies associated with
residential, commercial, industrial, and military sites
are invariably caused by radioactive material brought in
from outside the Coastal Plain by man. Such materials
are generally granitic road metal, facing stone, and
other hard rock. On the other hand, anomalies associ-
ated with agricultural, forest, wetland, barren, and
beach areas are in some instances explainable by con-
centrations of radioactive heavy minerals exposed at
the surface.

FIELD METHODS

Field investigation consisted of ground checks of
total-count aeroradiometric anomalies by two methods
and of sampling anomalous materials for laboratory
analyses. Geographic areas where the aeroradiometric
signature is greater than local background are outlined
on plate 4 and were also outlined on 7 1/2-minute
quadrangle topographic maps for the purposes of field
investigation.

First, an anomalous area was traversed by vehicle to
detennine the geographic extent of the anomaly; this
approach also verified that the anomaly registered by
the airborne system was real. During the vehicle

6 AERORADIOMETRIC MAPS IN EXPLORATION FOR HEAVY-MINERAL DEPOSITS

traverse, readings were taken over sediment at 0.1-mile
(0.16-kilometer) intervals by use of a total-count port-
able scintillometer to find the anomalous material. Sec-
ond, where this material was found, a four-channel spec-
tral scintillometer containing a largevolume (110 cubic
inches, 1,900 cubic centimeters) N aI detector was used
to measure the components of the gamma-radiation
field. To achieve constant geometry at each locality, the
detector unit of the instrument was suspended about
1.5 feet (0.5 meter) above the surface from a tripod.
After temperature equilibration and standardization
against a barium-133 gamma-ray source, the count rate
was measured at the following gamma-ray energies: (1)
2.62 MeV (million electron volts) from thallium-208 in
the thorium-232 series; (2) 1.76 MeV from bismuth-214
in the uranium-238 series; and (3) 1.46 MeV from
potassium-40. The counting time at each locality did not
exceed 6 minutes. The data are given in table 1. Sedi-
ment samples were taken immediately below the detec-
tor crystal by a soil auger to a depth of 3.3 feet (1 meter).
In the auger sample where dark minerals were observed
in quantities of 1 to 3 percent, more extensive sampling
was done.

Of 164 anomalous sites investigated (pl. 1), 75 were
sampled for laboratory analyses; the remainder were
not confirmed by field inspection or did not contain visi-
ble dark minerals. All aeroradiometrically anomalous
localities on the Delmarva Peninsula were estimated to
contain considerably less than 1 percent total heavy
minerals in the top 3.3 feet (1 meter) of sediment, and
consequently none was sampled for laboratory analyses.
The sample collected at loc. 20E was from 15 to 20 feet
(4.6 to 6.1 meters) depth.

EFFECTS OF FERTILIZER USE ON
AERORADIOMETRIC MEASUREMENTS

The most acute problem encountered during field in-
vestigation was related to land used for agricultural
purposes. Typical of this problem is an area north of the
North Carolina-Virginia line, where aeroradiometric
anomalies exceeding 250 counts per second could not be
located on the ground. The dominant lithology in this
area is a silty clay of the Sand Bridge Formation.2
Ground—radiometric characterization of this area (see A,
pl. 2) shows no consistently high values associated with
the aeroradiometric high. This lack of ground-
radiometric anomalies is typical of most agricultural
fields associated with aeroradiometric anomalies. In
these areas, corn, cash crops (for example, tomatoes,
strawberries), and “double crops” (for example, corn

”As mapped by Oaks and Coch (1963, 1973).

 

followed by grain) are the main products. The common
denominator of these crops is that they all need very
large amounts of mixed fertilizer frequently applied.
Mixtures commonly consist of variable amounts of
nitrate, potash, and phosphate either in liquid or
granular form. Potash and phosphate are radioactive.
Because both aeroradiometric surveys conducted near
the North Carolina-Virginia border registered high
count rates over these silty-clay deposits, we can
assume that, at the time of the surveys, anomalously
radioactive material was present. The inability of
ground surveys to prove these anomalies real poses a
problem for which no clear-cut explanation exists. Ex-
amination of flight recovery sheets and analog charts of
the northern part of the aeroradiometric survey shows
that the survey was conducted between October 1976
and February 1977, coinciding with fertilizer applica-
tion times. The most reasonable explanation involves
the temporal and geographic coincidence of fertilizer
application and of survey flights.

To test the assumption that fertilizer was responsible
for causing anomalies, gross fertilizer consumption by
counties (index map pl. 4) was compared with radio-
metrically anomalous areas in each county. As a result,
strong correlation was found to exist between the types
and quantities of fertilizer used in counties with
anomalies falling on agricultural fields, so that, for Acco-
mack County, Northampton County, Virginia Beach Ci-
ty (formerly Princess Anne County), and Chesapeake Ci-
ty, aeroradiometric anomalies on agricultural lands were
suspected to be caused by fertilizer. Subsequent field
and laboratory investigations strongly indicate the fer-
tilizer to be the cause of these anomalies. A possible
alternate explanation is that clayey material normally
contains potassium-40 in clay minerals such as mus-
covite, biotite, and illite and contains uranium-series
nuclides adsorbed on clay minerals; hence, areas where
clay is common should be anomalous with respect to
sandy terranes.

LABORATORY ANALYSES

Laboratory procedures were directed toward two
goals. First, I wanted to find the amount of economic
minerals in each sample. Second, because mined heavy
minerals are normally separated as coarse- to fine-
grained sand, by splitting the very fine grained sand-
and silt-size fraction away from the coarser sand and ex-
amining both size fractions independently for their
heavy-mineral content, insight would be provided as to
partitioning tendencies according to size of different
heavy minerals.

 

LABORATORY ANALYSES

TABLE 1.—Gmund-radiometric and aeromdiometric signatures of sample localities in the Coastal Pwin of Virginia

[Dash (—) means no data were collected. Sample localities are labeled on plate 1 by the sample numbers minus the prefix “AG"l

 

 

 

 

 

 

 

Ground Amuldio e’l‘h

radiometric metric eTh (ppmjl eU
Sample 3mm signatme Tom 144° Biz“ 11’“8 T K eU e’I'h (ppm)/ 60 (ppmV
number mu. Url (CPS)2 count count count count (min) (7..) (ppm) (ppm) K(%) (ppm) 10%)

Nufolk Fmfion’. shelf fine-sand fades
AG 12N —————————————— 10—12 17-20 220 23,780 1500 482 371 2 1.73 1.40 4.73 2.73 3.38 0.81
AG 14N —————————————— 4-5 7—8 250 20,500 1470 461 377 2 1.75 1.25 4.85 2.77 3.88 .71
AG 62N —————————————— — — 200 76.720 6289 2228 1864 6 1.68 2.32 8.48 5.05 3.66 1.38
AG 76N -------------- — — 250 — — — — — — —— — — — —
AG 77N —————————————— 7 12 250 — — —- — — — —— — — — —
AG 78N -------------- 6 10 340 — — — — — —- — — — — .—
mm and Holocene slimline sands
AG 13N -------------- 10 17 250 19.600 1290 366 247 2 1.59 1.03 2.86 1.80 2.78 0.65
AG 15N -------------- 6 10 195 23,500 1738 614 395 2 1.93 2.18 4.99 2.59 2.29 1.13
AG 48N -------------- - — 235 70,540 7388 1708 1367 6 2.43 1.70 5.98 2.46 3.52 .70
AG 50N —————————————— — — 210 78,530 6744 2240 2129 6 1.86 2.08 9.89 5.32 4.75 1.12
AG 61N -------------- 5 8 250 70.510 5010 2009 1608 6 1.23 2.10 7.20 5.85 3.43 1.71
AG 63N -------------- — — 214 73,230 5560 1990 1540 6 1.48 2.13 6.82 4.61 3.20 1.44
AG 72N —————————————— 5—6 8-10 200 75,800 3622 1910 1127 6 1.41 1.88 4.97 3.52 2.64 1.33
AG 80aN ————————————— 5—12 10-20 280 72.080 2708 2199 1434 6 .59 2.41 6.62 11.22 2.75 4.09
Sand Bridge Fmdm’. marsh and tidal-flat siltyclny fades

AG 16N —————————————— 8 13 280 — —- —- — — — — — — — —
AG 17N -------------- 6 10 290 25.980 2528 631 445 2 3.32 2.14 5.78 1.74 2.70 0.64
AG 18N —————————————— — — 250 22.050 1835 560 387 2 2.20 1.85 4.91 2.23 2.65 0 84
AG 19N —————————————— — - 250 — —— — — — — — — — —- -
AG 27N —————————————— 4—5 7-8 230 61.680 4900 1360 1152 6 1.48 1.17 4.97 3.36 4.25 .79
AG 28N —————————————— 4—5 7-8 250 66,870 5806 1563 1249 6 1.79 1.49 5.43 3.03 3.64 .83
AG 29N —————————————— — —— 250 66,500 5005 1526 1231 6 1.45 1.45 5.34 3.68 3.68 1.00
AG 30N —————————————— — — 260 67,660 5480 1555 1241 6 1.65 1.51 5.36 3.25 3.55 .92
AG 31aN ------------- 2-4 3-7 230 66,200 5297 1493 1209 6 1.60 1.40 5.22 3.26 3.73 .89
AG 32N -------------- — -— 240 73,110 6061 1813 1447 6 1.78 1.85 6.40 3.60 3.46 1.04
AG 33N -------------- — —- 250 64.700 5553 1542 1147 6 1.69 1.57 4.88 2.89 3.11 .93
AG 34N —————————————— —- — 160 56,210 4253 1266 887 6 1.25 1.25 3.56 2.85 2.85 1.00
AG 35N“ ------------- — — 300 76.930 7078 1941 1571 6 2.17 2.00 7.00 3.23 3.50 .92
AG 36N -------------- 6 10 275 69,040 6334 1722 1489 6 1.94 1.61 6.65 3.43 4.13 .83
AG 37N -------------- — — 225 75,650 5280 1724 1181 6 1.50 1.93 4.98 3.32 2.58 1.29
A'G 38N -------------- — — 255 82.220 7776 2181 1866 6 2.36 2.21 8.53 3.61 3.86 .94
AG 39N -------------- — — 265 81.060 6695 2062 1874 6 1.93 1.95 8.60 4.46 4.41 1.01
AG 40aN ------------- — — 235 66,740 5093 1728 1410 6 1.39 1.71 6.23 4.48 3.64 1.23
AG 41N —————————————— — —- 250 77,140 7849 1880 1670 6 2.53 1.77 7.57 2.99 4.28 .70
AG 42N -------------- — -— 240 69.210 6021 1716 1376 6 1.81 1.72 6.04 3.34 3.51 .95
AG 43N —————————————— — — 215 67.090 4586 1656 1342 6 1.20 1.62 5.87 4.89 3.62 1.35
AG 44N -------------- — — 210 74.590 6083 1917 1651 6 1.73 1.86 7.46 5.22 4.01 1.08
AG 45N -------------- 6 10 259 69.850 7306 1637 1382 6 2.42 1.53 6.10 2.52 3.99 .63
AG 46N -------------- — — 230 73,060 7424 1750 1618 6 2.40 1.54 7.32 3.05 4.75 .64
AG 47N —————————————— 8 13 258 74,740 5986 1728 1400 6 1.79 1.72 6.17 3.45 3.59 .96
AG 49N —————————————— 8 13 190 84,250 7856 2356 2222 6 2.29 2.23 10.37 4.53 4.65 .97
AG 58N —————————————— — — 175 53,500 2998 1214 875 6 .72 1.14 3.51 4.89 3.08 2.00
AG 59N —————————————— — — 275 58,710 3810 1523 1161 6 .93 1.51 4.95 5.32 3.28 1.62
AG 60N —————————————— — — 240 69,180 4976 1800 1416 6 1.31 1.85 6.24 4.76 3.37 1.41
AG 64aN ————————————— - — 275 87,150 7334 2072 1892 6 2.21 1.96 8.69 3.93 4.43 .89

See footnote at. end of table. p. B.

8 AERORADIOMETRIC MAPS IN EXPLORATION FOR HEAVY-MINERAL DEPOSITS

TABLE 1,—Gmund-radzbmetric and aemmdxbmetn'c signatures of sample localities in the Coastal Plain of Virginia—Continued

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

Ground Amndio e’I'h
radiometric metric eTh (ppmV eU
Sample 51mm signature Total K4o Biz“ ’11208 T K eU e’I‘h (ppm)! eU (ppmv
number uR/hr. Ur 1 (CPS? count count count. count (min) ('70) (ppm) (ppm) K(%) (ppm) K(%)
Sand dege Firm-fins, m and tidal-flat silty-clay {mice—Continued
AG 65N -------------- — — 260 72,470 6217 1886 1551 6 1.82 1.90 6.92 3.80 3.64 1.06
AG 66N —————————————— — —- 180 69,780 5192 1862 1597 6 1.37 1.80 7.19 5.25 3.99 1.31
AG 67N —————————————— — — 200 76,020 7381 2077 1869 6 2.23 1.97 8.56 3.84 4.34 .88
AG 68N -------------- 7—8 12—13 200 79,070 7503 1979 1954 6 2.32 1.69 9.85 4.25 5.83 .73
AG 69N —————————————— 6 10 250 — — — -— — — — —- — — ——
AG 70N —————————————— 6—7 10-12 250 —- — — — — — — —- — — —
AG 71N —————————————— 6—8 10-13 240 — — — — — — — — — _ _
AG 74N —————————————— 6—7 10-12 300 84,910 4244 2087 1455 6 1.81 1.61 6.87 3.80 4.27 89
AG 75N —————————————— 6—8 10-13 180 — — -— — — — — - — -— —
AG 79N —————————————— 6—7 10-12 260 — —— — — — — — — — — —
Sand Bridge anfiona. fluvial and lagoon silty-sand [Ides
AG 21N -------------- — — 75 69,580 5226 1626 1425 6 1.49 1.48 6.33 4.25 4.28 0.99
AG 22N —————————————— — —— 330 61.800 4264 1451 1226 6 1.15 1.29 5.33 4.63 4.13 1.12
AG 24N -------------- — —— 311 66,750 5176 1593 1349 6 1.49 1.48 5.93 3.98 4.01 .99
AG 25N —————————————— —— - 260 66,800 5570 1840 1509 6 1.55 1.84 6.71 4.33 3.65 1.19
AG 26N —————————————— — — 200 61.090 4155 1453 1211 6 1.11 1.32 5.23 4.71 3.96 1.19
AG 51N -------------- — — 275 70.280 4599 2106 1828 6 .98 2.10 8.33 8.50 3.97 2.14
AG 52N -------------- 6 10 180 70,760 5632 2173 1562 6 1.44 2.11 6.61 4.59 3.13 1.47
AG 53N -------------- — — 270 70,620 5084 2104 1789 6 1.20 2.12 8.14 6.78 3.84 1.77
AG 55N -------------- — — 275 65.440 4391 1522 1294 6 1.17 1.38 5.66 4.84 4.10 1.18
AG 56N -------------- — — 160 70.230 4876 1755 1352 6 1.28 1.82 5.90 4.61 3.24 1.42
AG 57N —————————————— — — 240 71,700 5027 2236 1771 6 1.12 2.43 8.00 7.14 3.29 2.17
UfitAbuflainlmdaspit—cmpkxfndes
AG 20E ———————————— . —— 6 10 275 92,970 4271 2517 1237 6 0.99 3.46 5.62 5.68 1.62 3.49
Holocene. fluvial and
AG 23N —————————————— 2-4 3—7 — — — —— — — — — — — — —
wwrm3.mmwmnatm
AG 54N —————————————— — - 240 73,250 6231 2053 1640 6 1.74 2.17 7.35 4.22 3.39 1.25
Sud mag. Emma. burial-smudge and mud-{Int [Ides
AG 73N —————————————— 5-6 8—10 215 74,990 3908 1752 1232 6 1.77 1.18 5.73 3.24 4.86 0.67
Ndolk Formntima, Inn-inc mm fndes
AG 81R —————————————— 6.5 11 290 86,470 3689 2266 1509 6 1.27 2.18 7.06 5.56 3.24 1.72
AG 82R -------------- 8 13 290 88,900 4278 2420 1531 6 1.55 2.41 7.15 4.61 2.97 1.55
Nufolk Fun-fins. fluvhl and estuarine clayey-sand fades
AG 83R4 ------------- 8 13 280 85.570 3976 2250 1550 6 1.50 1.99 7.33 4.89 3.68 1.33
AG 84R4 ------------- 8—9 13-15 275 91.110 3676 2390 1523 6 1.16 2.51 7.06 6.09 2.81 2.16
Undivided and and gravel (Inner Coast-l PHI)
AG 86R4 ————————————— 6-8 10-13 230 95,900 4255 2782 1926 6 1.37 2.72 9.29 6.78 3.42 1.99
AG 87R —————————————— 7—8 12-13 230 96,210 4374 2967 1837 6 1.30 3.35 8.66 6.66 2.59 2.58
AG 88R —————————————— 7-8 12—13 215 93.750 4907 2460 1482 6 2.01 2.41 6.88 3.42 2.85 1.20
AG 89R -------------- — — 225 84.510 3100 2566 1692 6 .64 2.86 7.96 12.44 2.78 4.47
UndividodsnndlndgraveHOumCoumlPhin)
AG 91R —————————————— 4—7 7-12 225 50,070 2441 1321 959 6 0.95 1.15 4.76 5.01 4.14 1.21
AG 90R —————————————— 8 13 283 94,160 4378 2735 1546 6 1.39 3.20 6.96 5.01 2.18 2.30

 

llur=0.6)LR/hr. A geological source of lux- premiums the same instalment. mspmse as
an identical scum cmtaining ally 1 part per minim uranium in radioactive equilbrium

(Intemational Atomic Energy Agency, 1976, p. 16).
7CPS=oounts per secmd.

3A5 mapped by Oaks and Coch (1963, 1973).
4No sediment sample collected.

1]

‘4

 

 

 

LABORATORY ANALYSES “ 9

TECHNIQUES

From 80 samples collected at aeroradiometrically
anomalous localities (except sample AG 23N), approx-
imately 14 ounces (400 grams) was split and dry-sieved
into gravel (>18 mesh); coarse- to fine-grained sand
(<18 mesh to >120 mesh, henceforth abbreviated
CFS); very fine grained sand to coarse silt (< 120 mesh
to >320 mesh, abbreviated as VFSS); and silt and clay
(<320 mesh) fractions. The CFS and VFSS fractions
were processed for their heavy-mineral content in
bromoform (s.g. 2.85). Samples of these two fractions
that contained more than 1 percent total heavy minerals
were studied to evaluate their economic value. First
separated into groups of four magnetic fractions on a
magnetic separator (hand-magnetic, 0.0 to 0.5-ampere,
0.5 to 1.0-ampere, and 1.0-ampere fractions), both CFS
and VFSS fractions were studied independently under
petrographic and binocular microscopes. Some opaque
minerals were identified by X-ray techniques. Amounts
of given mineral species were summed from each
magnetic and size fraction in which they occurred, and
their percentage of the whole heavy-mineral fraction
was calculated. Density was not compensated for.

RESULTS

Of the 80 samples collected from aeroradiometrically
anomalous localities, 65 samples were found to contain
less than 1 percent heavy minerals separated in bromo-
form (fig. 2). Table 2 shows the sieve analyses, total
heavy-mineral contents, and the heavy-mineral contents
of each sieve fraction analyzed. In more than half the
samples, the VFSS fraction contains a larger percentage
of heavy minerals than the CFS fraction of a sample.
Because the total heavy-mineral content of samples
listed in table 2 is so low, no qualitative studies were
made on their mineralogy, as they were considered to be
of no economic value.

Detailed mineralogical study of the 15 samples that
contained more than 1 percent total heavy minerals is
shown in table 3. Ten of the fifteen samples containing
more than 1 percent heavy minerals are from Oaks and
Coch’s Sand Bridge Formation. Sample AG 23N was
collected from a Holocene fluvial sand deposit that had
no aeroradiometric expression. As in the 65 samples
containing less than 1 percent heavy minerals, but more
consistently, the VFSS fraction was found to contain a
higher percentage of minerals than the CFS fraction of
the same sample. Whereas in the 65 samples, the
average ratio of weight percent VFSS-fraction heavy
minerals to the CFS-fraction heavy minerals per sample
is approximately 2.7 in a range of 0.0 to 25.0, in the 15
samples containing more than 1 percent total heavy

 

 

30

20

NUMBER OF SAMPLES

   

0.01 0.25 0.50 0.75 1.0 2.0
to to to to to to
0.24 0.49 0.74 0.99 1.99 12.0

WEIGHT PERCENT OF HEAVY MINERALS
HAVING SPECIFIC GRAVITIES >2.85

FIGURE 2,—Histogram of heavy-mineral percentages in 80 samples
from aeroradiometrically anomalous localities in the Coastal Plain
of Virginia (from tables 2 and 3).

minerals, the average ratio is approximately 9.15 in a
range of 0.7 to 27.1. To see if and how heavy minerals
partition according to grain size, mineralogical analyses
were plotted on circular percentage diagrams (pl. 5). For
each of the 15 samples analyzed, the total heavy-mineral
suite and the heavy-mineral suites of the CFS and VFSS
fractions are plotted independently to emphasize size-
dependent mineral distribution.

Samples from Oaks and Coch’s Sand Bridge Forma-
tion, marsh and tidal-flat silty-clay facies (AG 18N, AG
19N, and AG 66N) have heavy-mineral suites that are
dominantly CFS. Furthermore, most of the economic
minerals in the sample are CFS; for example, ilmenite,
sillimanite, kyanite, rutile, leucoxene, monazite, and zir-
con. The heavy-mineral suites of these samples contain
less than 2 percent economic minerals (table 3).

An auger sample from Unit A on the Delmarva Penin-
sula (AG 20E) yielded approximately 1.5 percent total
heavy minerals in a 5-foot(1.5-meter) section from 15 to
20 feet (4.6 to 6.1 meters) beneath the surface. As it is
deeply buried, the concentration could not have caused
the aeroradiometric anomaly. In this sample, more
heavy minerals are in the VFSS fraction, but most of
the economic minerals are in the CFS fraction. Ihnenite,
monazite, zircon, sillimanite, and kyanite tend to occur
in the CFS fraction: however, garnet, rutile, and leuco-
xene are more common in the VFSS fraction. Of interest
in this sample is the virtual lack of ilmenite in the VFSS
fraction. Noneconomic minerals dominate the suite of
heavy minerals in this sample by a wide margin, and,
consequently, the economic value of this deposit is
minimal.

A sample collected on the bank of the Elizabeth River
(AG 23N) 0.25 mile (0.4 km) north of the Portsmouth
Naval Hospital represents modern fluvial sediment
deposited over granitic material placed in the 1950’s to

10

AERORADIOMETRIC MAPS IN EXPLORATION FOR HEAVY-MINERAL DEPOSITS

TABLE 2.—Sieue analyses and heavy-mineral contents of 65 samples that contain less than 1 percent total heavy minerals collected from

[Sample localities are labeled on plate 1 by the sample numbers minus the prefix “AG." UTM, universal transverse mercamr. Some sieve analyses by Dwight E. Wallace]

aemradlbmetrically anomalous localities in the Coastal Plain of Virginia

 

 

 

 

 

 

 

so. so.
7 yr Thickness S.G. >2.85 >235
Sample UTM minute sampled vaell CFSI2 VFssl3 saw Clay >285 in CFS in VFSS
number coordinates quadrangle (mam) mm (Wm) mm (“7%) (wt.%) m7.» (wt.%)
Norfolk Fu-mdian‘, shelf fine-sand fades
AG 12N ———————— 4057520N Fentmess ---------------- 0.75 17 62 11 10 0.42 0.37 1.75
390840E
AG 14N -------- 4057740N Fentress ---------------- .75 2 78 11 8 .36 .32 1.01
392030E
AG 62N -------- 4081700N Little Creek -------------- 1.0 13 57 16 13 .05 .09 .03
395180E
AG 76N ———————— 4056380N Deep Creek -------------- 1.0 .7 64 21 14 .10 .07 .27
385720E
AG 77N ———————— 4055680N Fentress ———————————————— 1.0 0 55 28 17 .13 .14 .18
389220E
AG 78N -------- 4046900N Lake Drummond SE ——————— 1.0 0 66 21 13 .19 .16 .39
387440E
mm alluding and
AG 13N ———————— 4056540N Fentress ---------------- 0.75 l 84 9 6 0.88 0.49 5.13
392550E /
AG 15N ———————— 4055980N Fentress ———————————————— .35 25 51 14 10.8 .25 .29 .77
391250E
AG 48N -------- 4072600N Kempsville —————————————— 1.0 1 81 12 3 .98 .49 4.77
394180E
AG 50N -------- 4071530N Kempsville -------------- 1.0 10 72 10 7 .28 .34 .36
393190E
AG 61N ———————— 4083380N Little Creek —————————————— 1.0 27 56 7 9 .01 .01 .07
394480E
AG 63N -------- 4084000N Little Creek -------------- 1.0 9 54 17 18 .01 .01 .03
397650E
AG 72N -------- 4063420N Pleasant Ridge ----------- 1.0 .05 62 7 9 .59 .45 3.18
409140E
and Mg. anfim‘. m and tidal-M silty-chy fades
AG 16N -------- 4056610N Faun-ass ———————————————— 0.35 1 61 22 16 0.07 0.06 0.17
398400E
AG 17N ———————— 4056050N Fentress ———————————————— .50 24 34 25 17 .04 .01 .13
398620E
AG 27N -------- 4074620N Norfolk South ------------ 1.0 .4 74 12 14 .35 .39 .54
387000E
AG 28N -------- 4074420N Norfolk South ------------ 1.0 1.0 72 16 11 .41 .51 .23
385860E
AG 29N -------- 4075770N Norfolk South ------------ 1.0 7.0 70 11 12 .25 .31 .28
387720E
AG 30N -------- 4074920N Norfolk South ------------ 1.0 0 74 14 12 .46 .46 .85
388070E
AG 31aN ------- 4073330N Norfolk South ———————————— 1.0 1.0 74 14 10 .69 .70 1.23
386330E
AG 31bN ——————— 4073330N Norfolk South ———————————— 1.0 .3 98 1.5 .6 .28 .21 5.26
386000E
AG 32N ———————— 4074460N Kempsville -------------- 1.0 .08 66 20 14 .30 .38 .25
380770E
AG 33N ———————— 4078300N Kempsville —————————————— 1.0 .5 70 16 13 .74 .93 .55
391940E
AG 34N ———————— 4080720N Kempsville -------------- 1.0 3.1 81 8.5 7 .39 .33 1.49
3:93on
AG 36N -------- 4078650N Kempsville —————————————— 1.0 7.0 66 15 12 .31 .39 .34
392150E

See footnote on p. 12.

A

AA

 

 

 

A

LABORATORY ANALYSES

11

TABLE 2.—Sieve analyses and heavy-mineral contents of 65 samples that contain less than 1 percent total heavy minerals collected from

aemradiometrically anomalous localities in the Coastal Plain of Virginia—Continued

 

 

 

 

 

S.G. S.G.
7 Vr mm S.G. >285 >235
Sample UTM minute sampled Gravel1 CF81 2 Vl-‘ss1 3 Silt & Clay >285 in CFS in vsss
number coordinates quadrangle (metenl (wt%) (wh%) (wt-7o) (wt%l {wt%l (wt%l (wt.%)
Sand Bridge Fmfian‘. mm and tidal-flat :3:de [Ides —Confinued
AG 37N ———————— 4077320N Kempsville —————————————— 1.0 0.3 56 27 16 0.17 0.25 0.10
392000E
AG 38N ———————— 4075440N Kempsville —————————————— 1.0 .4 42 41 17 .02 .05 .01
396580E
AG 39N -------- 4074250N Kempsville —————————————— 1.0 7.0 60 23 9 .08 .13 .02
396970E
AG 40aN ------- 4074900N Kempsville -------------- 1.0 14 59 11 15 .10 .13 .21
396090E
AG 40bN ——————— 4074900N Kempsville -------------- 1.0 13 86 .16 .19 .00
396090E
AG 41N ———————— 4076540N Kempsville -------------- 1.0 .2 58 27 14 .30 .36 .31
394320E
AG 42N ———————— 4077440N Kempsville -------------- 1.0 7 67 14 12 .30 .35 .47
395680E
AG 43N ———————— 4075500N Kempsville —————————————— 1.0 7 70 11 11 .11 .14 .12
397810E
AG 44N -------- 4075000N Kempsville —————————————— .75 1 55 28 15 .13 .17 .11
398100E
AG 45N ———————— 4072040N Kempsville --------------- 1.0 4.5 61 21 14 .33 .38 .46
397280E
AG 47N ———————— 4069200N Kempsville —————————————— 1.0 .03 50 37 12 .04 .07 .01
398840E
AG 49N ———————— 4071050N Kempsville —————————————— 1.0 17 80 2 .86 .68 13.74
395620E
AG 58N -------- 4086700N Norfolk North ------------ 1.0 2 82 7 9 .16 .19 .07
385450E
AG 59N ———————— 4088150N Norfolk North ———————————— 1.0 2.5 71 11 15 .46 .63 .05
387900E
AG 60N ———————— 4089380N Norfolk North ------------ 1.0 3.5 63 22 11 .31 .42 .21
386600E
AG 64aN ——————— 4076090N Princess Anne ------------ 1.0 0 52 30 18 .06 .10 .02
407400E
AG 65N ———————— 4075430N Princess Anne ———————————— 1.0 .1 60 22 18 .21 .31 .09
405880E
AG 67N -------- 4070700N Princess Anne ———————————— 1.0 .1 46 39 14 .34 .42 .37
405800E .
AG 68N ———————— 4071380N Princess Anne ------------ 1.0 2 52 29 17 .11 .18 .03
401550E
AG 69N ———————— 4068220N Princess Anne ------------ 1.0 0 56 28 16 .44 .45 .68
401400E
AG 70N -------- 4064220N Pleasant Ridge ----------- 1.0 .09 55 29 15 .04 .04 .05
406000E
AG 71N ———————— 4073580N Princess Anne ———————————— 1.0 0 48 31 20 .01 .01 .02
409140E
AG 74N ———————— 4054300N Pleasant Ridge ——————————— 1.0 .2 62 21 16 .02 .02 .03
400780E
AG 75N ———————— 4061370N Fenlress ———————————————— 1.0 0 44 39 16 .01 .02 .01
392250E
AG 79N -------- 4050920N Lake Drummond ————————— 1.0 0 52 33 15 .09 .05 .17
379180E
Sand Bridge Fan-amt fluvial and lagoon silty—sand fades
AG 21N ———————— 4078580N Norfolk South ———————————— 1.0 0.2 91 6 4 0.37 0.36 1.57
382890E

See footnotes on p. 12.

1 2 AERORADIOMETRIC MAPS IN EXPLORATION FOR HEAVY-MINERAL DEPOSITS

TABLE 2,—Sieue analyses and heavy-mineral contents of 65 samples that contain less than 1 percent total heavy minerals collected from
aemmdlbmetn'cally anomalous localities in the Coastal Plain of Virginia—Continued

 

 

 

 

 

 

 

 

 

 

 

 

so. so.
7 ‘/r- Thickness 8.6. > 2.35 > 2.85
Sample U'l‘M minute sampled Gravsll CFS12 vr-‘ss1 3 Silt & Clay > 2.85 in CFS in vrss
number coordinates quadrangle (meters) mm) 1wt.%) (wt.%) (wwol (wt.%) (“'70) (wt.%)
Sand Bridge Fmfion“. fluvial Ind llgoul siltyclay {Ides —Continued

AG 22N ———————— 4077700N Norfolk South ———————————— 0.75 2.0 88 4 6 0.48 0.42 2.60
381870E

AG 24N ———————— 4075250N Norfolk South ———————————— 1.0 .2 69 16 14 .12 .15 .1 1
383780E

AG 25N ———————— 4073810N Norfolk South ———————————— .75 7.0 74 12 6 .59 .31 2.87
382670E

AG 26N ———————— 4072970N Norfolk South ———————————— ‘ 1.0 1.0 78 10 11 .39 .46 .35
382140E

AG 55N ———————— 4078960N Bowers Hill —————————————— 1.0 0 72 18 10 .92 .55 3.01
373840E

Unit A. barrier island «- spit-complex fades

AG 20E5 ——————— 4204230N Chincoteag-ue W. —————————— 1.50 0.3 60 30 1 1 0.35 0.39 0.39

(0-5’) ——————————— 458720E

AG 20E5 ------- 4204230N Chincoteague W. ---------- 1.50 1 66 27 4 .79 .47 1.81

(5-10') ---------- 458720E

AG 20E5 ------- 4204230N Chincoteague W. —————————— 1.50 2 33 53 12 .63 .39 .95

(10-15’) ————————— 458720E

Nmfdk Fa'mafim‘, marine sandy-cw fades

AG 81R ———————— 4104200N Hampton ——————————————— 1.0 5 65 16 13 0.80 1.00 0.93
378030E

AG 82R ———————— 4100000N Hampton --------------- 1.0 2 62 25 , 10 .93 1.07 1.04
3821 10E

Undivided sand and gravel (Inner Coastal Plain)

AG 87R ———————— 4168050N Urbanna ———————————————— 1.0 1 59 26 14 0.22 0.29 0.19
358860E

AG 88R -------- 4192860N Dunnsville —————————————— 1.0 .02 67 19 13 .97 .98 1.67
341520E

AG 89R ———————— 4187640N Dunnsville -------------- 1.0 1 61 22 15 .38 .41 .59
342060E

Undivided sand and gravel (Outer Coastal Plain)

AG 90R -------- 4200100N Heathsville —————————————— 1.0 9 64 13 13 0.26 0.30 0.52
379130E

llncludes dayballs. 4As mapped by Oaks and Cocll (1963. 1903).

2Coarsetovsryfinessna 5SampleaAG20EwerecollectedfromOtol5feet(0to4.6m)belawthesuriace.

3Very fine sand and silt. ,

stop shore erosion. A grab sample of this material
yielded about 1.5 percent total heavy minerals, most of
which are in the CFS fraction; only about 30 percent of
the heavy minerals are of economic value. Rutile, leu-
coxene, monazite, and zircon are concentrated in the
CFS fraction, and sillimanite and kyanite are more com-
mon in the VFSS fraction.

Some samples collected from Oaks and Coch’s Sand
Bridge Formation, fluvial and lagoon silty-sand facies
(AG 51N, AG 52N, AG 53N, AG 56N, AG 57N), yielded
1.25 to 1.92 percent total heavy minerals. Samples 51,
52, and 56 have heavy-mineral suites that are dominant-
ly VFSS, and samples 53 and 57 have heavy-mineral
suites that are dominantly CFS. The distribution of
economic heavy minerals in these samples is variable.

Samples 51, 52, and 56 have slightly more economic
minerals in the VFSS size fraciton, and samples 53 and
57 contain more in the CFS fraction. For all the samples,
only 35 to 45 percent of the heavy-mineral suites are of
economic value. Because of the small total heavy-
mineral contents of these samples, economic considera-
tions are precluded.

A sample from Oaks and Coch’s Sand Bridge Forma-
tion, barrier-sand-ridge and mud-flat facies (AG 73N),
contains about 1.2 percent total heavy minerals, 32 per-
cent of which are of economic value. Most economic
heavy minerals in this sample are in the VFSS fraction.
Garnet is the single most abundant mineral present,
and the concentration of titanium dioxide minerals is
very low.

 

SPECTRAL RADIOMETRIC CHARACTERIZATION OF ANOMALIES 13

Samples from a Holocene beach dune (AG 80aN) and
back dune flat (AG 80bN) yielded the largest total
heavy-mineral contents of all the samples collected in
the study area. Sample AG 80aN was collected from the
surface to a depth of 2.5 feet (0.75 meter). The total
heavy-mineral content of this sample is 5.1 percent, but
only about 30 percent of the total is of economic value.
The economically valuable heavy minerals are concen-
trated in the VFSS fraction. The overall economic value
of this deposit is diminished by the very small concen-
tration of titanium dioxide minerals. Sample AG 80bN
is a 33-foot (l-meter) channel sample from a dune face.
The total heavy-mineral content of this sample is 11.4
percent, but only about 30 percent of the total is of
economic value. Titanium dioxide minerals are more
abundant in this sample than in AG 80aN, particularly
in the VFSS fraction. In this sample, ilmenite, mona-
zite, and zircon are conspicuously more abundant in the
VFSS frac, but garnet, sillimanite, kyanite, rutile, and
leucoxene are more common in the CFS fraction. As in
sample AG 80aN, the economic heavy minerals are
mostly in the VFSS fraction. Because of the small size
of this deposit, and, perhaps more importantly, because
this deposit is in a national wildlife refuge, the economic
value of the deposit is considered marginal.

Sample AG 91R was collected from undifferentiated
sand and gravel (Outer Coastal Plain) in the northern
part of the study area near the Potomac River. This
sample yielded about 1.3 percent total heavy minerals,
of which about 65 percent are of economic value. The
total heavy-minerals suite is dominantly VFSS, and il-
menite is the single most abundant economic mineral
present. The very fine grained nature of the economic
heavy minerals and their small concentration limit the
economic value of this deposit.

Inasmuch as the heavy-mineral suites of the above-
described samples are dominated by noneconomic
minerals and contain only small percentages of
economic minerals, none of the areas sampled is consid-
ered to hold economic importance. Other factors that
diminish the economic importance of these deposits in-
clude the fine-grained nature of the host sediments,
small area and thickness of the deposits, and competing
land use.

SPECTRAL RADIOMETRIC CHARACTERIZATION
OF ANOMALIES

Previous studies on the applicability of spectral radio-
metric data to the exploration for heavy-mineral
deposits in coastal areas, particularly beach sands, have
shown that such deposits have radioelement spectra
having characteristic photopeaks where radioactive
heavy minerals such as monazite and zircon are in-

 

cluded in the heavy-mineral suite. The radioactive
elements are present either in the crystal lattices or as
inclusions in the stable heavy minerals, or both, and
therefore, secular equilibrium of daughter products with
the parent element is likely. Where an anomaly is not
caused by radioelement concentrations in resistate
heavy minerals, the assumption of equilibrium is prob-
ably not valid. Application of spectral aeroradiometric
data to the characterization of anomalies in the Charles-
ton, S.C., area (Force and others, 1978) showed that
heavy-mineral concentrations in that area have spectral
signatures in which either all photopeaks are anomalous
(that is, they show significant departures from K:U:Th
proportions typical of soils in the area), the Th peak is
anomalous and the U and K peaks are normal, or the Th
and U peaks are anomalous and the K peak is normal.
Robson and Sampath (1977) showed that the anomalous
radioelement of heavy-mineral concentrations in eastern
Australia is dominantly thorium. In a study of Atlantic
and Gulf Coast beach sands, Mahdavi (1964) showed
that thorium is the dominant radioactive element in-
volved in the radiometric signature of heavy-mineral
concentrations.

Under ideal circumstances, I expected the aeroradio-
metric anomalies caused by heavy-mineral concentra-
tions in the Coastal Plain of Virginia to have either high
U and Th values and minimal K values, or to have high
values in all spectra. Because Oaks and Coch (1973,
table 3, p. 38) indicated in their study that in the clay-
sized minerals, feldspar and illite or muscovite, were
present in amounts of trace to more than 20 percent and
trace to 35 percent, respectively, in post-Yorktown
Formation stratigraphic units, K spectral values were
expected to be fairly high for sediments in southeastern
Virginia. Force and Geraci (1975) found that the radio-
active heavy-mineral concentration in the most favor-
able lithology for economic heavy-mineral accumula-
tions (Pleistocene? shoreline sands) was low or absent.
This fact is borne out by the observation that none of
the deposits sampled in that study shows radiometric
(airborne or ground) contrast with the surrounding
lithology.

Knowledge of the low regional radioactive heavy-
mineral content of sediments coupled with information
on fertilizer applications in the study area strongly sug-
gested that radiometric contrasts, particularly as seen
by the spectrometer, may not be diagnostic of heavy-
mineral accumulations.

Ground-radiometric and aeroradiometric signatures
of sample localities are presented on table 1 grouped
into lithologic types. Attempts at graphic correlation of
aerial and ground radiation measurements failed to
show distinct patterns. The reason for this lack of cor-re
lation most likely lies in the fact that aeroradiometric

14 AERORADIOMETRIC MAPS IN EXPLORATION FOR HEAVY-MINERAL DEPOSITS

TABLE 3.—Sieve and heavy-mineral analyses of 15 samples that contain more than 1 percent total

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

S. G. S. G.
Thickness Silt S. G. > 2.85 > 2.85
Sample UTM 7 Vrminute sampled Gravel crs3 VI-‘ss4 and clay >235 in crs in vrss
number coordinate quadrangle (mam) (wt. ‘70) (wt. %) (wt. ‘70) (wt. ‘70) (wt. %) (wt %l (wt. %)
Sand Bridge Firm-limb. march
AG 18N ————— 4057590 N Fentress ------------- 0.75 21.1 58.8 9.6 9.6 1.32 1.53 4.34
395000 E
AG 19N ————— 4057590 N Fentress ————————————— 1.0 .8 87.0 4.5 7.3 1.74 1.35 12.72
395000 E
AG 66N ----- 4070150 N Princess Anne ———————— 1.0 0 78.7 10.7 10.4 2.44 2.83 2.00
407450 E
Unit A. bottle island
AG 20E6 -——— 4204230 N Chinooteague W. ------ 1.5 6.2 69.3 20.4 3.9 1.38 0.96 3.52
458720 E
Holocene. fluvial
AG 23N ————— 4079480 N Norfolk S. ------------ GRAB 8.4 89.5 1.4 0.3 1.43 1.24 23.98
386600 E
Sand Bridge Fmfion5. fluvill
AG 51N ----- 4079975 N Bowers Hill ---------- 1.0 0 77.4 15.2 7.0 1.92 1.01 7.50
368810 E
AG 52N ----- 4078585 N Bowers Hill —————————— 1.0 .1 75.8 14.5 8.9 1.25 .67 5.12
367580 E v
AG 53N ----- 4081050 N Bowers Hill ---------- 1.0 .1 78.4 13.0 7.9 1.40 .99 4.79
368675 E
AG 56N ————— 4083470 N Newport News S. ------ 1.0 .03 71.6 19.4 8.3 1.70 1.11 4.67
370040 E
AG 57N ————— 4083030 N Newport News S. —————— 1.0 .04 77.3 14.2 7.8 1.72 1.43 4.31
371570 E
Sand Bridge Emmi.
AG 54N ----- 4077980 N Bowers Hill ---------- 1.0 0 63.8 24.4 11.3 1.50 0.88 3.83
373780 E
. Sllld Bridge Fmtims. hm
AG 73N ————— 4058350 N Pleasant Ridge ——————— 1.0 0.2 70.8 12.9 15.4 1.18 0.67 5.42
- 408340 E
Hdoeae. slice
AG 80aN ———— 4048380 N Knotts Island --------- 0.75 0 93.7 6.0 0.08 5.10 2.29 49.65
421250 E
AG 80bN ———- 4048380 N Knotts Island -------- 1.0 .2 91.3 8.1 .2 11.40 5.70 76.96
421520 E '
Undivlded sand Ind
AG 91R ----- 4197090 N Burgess ------------- 0.75 6.5 74.4 17.5 0.9 1.26 0.23 6.24
381300 E
1SeamleAG23Nwasnotedlectedh'ananarea-sdiomstricallymamalwslocality. aCoal-setofinesand.
2Includes subadinnte pyroxene. ‘Vayfinesandandsil'.

lation most likely lies in the fact that aeroradiometric
values are extrapolated, as discussed above. Spectral
radiometric characterization (see section on “Field-
Spectrometer-Data Reduction”) of different lithologies
(table 4) in an attempt to distinguish lithologic type as a
function of spectral characteristics also yielded incon-
clusive results. The average potassium content of sedi-
ments in the study area is about 1.5 percent; thorium
and uranium are about 6.5 and 2 parts per million (ppm),
respectively. Averages calculated for localities yielding
samples that contain more, than 1 percent total heavy

 

minerals are about 1 percent potassium, 2 ppm ura-

nium, and about 7 ppm thorium. These results indicate
that, in general, anomalies caused by heavy-mineral
concentrations in the Coastal Plain area of Virginia tend
to have lower than the regional average potassium con-
tent, slightly lower than the regional average uranium
content, and slightly higher than the regional average
thorium content (table 4). However, the statistical preci-
sion (i 10 percent) and absolute accuracy (:20 percent)
of the spectral radiometric data may preclude the ab-
solute certainty of these observations.

More distinct differences in spectral signatures of
heavy-mineral-bearing samples can be defined if the

A‘

 

)r

SPECTRAL RADIOMETRIC CHARACTERIZATION OF ANOMALIES 15

 

heavy minerals collected from aemradbmetn‘cally anomalous1 localities in the Coastal Plain of Virginia

 

Weight percent of fractim having S.G. >285
P=>0.1 parent; —=none determined

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

Alteredt?) . Sllimanite
Magnitite ilmenite Epidote Amphibole‘ Garnet Tourmaline Stmrolite and Kaynite leueoxene Rutile Zircone Monazite Sphene Apafim
and tidal flat siltyday fades
0.2 50.2 7.7 4.6 4.7 0.5 l 1.4 1 1.6 2.3 2.6 4.0 0.5 P —
.4 53.6 5.9 8.3 1.6 .5 7.1 7.9 3.3 2.9 7.2 1.3 0.05 —
.2 54.5 1 1.1 5.6 .5 .34 3.4 16.2 1.3 1.5 3.9 .8 — .67
a spit ample: his: ,
0.1 16.4 5.9 34.8 10.7 0.9 13.2 10.5 0.7 1.9 3.8 1.1 — —
and fades ‘ "
3.8 27.0 11.0 28.6 0.8 0.05 7.7 7.7 4.3 3.1 5.6 0.5 -— —
and bacon silty-sand fades _7 :77! i
0.1 45.0 1 1.3 13.5 0.9 0.9 3.2 9.9 10.6 1.4 2.0 1.4 0.1 P
.13 23.7 9.2 24.2 .8 .8 3.3 23.3 7.1 2.5 4.2 .42 — —
.1 32.1 12.4 24.5 P .3 2.4 17.6 3.4 2.1 3.1 .3 —— —
.32 36.4 15.6 6.4 P .6 2.6 22.7 7.0 2.2 5.4 .64 — .6
.3 50.6 8.2 5.6 .6 .1 6.5 12.4 6.5 .9 8.2 6 — 3
estuarine and W fades 7
P 36.3 18.7 10.7 0.3 P 1.7 13.8 10.7 3.1 3.8 1.4 — _.
I-d-rldge and mud-flat fades
0.1 16.8 7.5 23.7 22.5 0.6 11.6 8.7 2.3 2.9 4.0 0.6 -- 0.1
line and .. 7
0.04 14.4 15.7 29.8 5.0 2.7 9.3 8.4 5.0 2.4 5.9 1.2 — —
.1 33.5 3.8 27.8 4.8 1.8 6.0 9.7 6.6 1.6 3.0 1.5 — —
gravel (Outs Coat-l PHI) , ,
0.1 58.4 1.8 1.4 P 2.1 5.3 2.5 18.5 3.6 5.3 1.1 — —

 

5A3 mapped by om and Coch (1963. 1363]:
oSampleAG 20E wascollectedfi'an l5to20feetl4.6wtalmeters)belowthesurface.

spectral signatures of mineralized samples (>1.0 per-
cent total heavy minerals) are compared with the spec-
tral signatures of nonmineralized (< 1.0 percent total
heavy minerals) samples within the same lithology. For
example, sample AG 80aN contains much less potas-
sium than, slightly more uranium than, and a similar
proportion of thorium as other samples from the same
or similar lithologies (table 1). Distinctive about this
sample, however, are the high thorium to potassium
ratio and the high uranium to potassium ratio with re
spect to those of the other samples of the same lith-
ology. Similarly, samples from Oaks and Coch’s Sand

 

Bridge Formation, fluvial and lagoon silty-sand facies,
also show a dichotomy of spectral characteristics.
Samples that contain more than 1 percent total heavy
minerals have smaller potassium contents and higher
uranium and thorium contents than do other samples
from the same lithology that contain less than 1 percent
total heavy minerals.

These results, then, indicate that heavy-mineral con-
centrations can be distinguished in a particular
lithology by their spectral radiometric signatures by
virtue of the fact that uranium and, more importantly,
thorium are the dominant radioactive elements and

16 AERORADIOMETRIC MAPS IN EXPLORATION FOR HEAVY-MINERAL DEPOSITS

TABLE 4.—Averages of equivalent thorium, equivalent uranium, and percent potassium in different lithologies measured in the field

 

 

Numbn- of e’l‘h (ppmll e’l’h (ppm)! eU (ppm)!

Lithologyr 7 samples K(%) eU (ppm) e’l'h (ppm) K(%) eU (ppm) K(%)

Norfolk Formationl, shelf fine-sand facies. ----- 3 1.72 1.66 6.02 3.52 3.64 0.97

Pleistocene and Holocene shoreline sands. ----- 8 1.57 1.94 6.17 4.67 3.17 1.52

Sand Bridge Formationl, marsh and tidal-flat 33 1.80 1.71 6.47 3.70 3.77 1.00
silty-clay facies.

Sand Bridge Formationl, fluvial and lagoon 11 1.27 1.76 6.56 5.30 3.78 1.42
silty‘sand facies.

Unit A, barrier island or spit-complex facies. -—-— 1 0.99 3.46 5.62 5.68 1.62 3.49

Sand Bridge Formationl. estuarine and 1 1.74 2.17 7.35 I 4.22 3.39 1.25
tidal-channel facies.

Sand Bridge Formationl, barrier-sand-ridge 1 1.77 1.18 5.73 3.24 4.86 0.67
and mud-flat facies.

Norfolk Formationl, marine sandy-clay fades. — 2 1.41 2.29 7.10 5.09 3.10 1.64

Norfolk Formationl, fluvial and estuarine 2 1.33 2.25 7.20 5.49 3.25 1.75
clayey-sand facies.

Undivided sand and gravel (Inner Coastal 4 1.33 2.84 8.20 7.33 2.91 2.56
Plain).

Undivided sand and gravel (Outer Coastal 2 1.17 2.18 5.86 5.01 3.16 1.76
Plain).

Average values for the study area. ——————————— 1.46 2.13 6.57 4.84 3.33 1.64

Average values of localities from which 1.33 1.92 6.69 5.71 3.57 1.67

samples yielded >1.0 percent total heavy
minerals (11 locations).

 

lAs mapped by Oaks and Coch (1963. 1973).

potassium is less abundant where radioactive heavy
minerals are present in the heavy-mineral suite.

SUMMARY AND CONCLUSIONS

Prospecting for placer heavy-mineral deposits by air-
borne and ground radiometric methods is severely
limited in efficiency by a number of factors, but none
precludes its usefulness. Even if all instruments func-
tion properly during a survey and the resolution of the
gathered data is good, the exploration targets must be
exposed at or within inches of the surface, or the radio-
metric response is attenuated by overburden. Other fac-
tors that limit the radiometric response of a heavy-
mineral deposit are ( 1) moisture content of the sediment
because water attenuates the gamma-ray flux to the
surface (and may force the radon precursor of Pb’“ and
Bi“ to deeper levels); (2) lack of or small concentration
of radioactive minerals in the heavy-mineral suite; (3)
fertilizer applications that mask the signature of a

 

deposit by increasing the regional background of non-

mineralized sediments; (4) dominance of radioactivity in
the clay-sized minerals in heavy-mineral-bearing sedi-
ments; and (5) thick vegetation cover or cultural over-
prints, such as roads made of granitic rock.

My experience in the Coastal Plain of Virginia has
shown that by very rigorous application of supple

.mentary data, such as land-use and land-cover maps,

fertilizer-use data. detailed geologic maps, and ground
spectral radiometric data, most anomalies on an aero-
radiometric map can be attributed to sources other than
heavy-mineral deposits prior to field investigation.

This study, as well as previous studies in the area,
shows that (no currently economic heavy-mineral
deposits are at or near the surface in the coastal Plain of
Virginia The deposits discussed in this and previous
reports are not considered economic because of the very
fine grained nature of the host sediments and the low
percentage of economic heavy minerals. These results
do not preclude the possibility of economically valuable
deposits at depth, as the lithology on the Delmarva
Peninsula is such that it may contain economic
deposits. Furthermore, inasmuch as the James River

I‘.‘{

1'

‘*

‘Igp‘rv

I‘Iv‘ Q

r?

FIELD-SPECTROMETER—DATA REDUCTION 17

drains an ilmenite-rich terrane (Minard and others,

1976), economic mineral deposits could possibly form in .
the offshore environment.

The aerial surveys used in this study were conducted
to aid exploration efforts in general, and, for that pur-
pose, they are well suited. If, however, aeroradiometric
surveys were to be conducted specifically for economic
placer deposits on coastal plain areas, a preliminary
ground reconnaissance should be performed to ascertain
radiometric contrasts, and the subsequent aerial survey
should be appropriately tailored to the geology.

FIELD-SPECTROMETER—DATA REDUCTION3
By KENNETH L. KOSANKE‘

At the completion of field work, the four-channel spec-
trometer was calibrated at the US. Department of
Energy’s facility at Walker Field in Grand Junction,
Colo. (Stromswold and Kosanke, 1978). The calibration
results indicated that the instrument was not function-
ing properly. The calibration parameters correcting the
uranium window for potassium and the thorium window
for uranium were both very much too large. A check of
the instrument revealed that its energy gain had in-
creased by approximately 6 percent relative to the K, U,
and Th windows (see fig. 3). This increase in gain was un-
fortunate because it meant that the gamma peaks from
K, U, and Th were falling about half outside their ap-
propriate windoWs and that the peaks from K and U
were falling partly within the windows for U and Th, re-
spectively. Field data were checked using these (pres-
ently con'ect) calibration parameters as well as calibra-
tion parameters appropriate for a typical spectrometer of
this type. The instrument’s condition was found to have
drifted systematically with time, and, thus, neither set of
calibration constants was satisfactory.

In order to salvage the field data, the performance of
the spectrometer had to be reconstructed as a flmction
of time during the period in which field data had been
collected. Fortunately, an accurate log of instrument
settings had been kept. Figure 4 is a plot of spectrom-
eter gain settings as an approximate function of time
during the period of field-data collection. During the
period of field-data collection, the instrument failed and
was repaired; this repair is indicated in figure 4 by a
dashed line. Gain settings before and after the malfunc-
tion are 5.9 and 7.7, respectively (corresponding to a
gain drift of about 4 percent). I assumed that this

’nfispmtofmewmk“msqumfledbyBemfixFfifliEnghaumg

”Baud!FhfldEnghueflnnglpuafion,GrundJundfllL(X)8150L

 

COUNTS, IN ARBITRARY UNITS

  

 

 

 

 

5L
§§

F.

1291

GAMMA-RAY ENERGY, IN KILOELECTRON VOLTS

FIGURE 3.—Gamma-ray spectrum of a bulk source of K, U, and Th,
showing the locations of improperly positioned spectrometer
windows.

change in proper gain setting was the result of a step-
wise deterioration of the instrument. Accordingly, the
field data were treated as two separate groups (before
and after repair) that had two sets of calibration con-
stants. Additional confirmation of a. change can be seen
in background measurements made during the field-
data collection period (see table 5).

The above information, coupled with data that had
been collected during the field operation, was sufficient
to allow a fairly accurate estimation of the two sets of
calibration constants. Calibration constants were
calculated by use of (1) results from spectral data from a
large pile of potash fertilizer that had been sampled and
could be used as a calibration standard; (2) data col-
lected at the same site before and after the instrument
change; and (3) computer simulation of the effects
caused by improper positioning of the K. U, and Th
spectral windows.

TABLE 5.—FieId-spectrometer background data before and after a
change in gain setting coinciding with a repair of the instrument.

 

 

 

Vthknvcnuntspernfinuts
K U Th
Before repair ---------- 74 58 27
After repair —————————— 49 48 30

 

 

18 AERORADIOMETRIC MAPS IN EXPLORATION FOR HEAVY-MINERAL DEPOSITS

 

 

I
I
|
| I
' 1
8— I.
7.7I—-——1————
o I o
| 0
Z I o
t: 7— I —
um: I
I
Z . 0 I
< I
0 6_ o ' I _
. 4i5.9
o i 0'
I
. I
o
5~ I —
I

 

 

 

TIME (NOT TO SCALE)

FIGURE 4.—Plot of sequence of spectrometer gain settings during the
field-data collection period. The elapsed time between gain settings
represented by points ranged from a day to a month. The time of in-
strument repair is shown as a dashed line. Horizontal lines repre-
sent average values.

The two sets of calibration constants are presented in
table 6 as calibration matrices. The calibration matrix is
a matrix of coefficients, each coefficient relating the
counting rate in a gamma-ray energy window to the con-
centrations of K, U, or Th present in the material being
evaluated. The matrix is derived from measurements
made over calibration “pads” consisting of background,
enriched-K, enriched-U, and enriched-Th concentrations
of the radionuclides in concrete.

TABLE 6.—Calibmt1bn matrices (A) for the spectrometer before and
after change in performance.

 

Calibratim matrix (fa rates in counts per minute)

 

 

Before After
275 75 20 255 74 13
0 80 15 27 72 20
0 5 32 0 6 30

 

SELECTED REFERENCES

Adams, J. A. S., Osmond. J. K., and Rogers. J. J. W., 1959, The geo-
chemistry of thorium and uranium, in Ahrens, L. H., and others,
eds., Physics and chemistry of the Earth: New York, Pergamon
Press, v. 3. p. 298-348.

Adams. J. A. S., and Weaver, C. E., 1958, Thorium-touranium ratios
as indicators of sedimentary processes—Example of concept of
geochemical facies: American Association of Petroleum Geolo-
gists Bulletin, v. 42, no. 2, p. 387-430.

Beck, H. L., 1972, The physics of environmental gamma radiation
fields, in Adams, J. A. S., Lowder, W. M., and Gesell, T. F., eds.,
The natural radiation environment II—Intemational Symposium

 

on the Natural Radiation Environment, 2d, Houston, Texas,
Aug. 7-11, 1972: US. Energy Rmarch and Development Ad-
ministration [Report] CONF 720805—P1, v. 1, p. 101—135.

Bick, K. F., and Coch, N. K., 1969. Geology of the Williamsburg,
Hog Island, and Bacons Castle quadrangles. Virginia: Virginia
Division of Mineral Resources Report of Investigations 18, 28 p.

Bunker, C. M., and Bush, C. A., 1966. Uranium, thorium, and radium
analyses by gamma-ray spectrometry (0184-0352 million elec-
tron volts): US. Geological Survey Professional Paper 550-3,
p. 3176—3181.

1967, A comparison of potassium analyses by gamma-ray spec-
trometry and other techniques: U.S. Geologcal Survey Profes-
sional Paper 575-B, p. 8164-8169.

Cazeau, C. J., 1974, Heavy minerals of Quaternary sands in South
Carolina, in Oaks. R. Q., Jr., and DuBar, J. R., eds., Post-Miocene
stratigraphy, central and southern Atlantic Coastal Plain: Logan,
Utah State University Press, p. 174-178.

Clark, S. P., Jr., Peterman, Z. E., and Heier, K. S., 1966, Abundances
of uranium, thorium, and potassium, in Clark, S. P., Jr., ed.,
Handbook of physical constants (revised edition): Geological
Society of America Memoir 97, p. 521-541.

Clark, W. B., and Miller, B. L., 1906, A brief summary of the geology
of the Virginia Coastal Plain, in Ries, Heinrich, The clay deposits
of the Virginia Coastal Plain: Virginia Geological Survey Bulletin
2. p. 11-24.

__1912, The physiography and geology of the Coastal Plain prov-
ince of Virginia: Virginia Geological Survey Bulletin 4, 274 p.

Coch, N. K., 1965, Post-Miocene stratigraphy and morphology, Inner
Coastal Plain, southeastem Virginia: US. Office of Naval Re
search. Geography Branch, Contract NONR 609 (40), Task Order
NR 388—064, Technical Report 6, 97 p. plus appendices. (Yale
University Ph.D. dissertation.)

__1968, Geology of the Benns Church, Smithfield, Windsor, and
Chuckatuck quadrangles, Virginia: Virginia Division of Mineral
Resources Report of Investigations 17, 39 p.

___1971, Geology of the Newport News South and Bowers Hill
quadrangles, Virginia: Virginia Division of Mineral Resources
Report of Investigations 28, 26 p.

Coch, N. K., and Oaks, R. Q., Jr., 1968, “Terrace-formation" concept
in Atlantic Coastal Plain stratigraphy [abs]: Geological Society
of America Special Paper 101, p. 253—254.

Duval, J. S., Schulz. K. A., and Pitkin, J. A., 1976, Preliminary
calibration of high-sensitivity gamma-ray systems being used in
the national uranium resource survey [abs]: Geologcal Society of
America Abstracts with Programs, v. 8, no. 6, p. 848.

Flint, R. F., 1940, Pleistocene features of the Atlantic Coastal Plain:
American Joumal of Science, v. 238, no. 11, p. 757-787.

Force, E. R.. and Bose, S. K., 1977, Gamma aeroradioactivity map of
parts of Norfolk and Eastville quadrangles. outer Coastal Plain,
North Carolina and Virginia: US. Geological Survey Miscellane-
ous Field Studies Map MF-863, scale 1:250.000.

Force, E. R., and Geraci, P. J.. 1975, Map showing heavy minerals
in Pleistocene(?) shoreline sand bodies of southeastern Virginia:
US. Geological Survey Miscellaneous Field Studies Map
MF-7l8, scale 1250.000.

Force, E. R., Grosz, A. E., Zietz, Isidore, and Loferski, P. J., 1978,
Utility of aeroradioactivity maps in the vicinity of Charleston,
South Carolina: US. Geological Survey Open-File Report 78—587,
56 p.

Gregory, A. F., 1960, Geological interpretation of aeroradiometric
data: Canada Geological Survey Bulletin 66, 29 p.

Harrison, Wyman, 1962, Pleistocene record in the subsurface of the
Norfolk area, Virginia. in Harrison, W., ed.. Guidebook for field
trips [Virginia Academy of Science, Geology Section, 40th An-
nual Meeting, May 11-12, 1962, Norfolk, Va]: Norfolk, Va, Nor-
folk College of William and Mary, p. 35-61.

Hobbs, C. H., 1974, Shoreline situation report. Newport News,

I.

[t‘azga

‘4

d

la‘r

he

\

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, +

 

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Finis-

SELECTED REFERENCES

Virginia: Virginia Institute of Marine Science Special Report in
Applied Marine Science and Ocean Engineering, no. 55, 77 p.

_1975, Shoreline situation report, City of Hampton, Virginia:
Virginia Institute of Marine Science Special Report in Applied
Marine Science and Ocean Engineering, no. 76, 63 p.

Hurley, P. M., 1956, Direct radiometric measurement by gamma-ray
scintillation spectrometer: Geological Society of America
Bulletin, v. 67, no. 4, p. 395—411.

International Atomic Energy Agency, 1976, Radiometric reporting
methods and calibration in uranium exploration: International
Atomic Energy Agency Technical Report 174, 57 p.

Johnson, G. H., 1969, Guidebook to the geology of the York-James
Peninsula and south bank of the James River. College of William
and Mary, Department of Geology, Guidebook 1, 33 p.

__1972, Geoloy of the Yorktown, Poquoson West, and Po-
quoson East quadrangles, Virginia: Virginia Division of Mineral
Resources Report of Investigations 30, 57 p.

___1976, Geology of the Mulberry Island, Newport News North,
and Hampton quadrangles, Virginia: Virginia Division of Mineral
Resources Report of Investigations 41, 72 p.

Mahdavi, Azizeh, 1964, The thorium, uranium, and potassium con-
tents of Atlantic and Gulf Coast beach sands, in Adams, J. A. S.,
and Lowder, W. M., eds., The natural radiation environment:
Chicago, University of Chicago Press, p. 87 -114.

Markewicz, F. J., Parrillo, D. G., and Johnson, M. E., 1958, The
titanium sands of southern New Jersey: American Institute of
Mining Engineers, Society of Mining Engineers Preprint
No. 5818A5, 10 p.

McCauley, C. K., 1960. Exploration for heavy minerals on Hilton
Head Island, South Carolina: South Carolina Division of Geoloy
Bulletin 26, 13 p.

McLean, J. D., Jr., 1966, Miocene and Pleistocene Foraminifera and
Ostracoda of southeastern Virginia: Virginia Division of Mineral
Resources Report of Investigations 9, 202 p.

Mertie, J. B., Jr., 1975, Monazite placers in the Southeastern Atlantic
States: US. Geological Survey Bulletin 1390, 41 p.

Meuschke, J. L., 1955, Airborne radioactivity survey of the Edisto
Island area, Berkeley, Charleston, Colleton, and Dorchester
Counties, South Carolina: US. Geological Survey Geophysical
Investigations Map GP—123, scale 1:62,500.

Minard, J. P., Force, E. R., and Hayes, G. W., 1976, Alluvial ilmenite
placer deposits, central Virginia: US. Geological Survey Profes-
sional Paper 959—11, 15 p., 1 pl.

Moncure, Richard, and Nicole, Maynard, 1968, Characteristics of
sediments in the James River estuary, Virginia: Virginia Insti-
tute of Marine Science Special Scientific Report, no. 53, 40 p.

Moore, W. E.. 1956, Pleistocene terraces south of the James River,
Virginia: Virginia Academy of Science Geology Section Guide-
book, 7 p.

1956, Stratigraphy of Pleistocene terrace deposits in Virginia
[abs]: Geological Society of America Bulletin, v. 67, no. 12, pt. 2,
p. 1755.

Moxham, R. M., 1954, Airborne radioactivity survey in the Folkston
area, Charlton County, Georgia, and Nassau County, Florida:
US. Geological Survey Geophysical Investigations Map
GP—119, scale 1:62,500.

_1960, Airborne radioactivity surveys in geologic exploration:
Geophysics, v. 25, no. 2, p. 408—432.

Neiheisel, James, 1958a, Heavy mineral beach placers of the South
Carolina coast: South Carolina Division of Geology, Mineral In-
dustries Laboratory Monthly Bulletin, v. 2, no. 1, p. 1—7.

_1958b, Origin of the dune system on the Isle of Palms, South
Carolina: South Carolina Division of Geology, Mineral Industries
Laboratory Monthly Bulletin, v. 2, no. 7, p. 46—51.

_l976, Heavy minerals in aeroradioactive high areas of the
Savannah River flood plain and deltaic plain: South Carolina
Division of Geology Geologic Notes, v. 20, no. 2, p. 45-51.

 

 

19

Neuschel, S. K., 1970, Correlation of aeromagnetics and aeroradio-
activity with lithology in the Spotsylvania area, Virginia: Geo-
logical Society of America Bulletin, v. 81, no. 12, p. 3575—3582.

Neuschel, S. K., Bunker, C. M., and Bush, C. A., 1971, Correlation
of uranium, thorium, and potassium with aeroradioactivity in the
Berea area, Virginia: Economic Geology, v. 66, no. 2, p. 302—308.

Oaks, R. Q., Jr., 1964, Post-Miocene stratigraphy and morphology,
Outer Coastal Plain, southeastem Virginia: US. Office of Naval
Research, Geography Branch, Contract NONR 609(40), Task
Order NR 388—064, Technical Report 5. 240 p. plus appendices.
(Yale University Ph.D. dissertation.)

Oaks, R. Q., Jr., and Coch, N. K., 1963, Pleistocene sea levels, south-
eastern Virginia: Science, v. 140, no. 3570, p. 979—983.

__1968, Post-Miocene tectonics of southeastern Virginia [abs]:
Geological Society of America Special Paper 101, p. 272.

_l973, Post-Miocene stratigraphy and morphology, southeast-
ern Virginia: Virginia Division of Mineral Resources Bulletin 82,
135 p.

Oaks, R. Q., Jr., Coch, N. K., Sanders, J. E., and Flint, R. F.. 1974,
Post-Miocene shorelines and sea levels, southeastern Virginia, in
Oaks, R. Q., Jr., and DuBar, J. R., eds., Post-Miocene stratig-
raphy, central and southern Atlantic Coastal Plain: Logan, Utah
State University Press, p. 53—87.

Onuschak, Emil, Jr., 1973. Pleistocene-Holocene environmental
geology, in Geologic studies, Coastal Plain of Virginia: Virginia
Division of Mineral Resources Bulletin 83, pt. 3, p. 103—153.

Perlman, S. H., Bunker, C. M., Bush, C. A., and Gohn, G. S., 1976,
Ground radioactivity measurements in South Carolina and
Georgia with radioelement analyses of surficial sediments: US.
Geological Survey Open-File Report 76—478, 10 p.

Pitkin, J. A., Neuschel, S. K., and Bates, R. G., 1964, Aeroradioactiv-
ity surveys and geologic mapping, in Adams, J. A.’S., and
Lowder, W. M., eds., The natural radiation environment: Chicago,
University of Chicago Press, p. 723-736.

Roberts, J. K., 1932, The lower York-James Peninsula: Virginia
Geological Survey Bulletin 37. 58 p.

Robson, D. F., and Sampath, N., 1977, Geophysical response of
heavy-mineral sand deposits at Jerusalem Creek, New South
Wales: BMR [Australia Bureau of Mineral Resources] Journal of
Australian Geology and Geophysics, v. 2, no. 2, p. 149-154.

Sanders, J. E., Flint, R. F., and Oaks, R. Q., Jr., 1962a, Preliminary
report on late Pleistocene and recent littoral and nearshore
marine sediments, south of Cape Henry, Virginia [abs]: Virginia
Journal of Science, new series, v. 13, no. 4, p. 286.

1962b, Preliminary report on the geoloy of southeastern
Virginia and adjacent coast and continental shelf, with remarks
on sediment sampling technique using vibro-drilling methods:
New Haven, Conn., Yale University, Department of Geology,
Annual Report on Contract NONR 609(40) NR—388—064 to US.
Office of Naval Research, Geography Branch, 45 p.

Schmidt, R. G., 1962, Aeroradioactivity survey and areal geology of
the Savannah River plant area, South Carolina and Georgia
(ARMS-1): US. Atomic Energy Commission Report
CEX-58.4.2, 42 p.

Sinnot, Allen, and 'I‘ibbitts, G. 0., Jr., 1954, Summary of geology
and ground-water resources of the Eastern Shore peninsula,
Virginia: Virginia Division of Geology Mineral Resources Circular
2, 18 p.

Spencer, R. V., 1948, Titanium minerals in Trail Ridge, Florida:
US. Bureau of Mines Report of Investigations 4208, 21 p.

Stockman, K. W., Pickering, S. M., Jr., Higgins, M. W., and Zietz,
Isidore, 1976, Application of contoured aeroradioactivity maps
for heavy-mineral exploration in Georgia’s coastal area [abs]:
Geologcal Society of America Abstracts with Programs, v. 8,
no. 2, p. 277-278. ,

Stromswold, D. C., and Kosanke, K. L., 1978, Calibration and error
analysis for spectral radiation detectors: Institute of Electrical

 

20 AERORADIOMETRIC MAPS IN EXPIDRATION FOR HEAVY-MINERAL DEPOSITS

and Electronics Engineers Transactions on Nuclear Science. v.
NS—25. no. 1, p. 782—786.

Taylor, S. R.. 1964, Abundance of chemical elements in the conti-
nental crust—A new table: Geochimica et Cosmochimica Acta,
v. 28, no. 8, p. 1273-1285.

Thomas, J. B., and Goodwin, B. K., 1973. Heavy minerals of the
Continental Shelf off the eastern shore of Virginia: Geologcal
Societyof America, Abstracts withPrograms,v. 5. no. 5, p. 444.

US. Geological Survey. 1977a, Land use and land cover and asso
dated maps for Eastville, Maryland, Virginia, and North

Carolina: US. Geological Sirvey Open-File Repat 77-62, scale
11250.000.
_1977b, Land use and land cover and associated maps for Non-
V folk. Virginia. North Carolina.- US. Geological Survey Open-File
Report 77-789, scale 1:250,000.

_1977c, Land use and land cover and assoa'ated maps for
Richmond, Virginia. Maryland: US. Geological Survey Open-File
Report 77-85. scale 1:250.ooo.

Wentworth, C. K., 1930. Sand and gravel mources of the Coastal
Plain of Virginia: Virginia Geological Survey Bulletin 32, 146 p.

I.
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NATIONAL GEODETIC VERTICAL DATUM OF 1929

INTERIOR—GEOLOGICAL SURVEY, RESTON, VA '1583’682621

25LMILES

35 KILOMETERS

GENERALIZED GEOLOGIC MAP OF THE OUTER COASTAL PLAIN OF VIRGINIA

 

37°00’

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Index map showing sources of data for coastal Virginia

EXP

Oaks and Coch, 1973

 

Force and Geraci, 1975

 

 

 

Bick and Coch, 1969

 

 

 

Coch, 1968

 

 

 

 

 

 

 

 

 

LANATION

Modified from Oaks and Coch, 1973
and Force and Geraci, 1975

 

 

 

 

Johnson, 1972

 

 

Johnson, 1976

 

 

 

Coch, 1971

 

 

 

R. B. Mixon, unpublished data

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

PROFESSIONAL PAPER 1263
PLATE 1

EXPLANATION OF MAP UNITS

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

Défi

Unconformity

 

 

21s
I70
I/

 

 

 

 

 

Holocene

Pleistocene

 

 

QUATERNARY

DESCRIPTION OF MAP UNITS

Shoreline sands (HoloceneI—Includes sediments on foreshore, berm,

dunes, and backdune flat. Sediments of the foreshore consist of white to
light—brown, fine- to coarse~grained quartz sand with some fine pebble
gravel, small amounts of mica, and locally abundant shell fragments.
Dark minerals commonly outline stratification. Coarser grained sand
commonly increases with depth. Sediments of the berm are chiefly
pale-yellow to light—brown, fine—medium-grained quartz sand with few
thin coarse-grained sand layzrs. Dark minerals outline small wind—driven
ripples within beds of the berm. Sediments grade from wave deposited
at the seaward berm ridge to eolian over the landward part. Dune areas
are irregular and marked by numerous closed depressions up to 12 feet
deep. Dune crests are commonly 18 to 27 feet above sea level; large
solitary dunes up to an elevation of 40 feet and 0.2 to 0.4 miles in
diameter are present landward of the shore dunes south of False Cape.
Shore dunes consist of pale—yellow to light—brown, medium- to very
fine grained quartz sand with 1 to 3 percent silt and 1 to 5 percent mica.
Elsewhere, dune stratification consists of nearly horizontal laminae that
have low—angle pinchouts. Sediments of the backdune flat are dark
brown to pale yellowish-brown, fine— to medium—grained sand with
variable admixtures of very fine sand, coarse sand, silt, clay, and vascu—
lar plant fragments. The abundant mica probably is blown in from the
dunes together with silt. Thickness ranges from 1 to 20 feet.(Oaks and
Coch, 1973, p. 99)

Undivided sediments (Holocene)——Beach, marsh, swamp, and stream
sediments. Along portions of existing streams above tidewater, clean to
clayey, fine— to medium-grained sand and sandy silt are being de-
posited. Along portions of streams directly affected by tides, especially
the James River, deposits of fine— to medium—grained sand with some
fragments of oysters and fine pebble gravel are common. Thickness
ranges from 1 to 20 feet. (Oaks and Coch, 1973, p. 104)

UnitA (upper? PleistoceneI—Medium- to coarse—grained sand, gravelly in
part, commonly strongly trough crossbedded; interfingers and overlaps
westward with very fine to fine clayey and silty sand and clayey silt. A
barrier island or barrier spit complex characterized by Carolina Bays.
(Robert B. Mixon, written commun., 1979)

Unit B (upper Pleistocene)—Medium— to coarse-grained gravelly sand,
interbedded with lesser amounts of fine— to very fine grained sand,
commonly trough crossbedded. A barrier island or barrier spit complex.
(Robert B. Mixon, written commun., 1979)

Unit C (upper Pleistocene)—Fine— to medium—grained sand, clayey and
silty in part, massively to horizontally bedded, characterized by a few
thin laminae or beds of clayey silt. Probably represents sands of an
ancestral Chesapeake Bay bottom. (Robert B. Mixon, written commun.,
1979)

Unit D (D1) (Late Pleistocene)—Fine— to medium-grained sand, gravelly
near base. Deposited on very near shore shelf environment. Unit D is
equivalent to Unit D1 and is shown in darker color. (Robert B. Mixon,
written commun., 1979)

Unit E (Pleistocene)—Fine— to medium—grained sand, clayey and silty in
the north; moderately sorted to well sorted in the south. Some gravel
near the base and along scarps that bound the unit on the east. Deposits
of an ancestral Chesapeake Bay. (Robert B. Mixon, written commun.,
1979)

Unit F (Pleistocene)—Beach ridge complex. Ridges of generally medium—
to coarse—grained sands and fine gravel. (Robert B. Mixon, written
commun., 1979)

Sand Bridge Formation1 (Pleistocene)——Estuarine and tidal channel:
clayey-sand facies. Composition of the clayey-sand facies ranges from
clayey sand, silt, and clay, to well—sorted, fine— to medium—grained sand.
Low—angle, planar cross laminae occur in sandier parts of the facies,
together with plates and angular chunks of clay. The base of the clayey—
sand facies slopes toward present streams so that the minimum total
thickness of this facies, exposed above sea level, is 25 to 30 feet near
present streams. The clayey-sand facies is oxidized yellowish orange to
depths of 3 to 8 feet, below which yellowish green, light gray, or bluish
white predominate. (Oaks and Coch, 1973, p. 92—94)

Sand Bridge Formation1 (Pleistocene)—Fluvial and lagoonal silty—sand
facies. Consists of fine— to medium—grained sand with 20 to 30 percent
silt. It lacks obvious stratification, sedimentary structures, and fossils.
This facies is 5 to 8 feet thick, and its top lies entirely above 18 feet. The
silty-sand facies is oxidized only 2 to 3 feet below the surface. The
clay—sized fraction includes more than 20 percent vermiculite, 5 to 20
percent illite, 5 to 10 percent feldspar, and considerably more than 20
percent quartz. (Oaks and Coch, 1973, p. 94)

Sand Bridge Formation1 (Pleistocene)——-Marsh and tidal flat silty—clay
facies. Consists of approximately equal amounts of clay and silt with 10
to 15 percent fine— to medium—grained sand, especially near its base.
The silty—clay facies in many places has a blocky, massive texture with-

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

80N.
1E0

 

 

out stratification and is cohesive; locally, however, rather continuous
but irregular wavy beds to 0.3 foot thick are present. Unoxidized color
varies from dark gray to very light gray. Yellowish-brown limonitic
mottling is common in many places. The clay-sized fraction consists of
vermiculite, chlorite, illite, and feldspar in amounts less than 5 to 20
percent each, 5 to 30 percent montmorillonite, and 60 and 80 percent
quartz. (Oaks and Coch, 1973,‘p. 94)

Sand Bridge Formation1 (Pleistocene)——Barrier- sand-ridge and mud—flat

complex. Ridges are underlain by linear bodies of white to light-gray,
fine to very fine grained quartz sand with a trace of silt, and 3 to 6 percent
mica, which overlies compact, light-yellow to yellowish—brown, fine- to
coarse-grained quartz sand with 1 to 5 percent fine pebble gravel. The
upper sand has a characteristic dry, fluffy feel similar to that of modern
dune sand, and it interfingers with the silty-clay facies in the western part
of Pungo Ridge. Stratification in cores is low angle, 0.5 to 2 inches thick,
and marked by concentrations of opaque minerals in both sand bodies.
Light yellow and yellowish-brown oxidation and clay enrichment ex-
tend to depths of 3 to 6 feet in the upper sand body. Thickness, 18 feet.
(Oaks and Coch, 1973, p. 94—95)

Kempsville Formation1 (Pleistocene) in southeastern Virginia—Beach and

dune sand and gravel in Hickory scarp; minor marsh clay among west
margin of unit. The Kempsville generally consists of fine- to coarse-
grained sand with minor amounts of fine pebble gravel. The sand is
composed chiefly of angular to subrounded quartz with 1 to 3 percent
opaque minerals and with rounded, broken shell fragments. Locally, the
Kempsville sand interfingers westward with silty, very fine sand, clay
and soft peaty clay. The Kempsville Formation is oxidized yellow to
reddish brown to depths of 3 to 4 feet. Clay in the weathered upper part
acts as a binder and makes the sand very hard when dry. (Oaks and
Coch, 1973, p. 81)
Elsewhere in the Coastal Plain of Virginia, the unit is Pleistocene(?)
shoreline sands, as inferred from strong relict beach morphology and
from county soils maps that show soil types with predominant sand in
the C horizon. All sandy environments of the shore and barrier complex
are probably included in the bodies mapped. Thickness, 10 feet. (Force
and Geraci, 1975)

Norfolk Formation1 (upper Pleistocene)—Shelf fine-sand facies. The
fine-sand facies consists of light bluish gray, fine sand to silty, fine— to
very fine grained sand, with 3 to 8 percent mica and 1 to 5 percent dark
minerals. Locally, clayey sand and sandy clay are abundant. The com-
position of this facies is highly variable laterally. Laminae are chiefly
horizontal and outlined by dark minerals but locally dip as much as 10
degrees. (Oaks and Coch, 1973, p. 76)

Norfolk Formation1 (upper Pleistocene)—Fluvial and estuarine clayey
sand facies. Sand, light-gray to tan, medium- to fine-grained, pebbly
locally, 1 to 4 percent heavy minerals, thin- to thick-bedded, with lenses
of clay (up to 8.2 feet) and peat and organic—rich clay (up to 7.6 feet)
locally. (Johnson, 1972, p. 32)

Norfolk Formation1 (upper Pleistocene )—Marine sandy clay. This facies is
the surficial stratigraphic unit over most of the Hampton flat and ranges
in thickness from 10 to 15 feet. A cobble zone with flattened quartzose
boulders up to 12 inches in diameter commonly marks the unconformity
between the sandy clay facies of the Norfolk and the underlying
Yorktown Formation. The sediments in the lower part of the sandy clay
facies are sandy clay and clayey sand with varying proportions of pea
gravel. These sediments are unsorted and in places stained reddish
brown. In the upper part of this facies, the amount of sand diminishes
eastward. Because of the high water table over much of the area,
oxidation has not taken place to depths greater than 3 feet. (Johnson,
1972, p. 36)

Norfolk Formation1 (upper P1eistocene)—Marine silty sand. This facies of
the Norfolk Formation crops out in ridges on Goodwin Islands, Crab
Neck, and Plumtree Island. The ridges are composed of fine- to
medium-grained sand, with less than 10 percent silt and clay. Small
quantities of gravelly sand occur locally on the ridges. On the broader
ridges, the sediments are slightly finer on the west side than on the east
side. The sand on the crest of the ridge lacks distinctive primary
sedimentary structures and shows shallow weathering profiles. This
facies is usually less than 10 feet thick. (Johnson, 1972, p. 37)

1A5 mapped by Oaks and Coch (1963, 1973)

Sediment and radiometric sample locality
Sediment sample locality (Force and Geraci, 1975)
Aeroradiometric anomaly investigated but not sampled

Radiometric sample

 

A/\
O ,,
UNITED STATES DEPARTMENT OF THE INTERIOR LII W \> Q PROFESSIONAL PAPER 1263
Q) h (L X PLATE 2

 

     
  

    

    
    
 

 

 
 
 
 
  

     

 

    

 

 

 
  
 

 

 
 

 

    

 

 

 

 

 
 
   
 

   
  
 
 

 

 

 

 
      

       
        
 

 

 

 

 

 

   

 

   

 

 

 

 
 
  
  

   

 

 

    
  

 

 
 

 

    
 
     

 
     

 
 
 

        
   

 

  
  
   

 

 

  
     

  
 

 

 

 

 

    

 
 
 
  
 

  

 

 

 

     
 

 

  

 

   
   

    

 

 

 

 

 

 

 

 

 

 

 

    

 

 

    

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

   
  
    

   
 
  

 

   
   
    

 

 
 
 
  
   
 
 

   
    
 
   

      
 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

GEOLOGICAL SURVEY 75°00’
77°00' 1 , 1 , 30' 15’ . , 1 I 38°00’
33°00' . 1 . 45 ,, ,, ,, 30 11 ' , .1 I I 1 .. I: ,7 1 .1 1 " x5
1“. Tit?! Islandfl ‘
1‘ CHINCO v21 1 Q 1 r ,1 cf
11 B . , ' r/ASSAIEA’GU ISLAND '
I 7 In he I I " ,N 1‘31, m , , 1\ I
, Itum I I ' ’ / '
‘ y Ts - It '
i T») T; ‘- T _‘
1 1 . NE Um W gs ,12’1'\ 5
. \ '2 I: \1‘. 1‘ I ‘3, ”co (/
If I a _. ‘7 I . C
Y T L,‘ «\I ” _ Shanksistzrd1
\ x40» .1
Ix, , ‘ 7 \.
,1/1 /
)2?)
1 go ‘‘‘‘‘‘ ;~«/»
77777 ,, ‘ 1 . "9 BEASLEY / \I
1/" BAY 15° /
3/ I
w :1 j; I)
:7 11
— 45’
I‘” _ . ‘1 ‘ 1' ‘ I E I7 . _‘ Y1 11: I: 11‘“ II,” 3, 1' 1« / / I Parker: Isiun
FInnGYI Isl 5‘ i ' . I 1 I . 1 1
Storkomugh Hand :is I _ _ 1 j 1 L I / ~1 I I . .V g 1/ , , ,, , . 1/1f, L L » ~ " '
Fl?rjlsland :1111\ \I 8 ;;
\ i 1\r\_\ 1 ‘1/
1 jWindmiil ' I 1 , Y
' P9,)” ,1 111x i I/f / 1_1
”I? ' t I ' EXPLANATION
\ ‘ .3 ’” COUNTS PER SECOND
Light . I . I
O ’ -1 I . I
‘Efi‘stingray 5 i E x I
1 “Wynn Istand 7 If/ I
1 III/'1 1 ,1...., _ 30’
/ ITS
. Emmi Light 1
t Mackinaw)
Inter \
5 Less than 24
I; _. . obit Isiamf
.‘ ' ‘1. ‘ I7 I x I; I 6 Radioactivity contours—Showing net intensity in
, I Q counts per second after removal ol cosmic compo—
x.” f I 1 I nent and adjustment for deviations from surveying
1 8‘ owns Bag; 7 V '1 1‘ I altitude Hachured to indicate areas of lower radio—
? Creek: [I '9 __ ‘ II activity. Contour interval 25 counts per second
n ' , ”E" V l ——77‘ Flight path—Showing location and spacing of data.
1' ~'1 1}( - «1 ; Flight—line spacing one mile. Flown at 500 feet
I , ‘ :I Y I, “ I 1 \1' :1 1 above ground in an east—west direction
a.“ T 39‘ 1‘1 \I 1 I 1
s T
sww/f—f’ E. Wit; sir 1 “I “My ,1 /
1Dcmc2ng i Barreis Pom? r’” . ”111 ..« 1 .- 1 s.‘ ‘1 1 / ,,,,, 1(9"I"T“=.__ _ ‘ \ I’1 {I \ 1‘ , l I)" 1 __ /I/ <1: 1
A ‘, 1. 1 no ‘\ I i ‘ \ a a I , ,, ,, . mm
‘P‘m’ I Jaw Em“ QC \ . . . I 7 \\ 1 " » H . g j __ »- -* “““ , I 111“‘_\__\ w” , 1' '1‘. ... 1 . 7 ' 1‘; ' “ """ 1I I 1 1
”” sURRY 05‘ ,1 1, p} I 1 , . ‘ ' ‘ , , - ' ' ' ' " ' " ' ' -1 _ ' 1 ‘ 1 1 ,
1 7;“ \I I I , I
‘Tv-Wjfim , “ 1
\f/j/ ‘ 1' I
I j; “'30 Km /1/1
E Jamestown "I » . 1’ l II I
\ Hand 9 30I
I 5;? 761 M1 1‘1 rossrri-7O1EEEK) (“mummy 7915’
I 1 ‘ _ If‘j7’il)” __4, “To :0
(1((331‘2‘11211’11‘? 1 - ” - ,\ TI/_
1T1 i‘: f)" ‘2; 1‘
T /N{ , I Lighthouse
“751mm :1-
37 00 _ r ‘1' Id EW’Ebrhfo; :111111“s_,1\ 1 K \1? 7 I j ‘ “I, i, ‘ 370001
1 pmw ROADS B}? vamnmwfib \\\\\ " III!) 1 __\ 1‘ ' ‘1 7
i _ 4 \ I _{II 7 \_ 7 7
§ \
32 ‘w
:5 t
:Alegam‘tjff 1~ 1 t A " . i_hIQA41_—__L_:;::‘__:___A_‘_:_*_:_:€fl*§§fifljfléfifl_hg:;_:__‘i. - ‘ “ ‘ * “ “ ’LE;———I-fiI—¥%§%fiaum
"Tfiérfimmum ———————————
a , ~ » 14) ” ~ "““T‘ 0 V2 I 2 MILES
36 45 \ ’ I 1 ads / IBaIIaLfsl _ 36°45' I I I I l I I I
-~ O» 0 If \-
(1 a I I ya )6 0 ‘5 1 2 3 KILOMETEHS
. 7 I, .
O; [y« / «dflvjx // CONTOUR INTERVAL 5 FEET
fl: 3 l/ é? '5“? 6 NATIONAL GEOOETIC VERTICAL DATUM OF 1929
11:; m‘éiIn-Rose ‘11; I.
53$”- 3‘7 I125’; . y 831.511
1, .K/
_ I
E X P L A N A T I O N
111111 , . 1 . ‘ 1 1 ' ‘ 11
_ . I, 1' '1 I ' I I ,. 1 , 4———‘— 250 counts per second contour—Hachures point tolower count rate *The International Atomic Energy Agency (1976) has replliced all
It , .1 1 8 _ _ _ I * previous measures of crustal radiation intensity wth the
I . Ground data locality Values m “I “Radioelement Unit" (pr), defined asthe radiation measund from
SURVEY 1’ AG78N © Sediment sample (see Table 1 for analyses) and radiometric sample a source containing 1 ppm uranium in equilibrium, and eq‘iivalent
1 ’ (see Table 3) to 046 [JR/hr.
' ’ \
§OUNDARYII (17/) Both 17px values are manmade anomalies such as rock road sur-
f ; T faces, concrete ponds, and stacked culvert (concrete) pipes
'
INSET A—Ma illustratin failure of round data to locate anomalous] radioactive sediment in an area used for
1 P Q 9 y
I agricultural purposes. In the area west of Lake Drummond Church,farmlands (cleared areas on map) are
‘2 fertilized by 600 pounds Of 5% nitrate, 30% potash, and 30% phosphate per acre several times a year
“A (depending on crop type) particularly in late fall and early spring
/ . I , _.

771000, 45, 30) 15/ 76°00' INTERIOR—GEOOGICAL SURVEY, HESTON, VA.—1983~682621
Base from US. Geological Survey Modified from data flown and
Eastville, 1959; NorfOIkI 1959; and SCALE 1:250 000 comolled by LKB Resources, Inc.
Richmond, 1973 5 o 5 IO 15 20 25 MILES Huntmqm Valle‘I. PA» “375—76

l—l I—I l—l I—————I I—-I I
5 O 5 10 15 20 25 3O 35 KILOMETERS
I—l l—I l—l I—I l——I l———————l I

 

 

CONTOUR INTERVAL 30 FEET
NATIONAL GEODEIIC VERTICAL DATUM OF 1929

TOTAL-COUNT CONTOURED GAMMA-RAY AERORADIOMETRIC MAP OF THE OUTER COASTAL PLAIN OF VIRGINIA

 

    
 
      

        

      
     
  
   
    
  
  
  
     

 
  

 
  

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

  
   
    

.«K I
V\ )1}?
UNITED STATES DEPARTMENT OF THE INTERIOR W \ PROFESSI‘C)’ ‘ L PAPER 1263
(A! ‘53” a . .~ 9 I “Q,
GEOLOGICAL SURVEY \E’éfs giizm‘gfiz PLATE 3
0 755°” 20' ““M 78°15’
37 00 \ / 37°UD’ EXPLANATION
\/ 1.00 Radioactivity contours—Showing net intensity in
counts per second after removal of cosmic compo—
O nent and adjustment for deviations from surveying
altitude. Hachured to indicate areas of lower radioac-
tivity. Contour interval 25 counts per second
/ +
I. — ¥ ~~ Flight path—Showing location and spacing of data.
Flight—line spacing 1 mile. Flown at500 feet above
HAMPTON ROADS ground in an east-west direction
BRIDGE-TUNNEL
Aeroradiometric high >300 counts per second
1 1 .
\ + _
Aeroradiometric low <100 counts per second
0
.9
162 — ‘
D n?
J A M E S R I V E R
0:
DJ
>
N
1r
55’

 

 
 
  

I 275
55'
Q
«I?

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\

 

 

  

 

 

 

 

SU‘ 50’

 

 

L L

36°45’ o , 35°45,
76 30 INTERIORiGEOLOGICAL SURVEY, RESTON. VA—1983—682621 76°15:
Modified from data flown and compiled
by Geodata international, Inc. 1976—77

| l l
l | | l | l | l

l .5 O 1 2 3 4 5 KlLOMETERS

1‘ 112 0 1 2 3 MILES
l

TOTAL-COUNT CONTOURED GAMMA-RAY AERORADIOMETIC MAP OF THE PORTSMOUTH AREA, VIRGINIA

  

    
      
   

 

UNITED STATES DEPARTMENT OF THE INTERIOR

GEOLOGICAL SURVEY
77°UU/

PROFESSIONAL PAPER 1263
PLATE 4

     

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‘ I 03“; ’
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61 / My. «‘7 “I, nr/v‘ o r
,2 {“I u A, 76 00 15' 75000’O 0,
:9 IN N J \m IxSmIth island JWES WW, , " x 76 I“ 3’} ‘Fd._'» (I; ,/ I / , é " , , 38 D
\ \ I .1 ' iI/ , , , 3mm) rsiand «I 'EIMHIII‘MR EL. - I , . f. .0 “' , “II 4” /, 'I I7
. a 7 a a f 922} W g: II . 5., CHINCC)TEA($U 2‘ , , “s‘ / _ * ;
N w n ’5‘ llama {Siam 8:41 Y / ASSA’TEAGIUEI ISLAND [37 * _ ,
A , Q /’ I \ » \ v 51' 1
rim “‘ .. ,5 “firm ["0 , ,
to , I X;- , ,r , ,I" ,' ~ . ,..
’ E Q? / c’ I l/INATIONAL SEASHORE , I‘ A D
- a, V > i at or [sin . / , " /
Shanks Mark/VG S U , I“ [ I I, I I ,/',
R772"? ,Igj L . I a- / I I
~ z *1. L. L Mama: ;
, , a I 1.1; $0. ,
. m 5:]: 4,15 é‘ttump IsIand at“: \
A Gr-mror ikad €53 w, I :I 4 I , , I,
. m <- 4» »- tr: - /' ,
r . 4.... . “I ‘ - o , ‘ ~ " ‘I ,
\\ ;/ fife/$5M?» 1 , Guam IsIonflkfi‘t , I 5} :n‘ne Pox Island; $3 54 “/2 $ , I in‘Tiftj’L11\I7a\‘y yr
4 it ,1? , , ’ Inlet "x, \_ / , )‘I / I , ~-
AI A“ I I 9 BE‘IIS‘I‘EY {It NACA 'WALLst lESLANiTII/I ’ I I "
Q , I HA 3; , . s . v I . I ) fl/ , /
'5" HI .5!
42:5 1
8/ . \
Simian irfondI I I;
, _, » M _ _ 45,
g; r' L (mans-«f . ‘
/ , 3 ,I , , .x / ~
Farrier: Triage _ I/
‘ . _ :11 r' , I imam Iriondvflfi
‘ ' ‘ iiwrtwrough trier-d fix
, z I
33;. n E 15' 75°00’
yum:
new: , z I
9*, I, W a2. EXPLANATION
\ A' A; I. fleets Isitmd I ‘ 13‘: w
V III" II], D , % Anomalies investigated
_ warrant: Wind”; ’1, y " I ‘ ,rI
Sfijiflm Paint I II) I I 'I .er
”1’,er n, M “is w “W“ n \ “I” Q , , ,, I
£31 I, (I N ’ TN"\ 34.; I 'I r’II ’I , III I C)
” ‘ x , I ' 1 URBAN OR BUILT-UP LAND
L73 I 11 Residential
, a - , , v, 12 Commercial and Services
I _ ,. j -, I I'/ 13 Industrial
>I":’I ’_ \ , / (II , . III I 14 Transportation, Communications,
I I ' j I II and Utilities
.’ “II I 15 Industrial and Commercial
\ , Complexes
I 2’ Q) QI’IIIIII IIIGIIII ,I I, 3 16 Mixed Urban or Built-up Land
_ In: 4 4‘ 9‘ Mm I I I ’ 17 Other Urban or Built-up Land
- - J 1’: ; a}, w a 7 {I I W _30I
- t}: I
I»..//., 5i ’I I
rfiflzwis If a)» '
1,. In ’ b” “ ’ ; I I ' I 2 AGRICULTURAL LAND
, 1 r, ‘ ,I . I
M: I 21 Cropland and Pasture
22 Orchards, Groves, Vineyards,
,I'I “ Nurseries, and Ornamental
Horticultural Areas
I 23 Confined Feeding Operations
IIIII 24 Other Agricultural Land
$3.453 Qfafigk‘fi II , ' 71%,;(Iy'mrz! Mzr‘ckr‘mmgo & 3 RANGELAND
\ mg. , ,‘ 1’18
M'Mwmmdfi Cu : (we; rim; I I
W!“ {.3 K , , . . . _ \ . Marl , / 31 Herbaceous Rangeland
. : , W. @x . , , I. 1 I‘ . ’ % 32 Shrub and Brush Rangeland
51/ 43} ,r/I It“ “EI 33 Mixed Rangeland
. "7 /’/ . ; ’ l
< / { Cleri'tar 570/, ‘ C I
I ”I ’1'" T :I I: ~ , I ‘ II
-315: efmm‘ I V . n. -1
”/w/M II J; 2 V ,mz‘I N A I In
a \ I , I; 5W ' - “ ' 2 4 FOREST LAND
rm; w I -. a __
, v , , , ' - 41 Deciduous Forest Land
:2? Cams (Issue , ‘I ' 42 Evergreen Forest Land
J tighten» 43 Mixed Forest Land
\rar mmg Ir gum: rim; , ‘- _ ‘ ‘ I — 15'
\\“\WI}QI”‘I AJAMEFE: C33: , _ :5 fr In . , :i _ I ‘ __ ' I, I I ,' _ , I
“WW M ”EILEEN? M "(KNEW ; "I f" . “ _ . /’ I I ‘ tight 'I , i
«w: a . . . , 2 -. a s x. 2 ‘I I w . . r 7, m. 5 WATER
Wéwrzms Pei, I ”:77fitstrm ” . '7’" 3:73:71?” II I ._ ‘ I I
21 “‘4 II II XIII“ IYL“'“"2~29, W [K I ,, E2 - 1 I: might _ ,I _ 1 _ I 51 streams and Canals
62 j’I egg u I 52 Lakes
I579“ 1/ ‘ “ I I '\ \ Mymmmr , _ , I , : , 53 Reservoirs
<9“ 1/ Z” Iv M _ I ' , I I I 54 Bays and Estuaries
m.“ r, " ”,1 York pom? \‘ 45' 30/ I I \
woman ' c F .. . .
rm} 3.» U M?” at ’i‘iiiili was)“: _
,3. \f, Eu ' “1.. NW 2»:
, i": I 5 4 6 WETLAND
‘gx~ ‘? INDEX MAP OF COUNTES
/ ‘ 77° 76° 61 Forested Wetland
M " 38° _ 62 Nonforested Wetland
>72» \\\ \‘~ P, 9
\t \\ Heathsville \\ “fly/7
a I 2.; ~r~ as
I r,» 7 BARREN LAND
71 Dry Salt Flats
72 Beaches
73 Sandy Areas Other than Beaches
74 Bare Exposed Rock
I, , I .1 75 Strip Mines Quarries, and
F” if “QI . g: s, III” mar; Gravel Pits
,2; -; writ MUNRUR; WI? 76 Transitional Areas
’6, «,1 // my??? / ”I V w ‘ 77 Mixed Barren Land
I) ,y- _ .4" I «I I
WfiMFFQN .2 37°00’
\g 5 I, ‘ \_
81 TUNDRA
81 Shrub .and Brush Tundra
82 Herbaceous Tundra
83 Bare Ground Tundra
E HfNB‘f‘ 84 Wet Tundra
" " grow .
4H ‘ I'F‘tifi‘f‘ WESLEY/$3.7 ION 85 Mixed Tundra
9 PERENNIAL SNOW OR ICE
— 91 Perennial Snowfields
92 Glaciers
mm 10 20 MILES For definitions of Level I and Level [1 categories see, US.
I Vil”Qinia Beach Geological Survey Professional Paper 964, A Land Use and
IINIDIIIIIJILINN; 214.7113? Princess Anne 0 10 20 30 KILOMETERS Land Cover Classification System for Use With Remote
1 0 Sensor Data, 1976, by Anderson, J. R., E. E. Hardy, J. T.
10 Roach, and R. E. Witmer. Minimum mapping units are: 4
hectares (10 acres) for Level II categories 11-17, 23-24, 51-54,
L _ am, 75, and urban occurrences of 76; and 16 hectares (40 acres)
.. TRAINING swim for all other Level II categories.
36045/ II ‘1' 4 . imam _____—__—__
NORTH CAROLINA
— 36°45’
NUMBER COUNTY/CITY GROSS TONS TONS P205 TONS KCl
, 1 ACCOMACK 29,400 2056 2740
> gwés /,2/ $2 I. 2 NORTHAMPTON 22,812 1459 1900
\JLI j A?” 3 NORTHUMBERLAND 10,950 1119 1859
\ 42' Xx. I; *6 4 LANCASTER 6,736 518 764
MINI IgI 1 “I21 5 MIDDLESEX 5,759 510 790
I 6 GLOUCESTER 6.256 379 708
7 MATHEWS 3.709 183 327
8 YORK NO RECORD NO RECORD NO RECORD
9 NORFOLK 7,228 670 468
10 VIRGINIA BEACH 24,159 1411 2281
11 CHESAPEAKE 16,532 955 1345
12 SUFFOLK 17,029 864 1457
II I Gross fertilizer consumption by counties in the Coastal Plain of Virginia, July 1976 to June 1977.
' Source: Virginia Polytechnic Institute, Agricultural Experiment Station, oral communication
I July 20, 1978 '
NGUEW Lej’i; g.
. II‘I‘IT 15' INTERIOR—GEOLOGICAL SURVEY, HESTON, VA —19837882621
Base from US. Geologlcal Survey - -
Eastville, 1969; Norfolk, 1969; and SCALE 1:250 000 Data from Land Use Series, Open-File reports:
Richmond 1973 77—0624, Eastvrlle quadrangle, VA., N. C., MD., 1972—73
' 5 0 10 15 20 25 30 MILES 77—085—1, Richmond quadrangle, VA., MD., 1973
. 1 . I ._._I r . . - I 77—789—1, Norfolk quadrangle, VA., N. (3., 1972

 

5 0 5 10 15 20 25 30 35 KILOIVIETEFIS
I—J l—l l—l . . g - - - 1

NATIONAL GEDDETIC VERTICAL DATUM OF 1929

LAND-USE AND LAND-COVER CLASSIFICATION OF AERORADIOMETRIC ANOMALIES IN THE OUTER COASTAL PLAIN OF VIRGINIA

 

PROFESSIONAL PAPER 1263

UNITED STATES DEPARTMENT OF THE INTERIOR
PLATE 5

GEOLOGICAL SURVEY

T

SAND BRIDGE FORMATIONI, BARRIER-
SAND-RIDGE AND MUD—FLAT FACIES
/_—/%\

SAND BRIDGE FORMATIONI, MARSH AND UNIT A, BARRIER ISLAND HOLOCENE SAND BRIDGE FORMATIONl, FLUVIAL AND HOLOCENE SAND BRIDGE FORMATIONl, ESTUARINE UNDIVIDED SAND AND GRAVEL

TIDAL—FLAT SILTY-CLAY FACIES OR SPIT COMPLEX FLUVIAL SAND LAGOON SILTY —SAND FACIES SHORELINE SANDS AND TIDAL-CHANNEL FACIES (OUTER COASTAL PLAIN)
A

 

 

Sample number AG 19N AG 20E2 AG 23N AG 53N AG 80aN AG 80bN AG 54N AG 91R

Size distribution of
economic heavy
minerals (sp. gr.
>285) in the total
sample

Distribution of
heavy—mineral
species in the
CFS fraction

garnet: P zircon = P

Distribution of
heavy—mineral
species in the
VFSS fraction

garnet=P garnet=P garnet=P garnet=P
monazite : P

EXPLANATION

A Percentage of total heavy minerals in the very ilmenite

fine grained sand to coarse silt-sized fraction
garnet

B Percentage of total heavy minerals in the coarse—
to fine~grained sand fraction

\ . .

. . . . . . i . zit n Zircon
Distribution of economic heavy minerals in the \_ mona e a d

very fine grained sand to coarse silt—sized \‘ \ fume and leucoxene

fraction

sillimanite and kyanite

epidote, amphibole, staurolite, magnetite, mus-
covite, :apatite, :limonite, :sphene, :go-
ethite

Distribution of economic heavy minerals in the
coarse- to fine—grained sand‘sized traction

1As mapped by Oaks and Coch (1963, 1973)
2Sample AG 20E was collected from 15 to 20
feet (46 to 61 meters) below the surface

P: <0.5 percent

 

 

 

INTERIOR—PGEOLOGICAL SURVEY, RESION, VA 719837682621

SIZE DISTRIBUTION OF HEAVY-MINERAL SPECIES IN 15 SAMPLES THAT CONTAIN MORE THAN 1 PERCENT TOTAL HEAVY MINERALS COLLECTED
FROM AERORADIOMETRICALLY ANOMALOUS LOCALITIES IN THE COASTAL PLAIN OF VIRGINIA

 

 

‘

Earth Flows:
. Morphology, Mobilization,
and Movement

* > By DAVID K. KEEFER and ARVID M. JOHNSON

 

kg GEOLOGICAL SURVEY PROFESSIONAL PAPER 1264

k A f;
' ”a A study of the factors controlling earth-flow occurrence
v fix: and the velocity of an active earth flow
4

 

 

 

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON : 1983

UNITED STATES DEPARTMENT OF THE INTERIOR
JAMES G. WATT, Secretary

GEOLOGICAL SURVEY
Dallas L. Peck, Director

Library of Congress Cataloging in Publication Data

Keefer, David Knight
Earth flows: Morphology, mobilization, and movement.

(Geological Survey professional paper 1264)

Bibliography: p. 53—56

Supt. of Docs. no.: I 19.16:1264

1. Eanhflows. 2. Earthflows—California. I. Johnson, Arvid M. II. Title. III. Series: United States. Geological Survey. Pro-
fessional Paper 1264.

QE599.A2K43 1983 551.3 W499

 

For sale by the Superintendent of Documents, US. Government Printing Office
Washington, DC. 20402

L)

CONTENTS

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

Page
Abstract 1 The Davilla Hill earth-flow complex—Continued
Introduction 1 Studies of active earth flows and ground-water conditions,
Previous studies 2 1974-78— Continued
Classification and terminology -------------------------------------- 4 Subsurface displacements ---------------------------------------
Acknowledgments 4 Water levels
Earth-flow morphology, materials, and settings -------------------- 5 Water contents
Topography, climate, and geology in areas of earth-flow oc- Unit weights, void ratios, and saturation -----------------
currence 5 Laboratory swell tests -------------------------------------------
San Francisco Bay region -------------------------------------- 5 Piezometric and tensiometric measurements, August
Other regions 8 1975—February 1978 -------------------------------------------
Morphology 8 Analysis of mobilization
Idealized earth-flow morphology ----------------------------- 10 Measurement of shear-strength parameters, using di-
Subsurface features: Basal shear surfaces, lateral shear rect—shear tests
surfaces, and associated zones of deformation ------- 13 Slope-stability analysis of earth flows ----------------------
Earth-flow complexes 16 Discussion
Materials 17 A model for earth flow
The Davilla Hill earth-flow complex ------------------------------------ 22 Earth-flow mobilization
Description 22 Model of a moving earth flow ---
Location and setting --------------------------------------------- 22 Velocity measurements ------------------------------------------
Morphology and history of movement ---------------------- 24 Shear surfaces and adjacent zones of disturbance ------
Material 27 Distribution of displacements ---------------------------------
Subsurface geometry — ------------------------------------------ 27 Surges
Studies of active earth flows and ground-water conditions, Velocity, boundary slip, and boundary roughness ------
1974-78 28 Internal deformation
Surface displacements ------------------------------------------- 28 Composite model
Correlations of precipitation with mobilization and rate Conclusions
of movement 29 References cited
ILLUSTRATIONS

[Plates are in pocket]

PLATE 1.

N

Map showing earth-flow complexes in Eden Canyon and vicinity, Alameda County, California
. Map and cross section showing the Hidden Valley earth-flow complex, Alameda County, California

3. Maps and cross section showing the Davilla Hill earth-flow complex, Alameda County, California

FIGURE 1. Location map of part of San Francisco Bay region

 

2. Photographs of typical earth—flow deposits

 

3. Bar graphs showing inclinations of slopes on which earth flows occur

4. Block diagrams showing idealized earth flow

 

 

5—7. Photographs showing:

5. Main scarps and crown cracks of earth-flow complexes

6. Deposits in zones of depletion of earth flows

 

 

7. Toe of earth-flow deposit in Bear Valley area

 

8. Diagram showing cross-section through pressure ridge flanking earth flow at Kirkwood Creek, Mont. --------------------------------

9. Photograph showing lateral ridge in earth flow at Melendy Ranch complex

10. Block diagrams showing formation of lateral ridges

11. Photograph of lateral ridge in earth—flow deposit in Cordelia, Calif, area

 

 

 

12. Photograph of slickensided lateral shear surface on inner flank of lateral ridge at Melendy Ranch complex --------------------------

[II

Page

29
32
32
33
33

36
38

39
40
41
41
43
43
43
46
47
47
48
51
51
52
53

IV

FIGURE 13.
14—16.

TABLE

17.
18.
19.
20.
21.

23.

25.
26.
27—28.

29.
30—33.

35.
36.
37.

39.
40.

gmmpwmw

CONTENTS

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

Page
Diagram showing multiple basal shear surfaces underlying earth-flow toe at Beltinge, England ----------------------------------------- 13
Photographs showing:

14. Types of cracks disrupting earth flows 14

15. Slickensided shear surfaces 14

16. Lateral shear surfaces in zones of depletion of earth flows at Davilla Hill complex 15

Block diagram and map of trench cut through earth-flow deposit in Cordelia, Calif., area 15

Photograph of convoluted contact between materials in trench cut through earth-flow deposit in Cordelia, Calif., area ---------- 15

Diagram showing earth-flow deposits exposed in wall of gully near Davilla Hill complex 16

Photograph of earth-flow complexes occupying long narrow channels 16

Aerial view of two earth-flow complexes in Cordelia, Calif., area 16

. Photograph of hummocky earth-flow complex near Pleasanton, Calif. 17

Chart showing plasticity indexes and liquid limits of earth-flow materials 22

. Geologic map of Eden Canyon area, showing location of Davilla Hill earth-flow complex 23

Chart showing daily precipitation in Eden Canyon, July 1974—February 1978 25

Photographs showing Davilla Hill earth-flow complex 26
Graphs showing:

27. Cumulative displacements of survey stakes on earth flows 30

28. Displacement of survey stakes on earth flows of Davilla Hill complex 31

Photograph and graph showing measurement of subsurface displacements on earth flow 1 of Davilla Hill complex ---------------- 32
Graphs showing:

30. Water levels in boreholes 33

31. Natural water contents in boreholes 34

32. Natural water contents in pit 1 35

33. Piezometric measurements 36

. Sketch illustrating forces acting on block on inclined plane 39

Graphs showing MohnCoulomb failure envelopes for direct-shear tests on samples from basal shear surface ------------------------ 40

Graphs showing factors of safety as functions of water-table depths 42

Graphs showing movement patterns of earth flows monitored with continuous-recording devices --------------------------------------- 47

. Diagram showing subsurface displacement profile of earth flow at Beltinge, England 48

Graphs and diagrams showing surface-displacement profiles of earth flows 50

Graph showing shear resistance as a function of velocity in direct-shear tests on basal-shear-surface material ---------------------- 52

TABLES

Page

. Bedrock in areas containing abundant earth-flow deposits in the San Francisco Bay region 7

. Bedrock in areas containing abundant earth-flow deposits outside the San Francisco Bay region ---------------------------------------- 9

. Composition of earth-flow materials 18

Clay-mineral content of materials from the Davilla Hill earth-flow complex 28

. Unit weights, void ratios, and saturation of Davilla Hill earth-flow material 35

. Laboratory swell tests on Davilla Hill earth-flow material 35

Earth-flow velocities 44

 

Any use of trade names and trademarks in this publication is for descriptive
purposes only and does not constitute endorsement by the US. Geological Sur-
vey.

 

 

EARTH FLOWS:
MORPHOLOGY, MOBILIZATION, AND MOVEMENT

By DAVID K. KEEFER and ARVID M. JOHNSON

ABSTRACT

The terms “earthflow” and “earth flow” (after Varnes, 1978) have
been applied by previous investigators to several mass—movement
phenomena involving fine soils. In this report we describe studies on
one type of earth flow—a type characterized by movement at veloci-
ties of a few meters or less per day that persists for several days,
months, or years. These studies were conducted to determine what
morphologic and material properties characterize earth flows of this
type, to analyze how such earth flows are mobilized, and to determine
what mechanisms control their movement.

During our studies, we examined numerous earth flows and
earth-flow deposits in the Coast Ranges of central California; moni-
tored earth-flow movements at three sites; and conducted detailed
field, laboratory, and analytical studies at one site—Davilla Hill, near
Hayward, Calif. Using information from these studies and from pub-
lished descriptions of other earth flows, we analyze earth-flow mobili-
zation and propose a model to explain what controls the velocity of an
active earth flow.

Our studies show that earth flows are commonly tongue or tear-
drop shaped. They have rounded, bulging toes and sinusoidal longitu-
dinal profiles, concave upward near the head of the earth flow and con-
vex upward near the toe. Many earth flows are flanked by lateral
ridges composed of immobile earth—flow material. Earth flows are
bounded by discrete slickensided shear surfaces. Materials in earth
flows consist primarily of silt and (or) clay and smaller amounts of
sand—size and larger particles; these materials are pervasively fissured
and contain a significant amount of entrained water.

At some sites, earth-flow deposits blanket areas of several square
kilometers. These large masses of earth-flow material are rarely de-
posited by a single episode of movement but, rather, are complexes
built up over several years by many individual earth flows. Earth
flows making up these complexes are mobilized out of older earth-flow
material, out of material previously deposited by other types of mass
movement, or out of material not previously disturbed by mass move-
ments.

We studied the earth-flow process over a 4-year period at Davilla
Hill. Mapping of earth—flow features and monitoring of active earth
flows showed that the Davilla Hill earth-flow complex is made up of at
least 34 individual earth-flow deposits. The earth flows active during
the study period were formed out of older earth-flow material. When
active, these earth flows moved several centimeters per day, and most
movement took place by shearing along boundary shear surfaces. The
earth flows were mobilized by increases in pore-water pressure caused
by infiltration of water into the soil during and after rainstorms.
Mobilization was accompanied by softening of the material and by an
increase in water content due to saturation; no significant volume
change or remolding of the material occurred.

Analyses of earth flows elsewhere show that they also are gener-
ally mobilized by increases in pore-water pressure. Although short

 

high-velocity surges occur on some earth flows, most are characterized
by persistent movement at velocities ranging from less than a millime-
ter to several meters per day. Most movement takes place by displace-
ment on or immediately adjacent to the boundary shear surfaces. We
conclude that the general absence of sustained acceleration in earth
flows is due to boundary roughness. Factors controlling earth-flow ve-
locity are: pore-water pressure, the characteristics of asperities on the
shear surfaces, and the properties of material on and adjacent to the
shear surfaces.

INTRODUCTION

Earth flows are among the most common mass-
movement phenomena in nature; they occur in many of
the world’s hilly and mountainous areas. In parts of the
California Coast Ranges near San Francisco Bay, where
our field studies were carried out, thousands of earth
flows and earth-flow deposits1 are present: they are
identified by dish-shaped scars, by bulging toes, and by
long narrow tongue- or teardrop-shaped forms. Some
features, such as bulging toes and smooth surface pro-
files, suggest that earth-flow movement involves a com-
ponent of fluidlike flow; other features, however, such
as slickensided shear surfaces, suggest that earth flows
behave more like rigid bodies that move by boundary
shear.

Earth-flow activity has been recognized as a dis-
tinctive type of mass movement since at least the early
20th century (Howe, 1909; Blackwelder, 1912). Since
then, numerous individual earth flows have been de-
scribed, and detailed investigations have been carried
out on a few. Using data from these descriptions and in-
vestigations and from our own field and laboratory
studies, we have analyzed several aspects of the earth-
flow process. In this report we discuss our studies, un-
dertaken to answer the following questions: What mor-
phologic features and material properties characterize
earth flows? How are earth flows mobilized? What
causes them to stop moving? At what velocities and how

1In this report, the term “earth flow" refers to a moving body of earth and water; once
movement has stopped, the material displaced by an earth flow is called an “earth~flow de-

posit."
1

2 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

long do earth flows move? What are the patterns of in-
ternal deformation within moving earth flows? What
factors control their velocities?

Our investigation was conducted in three stages. At
the outset, we identified areas in the San Francisco Bay
region of California that contain abundant earth-flow de—
posits, from reconnaissance in a small airplane (fig. 1).
We examined several tens of earth-flow deposits on the
ground in those areas to identify the morphologic fea-
tures, material properties, and topographic and geologic
settings characteristic of these deposits. We then com-
bined information from these field studies with informa-
tion from published studies to formulate a general de—
scription of earth-flow morphology, materials, and set—
tings.

During the second stage of our investigation, we
monitored movement of earth flows at three sites—
Melendy Ranch, Cycle Park, and Davilla Hill (fig. 1)—
and conducted more detailed field, laboratory, and
analytical studies of the earth flows at Davilla Hill.
These more detailed studies included: mapping of mor-
phologic features; augering and subsurface sampling;
measurements of ground-water level, pore—water pres-
sure, shear strength and other material properties, and
rates and distribution of movement; and slope-stability
analyses of mobilization.

In the third stage of the investigation, we formu-
lated general models to explain earth-flow mobilization
and the observed velocities and displacement distribu-
tions in earth flows.

PREVIOUS STUDIES

Previous investigations include descriptions of
numerous earth flows and earth-flow deposits, measure-
ments of displacement on several earth flows, and
analyses of mobilization of a few earth flows.

In the San Francisco Bay region, Krauskopf and
others (1939) described the geologic setting, movement,
and morphology of an earth flow in the Lomerias Muer-
tas area (fig. 1). Other earth flows and earth-flow de-
posits in part of the Lomerias Muertas area were map-
ped and described by Oberste-Lehn (1976), who also
measured displacement of one earth flow. Radbruch and
Weiler (1963) mapped and described earth-flow deposits
and tabulated the angles of slopes on which earth flows
occurred in several square kilometers of the central
Orinda, Ca1if., area (fig. 1); a similar study was con-
ducted by Turnbull (1976) in an area of about 6 km?‘
around Davilla Hill (pl. 1; fig. 1). Nilsen and Turner
(1975) studied the relations between landslide occur-
rence, amount and temporal distribution of precipita-
tion, and position of old landslide deposits in a region

 

that included part of the Orinda area. Earth-flow com-
plexes near Davilla Hill were mapped by Aboim-Costa
and Stein (1976) (pl. 2) and Spain and Upp (1976). Ellen
and others (1979) and Peterson (1979) discussed the re-
lation between soil and bedrock type and earth-flow oc-
currence in part of the Hicks Valley area (fig. 1). Our in-
vestigation was previously discussed in the reports by
Keefer (1976, 1977a, b, c) and Keefer and Johnson
(1978).

Earth flows are also abundant in parts of the Coast
Ranges north of the San Francisco Bay region. Cooksley
(1964) described the morphology and geologic setting of
one earth-flow complex2 near the Eel River. Kelsey
(1977, 1978) described earth flows and monitored move-
ment on one earth-flow complex along the Van Duzen
River. Descriptions of and movement data on several
earth flows in the Redwood Creek basin are included in
the reports by Harden and others (1978), Nolan and
others (1979), and Janda and others (1980).

Swanson and James (1975), Swanson and Swanston
(1977), and Swanson and others (1980) described and
measured movement and ground-water level on an
earth-flow complex in the western Cascade Range of
Oregon. Earth-flow movements in the Oregon Coast
Ranges and the Cascade Range were also discussed by
Swanston (1980).

A regional study of earth-flow occurrence in several
thousand square kilometers of northwestern Wyoming
and adjacent areas of Idaho and Montana was presented
by Bailey (1972). Individual earth-flow complexes in this
region were described by Blackwelder (1912), Waldrop
and Hyden (1962), Hadley (1964), Fraser and others
(1969), and Bailey (1972). Blackwelder’s description of
the Gros Ventre earth-flow complex in this region is one
of the earliest and clearest in the literature; this earth-
flow complex was used by Sharpe (1938) as a type exam-
ple of the earth-flow process.

Elsewhere in the United States, earth flows in
western Pennsylvania, West Virginia, eastern Ken-
tucky, and eastern Ohio were described by Sharpe and
Dosch (1942). The Slumgullion earth-flow complex in
Colorado was described by Howe (1909) and Crandell
and Varnes (1961); Crandell and Varnes (1961) also mon-
itored movement there. Two other earth-flow com-
plexes in Colorado were described by Varnes (1949).

Outside the United States, earth-flow complexes
have been described at sites in New Zealand (Benson,
1940), England (Ward, 1948; Hutchinson and Bhandari,
1971; Barton, 1973; Hawkins, 1973; Brunsden and Jones,
1976; Bromhead, 1978), the islands of Barbados (Prior
and Ho, 1972) and Vestspitzbergen (Chandler, 1972),

2An “earth-flow complex" contains several earth-flow deposits and associated features; it
may also contain one or more active earth flows.

L

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V V

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Y

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INTRODUCTION

122°

121°

 

0 10 20 30 40 KILOMETERS

{ figure 1

 

 

 

 

 

     
    

Area of

INDEX MAP

  

 

   

 

 

 

FIGURE 1.—San Francisco Bay region, showing limit of aerial reconnaissance, areas where earth-flow deposits are abundant, and sites
discussed in text. Base from U.S. Geological Survey, 1971, California Shaded Relief Map, scale 1:1,000,000.

4 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

Australia (Pilgrim and Conacher, 1974; Dunkerley,
1976), Czechoslovakia (Fussganger and Jadron, 1977),
and Scandinavia (Rapp, 1960; Prior, 1977). Descriptions
and movement data have been reported for earth flows
in the Panama Canal Zone (Cross, 1924), New Zealand
(Campbell, 1966), the USSR. (Ter-Stepanian, 1967;
Ter—Stepanian and Ter-Stepanian, 1971), England,
Japan, and Switzerland (Skempton and Hutchinson,
1969), and Czechoslovakia (Zéruba and Mencl, 1969).
Hutchinson and Bhandari (1971), Chandler (1972), Dun-
kerley (1976), Fussganger and J adron (1977), and Prior
(1977) performed slope-stability analyses of earth-flow
mobilization.

In addition to these studies, comprehensive site
studies of earth flows have been carried out at Beltinge,
England (Hutchinson, 1970), and Minnis North, north-
ern Ireland (Prior and others, 1968, 1971; Prior and
Stephens, 1971, 1972; Hutchinson and others, 1974).
These studies included descriptions of morphology,
monitoring of movement, measurement of several mate-
rial properties and of ground-water level or pore-water
pressure, and analyses of mobilization.

CLASSIFICATION AND TERMINOLOGY

The term “earth flow,” as used in this report, refers
to the “drier and slower” type of earth flow described in
the mass-movement classification of Varnes (1978, p. 19-
20, fig. 2.1r3). This type of mass movement is called
“earthflow” in the classification of Sharpe (1938) and
“slow earthflow” in the classification of Varnes (1958).
Most other previous investigators in North America
have applied the term “earthflow” to this type of mass
movement; the term “earthflow” was used in our previ-
ous reports (Keefer, 1976, 1977a, b, c; Keefer and
Johnson, 1978).

Our definition of the term “earth flow” encompasses
both “earthflows” and “mudflows” as defined in the
mass-movement classification of Skempton and Hutchin-
son (1969), which has been widely used in the United
Kingdom.3 In more recent publications in the United
Kingdom (Chandler, 1972; Hutchinson and others, 1974;
Bromhead, 1978), the term “mudslide” has been used for
the type of mass movement discussed in our report.

Some semantic confusion in the use of the unmod-
ified term “earthflow” stems from Sharpe’s (1938) use of
the term to refer to two different types of mass move-
ment. Our use of “earth flow” refers to the slower mov-
ing type described by Sharpe (1938, p. 53-55) and illus—
trated by the Gros Ventre earth-flow complex of

“The distinction between “earthflow” and “mudflow,” as defined by Skempton and Hutch-
inson (1969), is based on the amount of internal disruption: mudflows are more disrupted than
emhflows.

 

Blackwelder (1912). The other type of mass movement
(Sharpe, 1938, p. 50-53) commonly takes place in glacial
sediment and involves very rapidly moving bodies of
liquefied material. In more recent classifications, this
type of mass movement has been called “rapid earth-
flow” (Varnes, 1958), “rapid earth flow” or “quick clay
flow” (Varnes, 1978, p. 19, fig. 2.1r2), or “retrogressive
quick clay sliding” (Skempton and Hutchinson, 1969).

“Debris flow” is another kind of mass movement
commonly confused with earth flow. A debris flow, how-
ever, differs from an earth flow in several respects. In
a debris flow, granular soils that generally are admixed
with only small amounts of clay, entrained water, and
air are mobilized into materials that move very rapidly
on gentle slopes. Debris flows are not bounded by dis-
crete shear surfaces, and most movement takes place by
distributed internal shear. Debris-flow material is
coarser grained than earth-flow material. In addition,
debris-flow deposits generally form during a single
episode of activity, so that, within a few minutes or
hours, a typical debris flow mobilizes, flows through a
channel to the surface of an alluvial fan, forms a deposit,
and dries. Such deposits are rarely remobilized. More
detailed descriptions of the debris-flow process were
given by Johnson (1965, 1970).

In our report, the term “soil” refers to any un-
cemented or poorly cemented aggregate of mineral
grains, with or without organic constituents. This usage
conforms to the definitions of Peck and others (1974) and
Varnes (1978). Following the definitions of Varnes
(1978), “earth” is soil in which 80 percent or more of the
grains are smaller than 2 mm in diameter, and “debris”
is soil in which 20 to 80 percent of the grains are larger
than 2 mm in diameter.

ACKNOWLEDGMENTS

This work was funded by the US. Geological Sur-
vey and was carried out in cooperation with Stanford
University, Stanford, Calif; we gratefully acknowledge
the assistance of numerous people from both institu-
tions. In particular, Edwin L. Harp of the US. Geologi-
cal Survey aided in many aspects of the research, in—
cluding the initial fieldwork and modification of some
field and laboratory equipment. Deane Oberste—Lehn,
Richard Turnbull, and Douglas Yadon, all formerly of
Stanford University, showed us many interesting earth-
flow deposits in the field, and Douglas Yadon made ar-
rangements for digging the trench in the Cordelia,
Calif, area. Clay-mineral determinations were made by
John Sarmiento, and Atterberg—limit and grain-size de-
terminations were made by John Sarmiento, Michael
Bennett, and Joseph Heffern, all of the US. Geological
Survey. The manuscript was read by G. Wayne Clough,

A
A

EARTH—FLOW MORPHOLOGY, MATERIALS, AND SETTINGS 5

Bernard Hallet, Richard H. Jahns, and Ernest I. Rich
of Stanford University and by Robert W. Fleming,
Edwin L. Harp, and David J. Varnes of the U.S.
Geological Survey; their reviews and constructive criti-
cisms led to numerous improvements. Seena N. Hoose,
Gerald F. Wieczorek, Raymond C. Wilson, and T. Les-
lie Youd, all of the U.S. Geological Survey, provided
helpful insights during many stimulating discussions.

Ralph Azevedo, Charlotte Berberick, Mr. and Mrs.
Perry Davilla, Mr. and Mrs. A. A. Fields, the Paul
Hudner family, and Carl Leberer kindly gave permis-
sion for field studies on their land; without their gener-
ous cooperation this work would not have been possible.
We are particularly indebted to Mr. and Mrs. Perry
Davilla and Mr. and Mrs. A. A. Fields for permission to
conduct extensive studies on their property, and to Mr.
and Mrs. A. A. Fields for providing records from their
rain gage.

We especially thank Karen H. Keefer, who assisted
with the fieldwork on numerous occasions, drafted most
of the original figures, and contributed many helpful
suggestions and insights.

EARTH-FLOW MORPHOLOGY, MATERIALS,
AND SETTINGS

In parts of the California Coast Ranges near San
Francisco Bay the hillsides are dotted by thousands of
earth-flow deposits (fig. 2). Similar earth-flow deposits
occur in many parts of the world, and, in some regions,
earth flow is the dominant mass-movement process
(Radbruch and Weiler, 1963; Bailey, 1972; Dunkerley,
1976; Turnbull, 1976).

TOPOGRAPHY, CLIMATE, AND GEOLOGY IN AREAS
OF EARTH-FLOW OCCURRENCE

SAN FRANCISCO BAY REGION

In the San Francisco Bay region, earth flows occur
within a belt of northwest-southeast-trending mountain
ranges collectively called the “California Coast Ranges.”
Topography in these mountains ranges from gently roll-
ing hills to steep rugged ridges separated by narrow
canyons; elevations range from sea level to about 1,200
m.

Within the Coast Ranges, we identify six areas
where earth-flow deposits are abundant (fig. 1). In these
areas, the topography is characterized by rolling hills or
by ridges with broad rounded crests and moderately
steep flanks. Earth flows in these areas commonly come
to rest on or at the bases of the hillsides or ridges; few
travel more than a few meters on the gentle valley
floors. Many earth flows come to rest part way down ap-
parently uniform slopes.

 

 

 

FIGURE 2.—Typical earth-flow deposits. A, Near La Honda, Calif.
Photograph by G. F. Wieczorek, U.S. Geological Survey. B,
Near Berkeley, Calif. Photograph by G. K. Gilbert, U.S. Geolog-
ical Survey.

Inch'nations of the slopes on which earth flows occur
have been systematically measured in two parts of the
Orinda, Calif, area (fig. 1). In the central Orinda area,
earth flows occur only on slopes of 10° or steeper, and 99
percent of all earth-flow complexes are on slopes of 15°
or steeper (fig. 3A). In the area around Davilla Hill,
earth flows occur only on slopes of 8° or steeper, 99 per-
cent of earth-flow complexes are on slopes of 12° or
steeper (fig. 3B), and all earth-flow complexes on slopes
of less than 14° are associated with such locally steep
slopes as roadcuts, gully walls, or streambanks
(Turnbull, 1976). Although systematic studies of slope
inclination have not been carried out elsewhere in the
San Francisco Bay region, the slopes in the other five
areas with abundant earth flows appear to be compara-
ble to those in the Orinda area.

25

24

23

22

21

20

19

18

17

16

15

PERCENTAGE OF EARTH-FLOW COMPLEXES DEVELOPED
ON SLOPES WITH GIVEN INCLINATION IN PERCENT

EARTH FLOWS: MORPHOLOGY,

 

 

 

 

 

 

 

 

 

 

 

MOBILIZATION, AND MOVEMENT

 

 

I

 

 

 

 

0

B

5

I I I | | | I I
— — —I
— fl
.. _J
—— —«
_ _I
_ —¢
_. —I
I— — 14
,_. — 13
_ — 12
— — 11
— — 10
— — 9
— " — 8
I— — 7
_ -— 6
— — 5
— — 4
I— — 3
— E — 2
— — 1

II II J . 1 . . II 0
0 10 15 20 25 30 35 40 45 50

SLOPE INCLINATION, IN DEGREES

A

10

15

FIGURE 3.—Inclinations of slopes on which earth flows occur. Locations of areas shown in figure 1. A, Central Orinda, Calif., area.
Total number of earth-flow complexes included in study, 185. Data from Radbruch and Weiler (1963). B, Eden Canyon and vicin-

ity. Total number of earth-flow complexes included in study, 582. Data from Turnbull (1976).

 

 

EARTH—FLOW MORPHOLOGY, MATERIALS, AND SETTINGS 7

The San Francisco Bay region has a Mediterranean-
type climate with warm dry summers and cool rainy
winters. Annual precipitation varies greatly from place
to place; in the six areas containing abundant earth-flow
deposits, average annual precipitation ranges from 36
cm (in parts of the Bear Valley area) to 101 cm (in parts
of the Hicks Valley area) (Rantz, 1971). Nearly all pre-
cipitation falls as rain in the months October through
April, when storms generated over the Pacific Ocean
pass through the region; the amount and distribution of
precipitation during any given year depend on the
number and intensity of these storms. Thus, the annual
precipitation at any one place also varies considerably
from year to year. For example, in the 125-year period
from 1851 to 1976, annual precipitation in downtown San
Francisco ranged from 18 to 125 cm (San Francisco
Chronicle, 1976, p. 1). Severe storms can drop a large
proportion of the average annual precipitation within a
few hours or days. Temperatures in the San Francisco
Bay region rarely fall below freezing, and snow accounts
for a negligible amount of the total precipitation.

In the San Francisco Bay region, most or all earth
flows move only during the rainy, winter months; most

 

earth flows are mobilized during winters having above-
average precipitation; and earth flows are especially
common after severe, late-winter rains (Radbruch and
Weiler, 1963; Nilson and Turner, 1975; Oberste-Lehn,
1976; Turnbull, 1976; Keefer, 1977a). During the warm
dry summers in the San Francisco Bay region, most hill-
side materials, including earth-flow deposits, are thor-
oughly desiccated.

Table 1 lists bedrock types in the six areas contain-
ing abundant earth-flow deposits. Five of these six
areas are underlain primarily by poorly consolidated
Tertiary shale, mudstone, sandstone, and conglomerate.
Parts of the Hicks Valley area (fig. 1) are underlain by
heterogeneous tectonic melange with a matrix of highly
sheared shale; other parts of this area are underlain by
relatively unsheared sandstone and shale or by bluesch-
ist-facies metamorphic rocks.

Within each of these six areas, both the abundance
of earth-flow deposits and the bedrock lithology vary.
Detailed mapping in parts of the Orinda area (Radbruch
and Weiler, 1963; Turnbull, 1976), the Lomerias Muer-
tas area (Oberste—Lehn, 1976), the Cordelia area (Doug-
las Yadon, unpub. data, 1976) and the Hicks Valley area

TABLE 1.—Bedrock in areas containing abundant earth-flow deposits in the San Francisco Bay region

[See figure 1 for locations]

 

Area Bedrock

Reference

 

Orinda ------------ Orinda Formation (Miocene): Freshwater conglomerate, sandstone, siltstone,

and shale containing minor amounts of volcanic tuff, limestone, and lignite.

All the rocks are soft, poorly consolidated, and closely jointed; all the
clastic rocks have a clay matrix.

Cordelia ---------- Kreyenhagen Formation (Eocene): Silty clay—shale and poorly consolidated
micaceous te dspatfiic sandstone.

Unnamed formation (Jurassic or Cretaceous): Mudstone and shale containing
minor amounts of conglomerate and sandstone.

Bear Valley ------- Etchegoin group (Miocene and Pliocene):
sandstone, $1 tstone, and shale.

Poorly consolidated conglomerate,

Serpentinite (Jurassic or younger): Green, bluish-green, and bluish—gray;
perva51ve y sheared.
Cienega Valley---— Etchegoin Formation (Miocene and Pliocene): Thick-bedded sandstone con—
taining minor amounts of micaceous siltstone, brackish-marine and marine.

Nonmarine sedimentary rocks (Pliocene): Claystone, siltstone, and inter-
Beaaea friaS e sandstone.
Lomerias Muertas—- Purisima(?) Formation (Pliocene): Shale, claystone, siltstone, and friable
sandstone containing minor amounts of conglomerate and tuff.

Hicks Valley ------ Franciscan assemblage (Jurassic and Cretaceous): Blueschist-facies meta—
morphic rocks; tectonic melange with matrix of pervasively sheared shale
and sandstone; relatively unsheared shale and sandstone. Earth-flow
deposits are most abundant in clayey soils developed on melange.

Ham (1952), Hall
(1956), Robinson
(1956), Radbruch and
Neiler (1963),
Turnbull (1976).

Sims and others
(1973), Douglas
Yadon (unpub. data,
1976).

Wilson (1943).

Dibblee (1975).

Oberste-Lehn (1976).

Blake and others
(1974), Ellen and
others (1979),
Peterson (1979).

 

8 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

(Ellen and others, 1979; Peterson, 1979) showed that
most earth flows occur where the bedrock contains a
large proportion of clay.

OTHER REGIONS

In two areas outside the San Francisco Bay region,
the minimum inclinations of slopes on which earth flows
occur are rather low. In the Gros Ventre Mountains of
Wyoming, the minimum inclination is 2° (Bailey, 1972);
along the coast of southeastern England, the minimum
inclination is 4° (Hutchinson and Bhandari, 1971). Else-
where, slopes on which earth flows occur were reported
to range from 6° (Hadley, 1964) to 35° (Harden and
others, 1978). Minimum slope inclinations have been de-
termined from systematic studies in two areas: in the
Redwood Creek basin in northern California, where the
minimum inclination is 9° (Harden and others, 1978); and
in the Razorback Range in Australia, where the mini-
mum inclination is 10° (Dunkerley, 1976). Thus, the in-
clinations of slopes on which earth flows occur in most

 

 

other areas are comparable to those in the San Fran-
cisco Bay region.

Earth flows occur in areas where precipitation or
ground-water flow is sufficient to saturate surficial hill-
side materials at least intermittently. Previous studies
have shown that earth flows mobilize. during periods of
high precipitation (Campbell, 1966; Prior and Stephens,
1972; Hutchinson and others, 1974; Harden and others,
1978; Swanson and others, 1980) or during periods with
a surplus of precipitation over evapotranspiration
(Hutchinson, 1970).

Outside the San Francisco Bay area, earth flows
also occur in areas where the bedrock contains a large
proportion of clay. Bedrock types in areas of earth-flow
occurrence include stiff fissured clays, flysch, altered
and weathered volcanic rocks, tectonic melange, glacial
till, and poorly consolidated shale, mudstone, and
sandstone (table 2).

MORPHOLOGY

Figures 4A and 4B illustrate an idealized earth
flow, a model synthesized from our field observations

. u,“ '5 , . ,,

, 1,“ "plum, .
mwi

I|

.

,3 if’pl‘il’l‘.
a v.
[/i’llii ,1)”

\\\
w

 

 

FIGURE 4.—Idealized earth flows, showing features described in text. A, Surface features. B, Subsurface features.

 

 

EIXRTWI—FWLOVV'LiOltPIIOIJO(3Y} BLAUWERIAJJS,.AIJI)SIET111JGS

TABLE 2,—Bedrock in areas containing abundant earth-flow deposits outside the San Francisco Bay

Area Bedrock

Franciscan assemblage (upper Mesozoic): Unmetamorphosed argillaceous rocks

and slight y metamorphosed highly sheared sedimentary rocks.

Redwood Creek
basin, northern
California.

Franciscan melange (Mesozoic and early Cenozoic): Highly sheared clayey

$1 I ES Eone.

Van Duzen River
basin, northern
California.

Franciscan Group of Cooksley (1964) (upper Mesozoic): Interbedded black

shale, graywacke, and conglomerate; folded, faulted, and locally intruded
by serpentinite; deeply weathered to abundant clay.

Eel River basin,
northern
California.

Volcaniclastic rocks, lava flows, ash flows, and intrusive rocks (Tertiary),
Highly altered to form clay and saprolite; earth flows are particularly
common where soft volcaniclastic rocks are capped by harder lava flows and

(or) ash flows.

Western Cascade
Range, Oregon.

Soft shale and claystone (upper Mesozoic to Eocene) with interbedded sand-
stone and ST tstone containing minor amounts of conglomerate, limestone,
and coal; much of the shale and claystone is highly plastic and bentonitic.

Northwestern
Wyoming and
adjacent parts
of Montana and
Idaho.

 

Glacial till (Quaternary)

Gardiner area, Dacite breccia (Eocene) containing abundant highly plastic bentonite ———————————

Montana.
Landslide Creek Formation (Upper Cretaceous): Conglomeratic sandstone

interbedded with Bentonitic mudstone and claystone.

Clay, clay—shale, and coal (Pennsylvanian and Permian): Earth flows are

particu ar y abundant where rock creep has formed impermeable layers with
dips parallel to slopes; many earth flows form at sags in impermeable layers.

Upper Ohio River
Valley; Ohio,
West Virginia,
Pennsylvania,
and Kentucky.

Barbados —————————— Shale, mudstone, and sandstone (Tertiary) --------------------------------------

Cucaracha Formation (early Miocene): Sandy claystone, very poorly cemented

an very closely fractured, deeply weathered to soil containing abundant
kaolin, mica, chlorite, and iron oxides; rock disintegrates after oven

drying at 100°C and reexposure to water.

Panama Canal
Zone.

Dunedin district, Abbotsford mudstone (Cretaceous)

New Zealand.

Razorback Range, Nianamatta Group (Triassic):

Australia.

Decomposed claystone, chalk, basalt, marl (Triassic), and glacial till
(Quaternary): Earth flows occur on y where Triassic strata dip steep y
inland owing to disturbance by rotational slumping; steep dip allows

ground water to percolate along bedding planes.

Antrim coast,
northern
Ireland.

Coastal cliffs London clay (Eocene), Barton clay (Eocene), and Hamstead beds (Oligocene):
in southeastern Stiff fissured overconsolidated clay with interbedded 51 f and fine sand.

England.

Mo clay and Lille Baelt clay (Eocene): Stiff fissured overconsolidated

Coastal cliffs
_’clay containing some vo canic tuff and diatomaceous beds.

in Denmark.

 

Flyschoid strata of Zakopane facies (Paleogene): Claystone containing

Okolicné area,
lesser amounts ET sandstone; earth flows form where confined aquifers occur.

Czechoslovakia
Soft tectonically disturbed flysch, sandstone, and pelitic and argillaceous

 

Carpathian
Mountains, shale (Paleogene).
Czechoslovakia.

Tuff, agglomerate, coal, and poorly consolidated clay, silt, sand, and
gravel (Neogene).

 

Glauconitic mudstone ------------------------

Shale interbedded with lithic sandstone ----------

region

Reference

Harden and others
(1978).

Kelsey (1977).

Cooksley (1964).

Swanson and James
(1975), Swanson and
Swanston (1977),
Swanson and others
(1980).

Bailey (1972).

Fraser and others
(1969).

Sharpe and Bosch
(1942).

Prior and Ho (1972).

National Academy of
Sciences (1924).

Benson (1940).
Dunkerley (1976).

Prior and others
(1968).

Ward (1948),
Hutchinson (1970),
Hutchinson and
Bhandari (1971),
Barton (1973),
Bromhead (1978).

Prior (1977).
Fussganger and
Jadron (1977).

Zéruba and Mencl
(1969).

10 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

r

and from descriptions by other investigators. The model
shows the characteristic features of earth-flow morphol-
ogy; examples of these features are shown in figures 5
through 19. Not all earth flows contain every feature in
this model. Individual earth flows may also differ from
the model in shape: some have irregular boundaries that
conform to the local topography. Earth flows range in
size from bodies a few meters long, a few meters wide,
and less than 1 m deep to bodies more than 1 km long,
several hundred meters wide, and more than 10 m deep.
Many earth flows also occur as parts of larger earth-flow
complexes.

IDEALIZED EARTH-FLOW MORPHOLOGY

On the hillside above an earth flow, a main Scarp
and a zone of crown cracks mark the upslope boundary
of the area disturbed by mass movement. The crown
cracks break the earth in the crown region into blocks
(fig. 5A, B); from time to time, blocks become detached
and fall, slide, or slump down the main scarp. The main
scarp is steep; it is linear (fig. 5B) or arcuate (fig. 50) in
plan view. Main scarps associated with some large earth
flows are higher than 100 m. The main scarp bounds a
natural amphitheatre, called the zone of depletion. The
zone of depletion comprises the head and proximal part
of an earth flow as well as deposits of blocks that have
become detached from the crown. In some zones of de-
pletion, these deposits are slices of earth that have slid
on curved failure surfaces and rotated backward to form
slump deposits (figs. 50, 6A); in other zones of deple-
tion, the deposits are irregular blocks that result from
earth falls or earth block slides4 (fig. GB). Still other
zones of depletion are relatively free of fall, slide, or
slump deposits. The head of an earth flow can be adja-
cent to the main scarp or at some distance downslope, in
which case part of the basal shear surface is exposed on
the floor of the zone of depletion.

An earth flow itself is tongue or teardrop shaped
(fig. 2). The length of an earth flow, measured from
head to toe, is greater than its width, measured from
flank to flank; and its width is greater than its depth.
The flanks are parallel or diverge slightly downslope.
The average surface slope of an earth flow is gentler
than the slope of the adjacent, undisturbed part of the
hillside. In the zone of depletion, the earth-flow surface
is at a lower elevation than the ground on either flank.
Farther downslope, in the zone of accumulation, the
surface of an earth flow bulges above the undisturbed
ground on either flank. The distal margin of an earth
flow is a rounded, bulging toe (fig. 7).

‘In earth falls, grains or coherent blocks of earth descend slopes by free fall, bounding,
leaping, or rolling. In earth block slides, blocks of earth move downslope translationally along
distinct failure surfaces.

 

 

FIGURE 5.—Main scarps and crown cracks of earth-flow complexes.
Crown cracks break earth upslope from crown into angular
blocks. From time to time, blocks become detached and fall,
slide, or slump from crown, so that scarp retreats upslope. A,
Crown cracks adjacent to main scarp (upper left) of earth-flow
complex in Bear Valley area. B, Linear main scarp and crown
cracks of earth-flow complex in Bear Valley area. C, Arcuate
main scarp of earth-flow deposit near Fields Barn complex. In
zone of depletion at base of scarp are several blocks of earth that
have slumped from main scarp. Photograph by E. L. Harp, U.S.
Geological Survey.

A

EARTH—FLOW MORPHOLOGY, MATERIALS, AND SETTINGS 11

In longitudinal profile, an earth flow is concave up-
ward in the zone of depletion and convex upward in the
zone of accumulation; in overall profile it resembles a

u. »

l single sinusoidal wave. On some earth flows, this con-
' cavity and convexity are so pronounced that a closed to—
. pographic depression is formed near the head; on other

earth flows, the longitudinal profile is nearly linear ex-
" cept for a steepening of slope associated with the toe.
k. ,, Small local irregularities may be superimposed on this

profile, but the sinusoidal form itself appears to be near—
ly universal among earth flows. The near-universality of
this form indicates that it is not due to local site condi-
tions but is characteristic of earth-flow movement. The
sinusoidal profile probably results from thinning of the
earth flow near the head and thickening near the toe
during movement.

In the zone of accumulation, an earth flow is flanked
by lateral ridges that form in three different ways. One
type of lateral ridge is pushed up as a pressure ridge
(fig. 8); these pressure ridges have slickensided shear
surfaces on their outer flanks. A second type of lateral
ridge forms by overflow of earth-flow material onto the
adjacent ground surface (fig. 9). Ridges of a third type
form where a toe overrides the former ground surface
and the earth flow then stops moving, forming an earth-
flow deposit. If a later earth flow remobilizes part of this
deposit, some material adjacent to the flanks may be left
in place and form lateral ridges (figs. 10, 11). Lateral

 

 

FIGURE 6.-——Deposits in zones of depletion of earth flows. A, Slump
deposits in complex in Bear Valley area. 3, Small sod-covered
blocks are deposits of earth block slides in Fields Barn complex.

     
   

Striations plunging

10215“
A .
Ci

 

Ground undisturbed by
earflhflow movement

Earth llow

 

 

 

Lateral shear surlace
FIGURE 7.—Toe of earth—flow deposit in Bear Valley area. In plan
view, toe is rounded and curved concave sourceward; in profile,
toe has steep frontal slope. Toe has overriden former ground sur- FIGURE 8.—Cross section through pressure ridge flanking earth
face and is underlain by a surface of separation. Field notebook flow at Kirkwood Creek, Mont. (after Hadley, 1964). Note lateral
on toe is 22 by 30 cm. shear surface on outer flank.

 

12 ’ EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

 

FIGURE 9.—Lateral ridge (between arrows) formed by local over-
flow of material from earth flow at Melendy Ranch complex.
Earth—flow surface is light area to left of lateral ridge.

ridges of the second and third types have slickensided
shear surfaces on their inner flanks (fig. 12). The lateral
ridges flanking many earth flows are composite fea-
tures, made up of several segments of one or the other
of these types.

The distal margin of an earth flow is a steep, bulg-
ing, rounded toe (fig. 7). Even though this toe is not
confined laterally, it is only slightly wider than those
parts of an earth flow that are so laterally confined; only
minor spreading of the material near the toe occurs.
Some earth-flow toes override the ground surfaces
downslope from them (fig. 7); the boundaries between
the distal parts, or feet, of these earth flows and the
former ground surfaces are called smfaces of separa-
tion. Some toes contain recumbent folds (Blackwelder,
1912; Krauskopf and others, 1939); other toes are under-
lain by multiple shear surfaces arranged like imbricate
thrust faults (fig. 13). Both the folds and the multiple
shear surfaces indicate thickening and shortening in the
zones of accumulation.

Although an earth flow moves as a generally cohe-
rent body, it is disrupted by numerous cracks. One type
of crack, which occurs throughout the earth flow, is
caused by differential movement (fig. 14A). A second
type, which occurs in echelon sets adjacent to the
flanks, is caused by drag along the flanks (fig. 143). A
third type of crack is caused by desiccation (fig. 140);
these desiccation cracks resemble the cracks formed by
differential movement.

The main body of an earth flow is bounded by slick-
ensided basal and lateral shear smfaces (figs. 12, 15).
In the zone of depletion, the surface traces of lateral
shear surfaces are scarps connecting with the main

 

Lateral ridge

 

FIGURE 10.—Formation of lateral ridges when toe overrides former
ground surface and is later remobilized (third type mentioned in
text). A, Earth flow toe overrides former ground surface; earth
flow becomes inactive and forms an earth-flow deposit. B, Part
of earth-flow deposit is remobilized by another earth flow. C, As
second earth flow advances, part of deposit that was not re-
mobilized is left in place to form a pair of lateral ridges.

 

 

 

 

EARTH—FLOW MORPHOLOGY, MATERIALS, AND SETTINGS 13

 

FIGURE 11.—Earth—flow deposit in Cordelia, Calif, area,
showing lateral ridge (dashed outline) formed as illustrated
in figure 10. Deposit is to right of lateral ridge. Ground to
left of lateral ridge is undisturbed by earth-flow move-
ment. Notebook is 22 by 30 cm.

 

FIGURE 12.—Slickensided lateral shear surface on inner flank of
lateral ridge at Melendy Ranch complex. Notebook is 22 by 30
cm.

 

Bands of blue—gray earth—flow material

Basalshear
surfaces

 
 
 
    

Brown earth—flow material 9/
5)» ,9 B h
9» 9, eac

ux>9¥ L\>—‘) xx“? .7? :v. 3;":1?

       

V
\)

  

 

Basal shear surface

Blue-gray London Clay

FIGURE 13.—Multip1e basal shear surfaces underlying earth-flow
toe at Beltinge, England (after Hutchinson, 1970), indicating
thickening and shortening in zone of accumulation. Reprinted
with permission of J. N. Hutchinson.

scarp (fig. 16A), or open cracks within the amphitheatre
(fig. 163). In the zone of accumulation, lateral shear sur-
faces are commonly scarps bounding lateral ridges (figs.
8, 1‘2). Subsurface geometries of basal and lateral shear
surfaces are described in the next section.

SUBSURFACE FEATURES: BASAL SHEAR SURFACES,
LATERAL SHEAR SURFACES, AND ASSOCIATED
ZONES OF DEFORMATION

During our study, we examined basal and lateral
shear surfaces and associated features in a trench into
an earth-flow deposit in the Cordelia, Calif, area, in a
gully eroded through an earth—flow complex near Davilla
Hill, and in two pits excavated into the Davilla Hill
earth-flow complex.

The trench in the Cordelia area was excavated
through a lateral ridge and part of the zone of accumula-
tion of the earth-flow deposit. The long axis of the
trench was perpendicular to the direction of movement
of the earth flow; thus, the trench provided a transverse
cross section through the deposit (fig. 17A). Figure 178
is a map of the trench wall, showing traces of the shear
surfaces and associated features. The basal shear sur—
face is pervasively slickensided (fig. 15A). The trace of
this shear surface, though generally linear, contains one
major asperity (B, fig. 178) and two subsidiary shear
surfaces (A and B, fig. 178). The basal shear surface is
underlain by a poorly consolidated yellowish-brown
mudstone broken by fractures spaced a few millimeters
apart. Whereas most of the earth-flow deposit is com—
posed of dark-brown silty clay, the layer immediately
above the basal shear surface is composed of softened
yellowish-brown mudstone. The boundary between the
mudstone and the clay, though distinct, is highly convo—
luted (fig. 18); the softened-mudstone layer also contains
rounded lumps of the clay. The convoluted boundary
and the inclusion of lumps of the clay in the softened
mudstone indicate that significant internal deformation
had taken place in the zone adjacent to the shear sur-
face.

14 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

 

FIGURE 14.—'I‘ypes of cracks disrupting earth flows. A, Cracks
caused by differential movement at Melendy Ranch complex.
Such cracks occur throughout most or all earth flows and earth-
flow deposits. Photograph by E. L. Harp, US. Geological Sur-
vey. B, Echelon cracks along flank of earth flow at Davilla Hill
complex. Such cracks are caused by drag along flanks. Lens cap
is 5 cm in diameter. C, Dessication cracks in earth-flow deposit
at Davilla Hill complex. Cloth flag in center foreground (arrow)
is 5 by 10 cm.

 

A steeply dipping lateral shear surface joins the
basal shear surface at a sharp bend (C, fig. 173); in
transverse cross section, therefore, this earth-flow de-
posit is approximately rectangular. Near the ground
surface, the lateral shear surface splits into two bran-
ches. One branch forms the outer flank of a lateral ridge
(D, fig. 173); the other branch may form the inner flank
of this lateral ridge, but this relation could not be deter-
mined because most of the lateral ridge was covered by
spoil from the trench. No slickensides were observed on
the lateral shear surface.

In a gully near Davilla Hill, two earth-flow de—
posits, one on top of the other, are exposed in longitudi-
nal cross section (fig. 19). The basal shear surface un-
derlying the upper deposit is broadly planar. The appar-
ent dip of the shear surface decreases gradually from 17°

 

FIGURE 15.—Slickensided shear surfaces. A, Basal shear surface
exposed on floor of trench in Cordelia, Calif. area. Slickensides
trend parallel to inferred direction of earth-flow movement.
Slickensided area is approximately 50 cm on a side. View verti-
cally downward. B, Lateral shear surface at Davilla Hill site. Di-
rection of earth-flow movement is toward left. Slickensides are
on margin of earth flow which is raised above adjacent, immobile
material visible along lower edge of photograph. View obliquely
downward.

 

under the head to horizontal at a point about 1 m up—
“ slope from the toe; downslope from there, the dip re-
verses, and, under the toe itself, the shear surface dips
‘ upslope about 15°. This shear surface is underlain by a
f ' layer, 20 to 30 mm thick, of very soft sheared silty clay.
‘ The shear surface underlying the lower deposit is under-

i, lain by a zone of sheared material, 3 mm thick. Both
‘ ' shear surfaces are slickensided.
L, ,, Two pits dug at the Davilla Hill site exposed areas,
‘ a few centimeters on a side, of the slickensided basal
1 ’ shear surface of one earth-flow deposit. We did not ob—
, serve any structures indicative of deformation in mate-
rial adjacent to this shear surface.
L i
, )-
,5»
L s

 

 

FIGURE 16.—Lateral shear surfaces in zones of depletion of. earth

flows at Davilla Hill complex. A, Scarp forming lateral shear sur-

~ 9 face. Fresh crack (between arrows) at base of scarp is boundary

between active earth flow and immobile material. Scarp is ap—

«A proximately 0.5 m high. B, Trace of lateral shear surface as an

open crack. Lateral shear surface is parallel to and approxi-

mately 20 cm from base of scarp bounding zone of depletion.

Scarp is visible on other side of lens cap (5 cm in diameter) from

lateral shear surface. Direction of earth-flow movement is toward
right. View obliquely downward.

 

EARTH—FLOW MORPHOLOGY, MATERIALS, AND SETTINGS 15

 

Differentially
displaced cracks

Ground

   

 

 
 
  

surfaces

Dark-brown
silty clay

Trench
boundary

    

    
 

Dark-brown
,r’g’o’hene‘d‘\ silty clav /
,4 mudslone ‘ —————— ”

     
  
   

Yellowishbrown

m
Basal shear udstone

surfaces

1 METER B

FIGURE 17.—Trench cut through earth-flow deposit in Cordelia,
Calif., area. A, Location and orientation of trench in relation to
earth-flow features. B, Trench wall, showing earth-flow basal
and lateral shear surfaces, boundaries between different mate-
rials, and differentially displaced cracks (sense of displacement
indicated by arrows). View upslope.

1 METER

 

FIGURE 18.~Convoluted contact (dashed line) between softened
yellow-brown mudstone (lighter material) and dark-brown silty
clay (darker material) in sample taken from trench in Cordelia,
Calif., area (see fig. 17). Slickensided basal shear surface is visi-
ble on left boundary of sample. Arrow points toward original top
of sample. Pencil is 15 cm long. Photograph by G. F. Wieczorek,
U.S. Geological Survey.

16 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

EARTH-FLOW COMPLEXES

On many hillsides, earth-flow deposits occur in long
narrow channels (pls. 2, 3; fig. 20; earth-flow complex 1
in fig. 21). At the upslope margin of each channel is a
scarp, and at the downslope margin is at least one earth-
flow toe. Within the channels, the topography appears
to be a chaotic jumble of bulges, depressions, and blocks
separated by scarps and cracks. Such topography led
Blackwelder (1912, p. 490—491) to describe a feature of

1 METER

 

Earth-flow head

 
 

1 METER

Ea rth—flow toe

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\ features not visible
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FIGURE 19.—Wall of gully near Davilla Hill complex. Two earth-
flow deposits, underlain by basal shear surfaces, are exposed in
longitudinal cross section. Dashed lines indicate contacts between
different materials; solid lines indicate basal shear surfaces. 1,
earth-flow deposit 1, dark-yellowish-brown silty sand; 2, earth-
flow deposit 2, dark-yellowish-brown silty sand; 3, earth-flow de-
posit 2, light-olive-gray silty sand; 4, dark-yellowish-brown silty
sand; 5, moderate-brown silty sand; A, basal shear surface un-
derlain by 20 to 30 mm of very soft silty clay; B, basal shear sur-
face underlain by 3-mm-thick shear zone.

 

FIGURE 20.—-Earth-flow complexes occupying long narrow channels
on a hillside near Davilla Hill complex. Arrows mark upslope
margins of channels, each of which contains several earth-flow
deposits.

 

this kind as a “long glacier-like tongue” with a surface
composed of “orderless humps and hollows.”

This and many other published descriptions imply
that all the material in a channel is part of a single large
earth-flow deposit. Our examination of several tens of
these channels and detailed mapping of these features at
three sites, however, show that most such channels con-
tain several smaller earth-flow deposits (see section en-
titled “The Davilla Hill Earth-Flow Complex”; pls. 2, 3;
Aboim-Costa and Stein, 1976; Keefer, 1976; 1977a; Spain
and Upp, 1976). The Hidden Valley earth-flow complex
(pl. 2) contains at least 13 distinct earth-flow deposits,
and the Davilla Hill earth-flow complex (pl. 3) at least
34. Alternating zones of depletion and accumulation,
formed by numerous individual earth flows, are shown
on the cross section in plate 2. Deposits that are com-
pletely preserved are similar in form to those shown in
figures 2 and 4. Many deposits, however, are highly dis-
rupted by cracks or are partially destroyed by erosion,
by renewed earth-flow movement, or by other mass-
movement processes. In some channels, earth-flow de-
posits are piled one on top of another, piggyback fash-
ion; in others, they are separated by zones containing
the deposits of other types of mass movement.

Most earth-flow deposits, therefore, occur as parts
of earth-flow complexes. A given earth-flow complex
may contain several earth-flow deposits, as well as other
types of mass-movement deposits. Some earth-flow
complexes consist of a single sinuous channel (pls. 2, 3),
others of several coalescing channels (fig. 20; earth-flow

 

FIGURE 21.—Two earth-flow complexes in Cordelia, Calif., area.
Earthvflow complex 1 consists of several earth-flow deposits in a
network of sinuous coalescing channels. Earth-flow complex 2
consists of a broad slump deposit (S) and several earth-flow de-
posits (E) adjoining slump deposit. Note road near base of hill-
side for scale.

 

EARTH—FLOW MORPHOLOGY,

complex 1 in fig. 21); still others are irregular in outline
and blanket broad areas with hummocky, disrupted to-
pography (earth-flow complex 2 in fig. 21; fig. 22). Some
earth flow complexes have surface areas of several tens
of square kilometers (Bailey, 1972).

The channels of earth-flow complexes are flanked
by a combination of scarps and lateral ridges. In many
places, the lateral ridges are made up of several seg-
ments arranged in echelon. These echelon patterns
document the passage of several earth flows through
each channel; each subsequent earth flow came to rest
farther downslope, as illustrated in figure 10.

At any given time, an earth-flow complex may con-
tain one or more active earth flows moving through any
part of the complex. In many earth-flow complexes,
earth flows are intermittently active for several years
(Radbruch and Weiler, 1963; Campbell, 1966; Hutchin-
son, 1970; Hutchinson and others, 1974) or even several
centuries (Crandell and Varnes, 1961; Swanson and
Swanston, 1977). Within complexes, earth flows com-
monly are formed out of older earth-flow deposits; the
younger earth flows, however, do not have the same
boundaries as the older deposits. New scarps and lateral
shear surfaces commonly cut across the older bound-
aries, and earth flows formed out of older earth-flow
material leave behind irregular remnants of the older
deposits.

MATERIALS

Earth flows occur in various fine materials, includ-
ing residual soils, colluvial soils, old earth-flow deposits,
and deposits of other types of landslides. Sand-, silt-,
and clay-size grains predominate in earth-flow mate-
rials, and, in most earth flows, silt and clay make up

FIGURE 22.—Earth-flow complex near Pleasanton, Calif., showing

 

hummocky topography typical of large earth-flow complexes.

MATERIALS, AND SETTINGS 17

more than 50 percent of the matrix (table 3). Some earth
flows also incorporate larger rock fragments, lumps of
stiff clay, or such miscellaneous debris as tin cans or
sheep carcasses. The earth flows and earth-flow de-
posits that we examined in the San Francisco Bay re-
gion are composed of silty clay, clayey silt, or silty sand
(table 3). The silty sand of the Melendy Ranch earth-
flow complex is the coarsest material we observed, and
the silty clay of the Cycle Park earth-flow complex is the
finest (table 3).

The clay mineralogy of earth-flow materials varies
from site to site (table 3). In the Davilla Hill and Fields
Barn earth-flow complexes in the Orinda, Calif, area, il-
lite and chlorite are the predominant clay minerals, and
in the Cycle Park earth-flow complex montmorillonite
and illite; in the Melendy Ranch earth-flow complex, talc
and serpentine are the only clay minerals present.
Montmorillonite predominates in the clay of earth flows
at Slumgullion, Colo. (Crandell and Varnes, 1961), in
the Gros Ventre Mountains of Wyoming (Bailey, 1972),
and in the Lomerias Muertas area (Oberste-Lehn, 1976).
Earth flows at Minnis North, northern Ireland, contain
large proportions of montmorillonite and illite and minor
amounts of kaolinite (Prior and others, 1971). Earth
flows on the Caribbean island of Barbados contain a
mixture of kaolinite, illite, chlorite, and montmorillonite
(Prior and Ho, 1972).

The wide range in the plasticity of earth-flow mate-
rials, as measured by the Atterberg liquid and plastic
limits and plasticity index, reflects variations in both
clay-mineral content and grain-size distribution (fig. 23;
table 3). On the plasticity chart of the Unified Soil Clas-
sification System, most earth-flow materials plot near
the A-line, which separates plastic from nonplastic soils
(fig. 23). Of the 19 earth-flow materials plotted (fig. 23),
8 are CH soils (clays with high plasticity), 5 are CL soils
(clays with low plasticity), 5 are MH soils (silts with
high plasticity), and one is a CL-ML soil.

Active earth floWs contain a significant amount of
entrained water; reported water contents5 range from
27 to about 60 percent (table 3). Earth-flow mobilization
is accompanied by an increase in water content (Hutch-
inson, 1970; Prior and others, 1971; Keefer, 1977a) and,
at many sites, by softening of the material (Ward, 1948;
Prior and others, 1968; Hutchinson, 1970; Hutchinson
and Bhandari, 1971; Prior and others, 1971; Hutchinson
and others, 1974; Keefer, 1977a). Most active earth
flows are so soft that they will not support a person’s
weight; others contain a stiff crust over a soft interior,
and still others are relatively stiff throughout. Earth

“Water content = 100(weight of water)/(weight of dry soil).

MORPHOLOGY, MOBILIZATION, AND MOVEMENT

EARTH FLOWS

18

 

 

 

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22 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

 

8°IIIIIIIITII

70—
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CL

   
 

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20 —

PLASTICITY INDEX, IN PERCENT

 

 

. I I I L n L
0 10 20 30 40 50 60 7O 80 90 100 110 120

LIQUID LIMIT, IN PERCENT

 

FIGURE 23.—Plasticity indexes and liquid limits of earth-flow mate-
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uid limit-plastic limit. Data points (see table 3 for references): 1,
Davilla Hill; 2, Fields Barn; 3, Cycle Park, silty clay; 4, Cycle
Park, clayey silt; 5, Gros Ventre Valley, sample 1; 6, Gros Ven-
tre Valley, sample 2; 7, Gros Ventre Valley, sample 3; 8, Be]-
tinge, England; 9, Isle of Sheppey, England; 10, Minnis North,
northern Ireland, earth flow 1; 11, Rojne Klint, Denmark; 12,
Rosnaes, Denmark; 13, Helgenaes, Denmark; 14, Castle Hill,
England; 15, Bruce Vale, Barbados; 16, Saddleback, Barbados;
17, Bathsheba, Barbados; 18, Bellplaine, Barbados; 19, Minnis
North, northern Ireland, earth flow 2. A-line, separates plastic
from nonplastic soils; CH, clays with high plasticity; CL, clays
with low plasticity; MH, silts with high plasticity; ML, silts with
low plasticity; 0H, organic soils with high plasticity; 0L, organic
soils with low plasticity.

flows, in general, are softer and more fluid than mate-
rial in earth slumps or earth block slides, as defined by
Varnes (1978), and stiffer and less fluid than material in
debris flows and rapid earth flows such as those de-
scribed by Johnson (1970) and Varnes (1978), respec-
tively.

The shear strengths6 of earth-flow materials are af-
fected both by material composition and by the presence
of planes of weakness, including cracks and slickensided
shear surfaces. Drained7 shear strength along cracks
and shear surfaces is much lower than the peak (upper
bound) shear strength of a given earth-flow material
(Hutchinson and others, 1974; Keefer, 1977a; Prior,
1977). Shear strengths along many slickensided shear

l‘Soil in an earth flow is assumed to be a Coulomb material, for which Temghi’s theory
of effective stress is valid. Thus

1 = (on—unan$ +6,

where 'r = the effective shear strength on s plane within the soil, 0,. = the normal stress on
the plane, at = the pore-water pressure, 3) = the angle of internal friction, E = the cohesion.

7Under drained conditions, excess pore-water pressures generated by strain are allowed
to dissipate by free movement of pore water, under undrained conditions, these pore~wster
pressures are not allowed to dissipate.

 

surfaces are nearly equal to the residual (lower bound)
strength of the material (Hutchinson and Bhandari,
1971; Fussganger and Jadron, 1977; Prior, 1977). Under
undrained7 conditions, field-vane shear strengths meas-
ured at two sites—Davilla Hill and Beltinge, England—
showed that the earth-flow materials have low sen-
sitivities.8 At Beltinge, the average sensitivity is only
slightly greater than 1.0 (Hutchinson, 1970); and at
Davilla Hill, 2.2.

THE DAVILLA HILL EARTH-FLOW COMPLEX

In this section we describe field, laboratory, and
analytical investigations of the Davilla Hill earth-flow
complex that were undertaken to study the earth-flow
process at a representative site. Field studies were con-
ducted from September 1974 to February 1978. During
February 1975, four earth flows mobilized within the
complex, and movements continued for several weeks.
During these several weeks, surface and subsurface
movements and ground-water levels were measured,
samples were obtained for laboratory testing, and the
undrained strength of the earth-flow material was meas—
ured with a field vane. Morphologic features of the
earth-flow complex were mapped in June and July 1975.

Beginning in August 1975 and continuing through
February 1978, pore pressures and ground-water levels
were monitored with portable tensiometers and open-
standpipe piezometers; sampling also continued during
this time. Two earth flows were active in the complex
during January and, possibly, February 1978, but no
earth flows were active during the exceptionally dry
winters of 1975-76 or 1976—77. During January and Feb-
ruary 1978, studies consisted of visual observations and
of monitoring of water levels and pore-water pressures.

Laboratory tests were carried out to measure
strength parameters for use in analysis of mobilization;
to measure changes in water content, saturation, and
volume associated with mobilization; and to determine
grain sizes, Atterberg limits, clay mineralogies, and
consolidation properties of the earth-flow material.
Earth-flow mobilization at Davilla Hill was analyzed
using a modified form of the slope-stability analysis de-
veloped by Morgenstern and Price (1965, 1967).

DESCRIPTION
LOCATION AND SETTING
The Davilla Hill earth-flow complex is near the
head of Eden Canyon in the southern part of the Orinda,

BSensitivity = (peak undrained shear strength) / (undrained shear strength of remolded
material).

THE DAVILLA HILL EARTH—FLOW COMPLEX

 

 

 

 

37° r.
42’
30"

 

Oakland

 
 

San Francisco 0

 
 
     

5’" J° 'Area of

figure 34 ‘ .

   

 

O
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1

 

    

 

 

 
 

 

 

 

 

INDEX MAP
0 1 KILOMETER
CONTOUR INTERVALS 20 AND 40 FEET
EXPLANATION
Folds4howing axial trace and di-
Orinda Formation (Miocene)—Non- rection of plunge

marine conglomerate, sandstone.

siltstone, and shale containing f—3— Anticline

minor limestone. tuff, and lig- .

nite 4—*— Synclrne

4—fi— Overturned syncline
Briones Formation (Miocene)—

Marine sandstone and shell Strike and dip of bedding

reefs
—— Contact—Dashed where indefinite 60 lnClined
T— Fault—Dashed where indefinite. fig Overturned

U, upthrown side; D, down-
thrown side

FIGURE 24.—Geologic map of Eden Canyon area (from Turnbull, 1976), showing location of Davilla Hill earth—flow complex.
Base from US. Geological Survey, Dublin and Hayward quadrangles, 124,000.

23

24 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

Calif, area (pl. 1; figs. 1, 24). In Eden Canyon, earth
flows are more abundant than any other type of mass-
movement phenomenon; an area of about 6 km2 contains
several hundred earth-flow complexes (pl. 1).

These earth-flow complexes are in colluvial and re-
sidual soils that mantle long northeast-trending ridges;
the soils are silty clay, clayey silt, and sandy silt, gener-
ally 0.5 to 1.5 m thick but more than 4 m thick in the
larger hillside reentrants and valleys tributary to Eden
Canyon. Topography on the ridges is rounded, subdued,
and hummocky owing to the pervasive degradation of
slopes by earth flows. The earth flows occur on slopes
with inclinations ranging from 8° to 45° (fig. 3B).

Bedrock underlying the area containing abundant
earth flows belongs to the Orinda Formation, a se-
quence of Miocene freshwater conglomerate, sandstone,
siltstone, and shale, and minor amounts of volcanic tuff,
limestone, and lignite (Hall, 1956; Robinson, 1956;
Turnbull, 1976). The rocks are soft, poorly consolidated,
and closely jointed, and all the elastic rocks contain a
matrix of clay (Turnbull, 1976).

Orinda Formation that crops out in a gully a few
meters south of the Davilla Hill earth-flow complex is a
medium—gray calcareous shale containing scattered egg-
size concretions and ocher staining on the joints. The
shale is soft, highly fractured, and highly weathered;
bedding is lenticular and deformed by small-scale fold-
ing. In a bulldozer cut about 100 m west of the complex,
the rocks are thin-interbedded shale, siltstone, and very
fine grained to fine—grained sandstone. All these rocks
are poorly consolidated, broken by closely spaced ortho-
gonal joints, and weakly cemented by calcite.

Auger borings indicate that the rock beneath most
of the earth-flow complex consists of interbedded
mudstone and shale and small amounts of fine-grained
sandstone. Auger borings on the steep slopes a few met-
ers east and north of the complex indicate that the rock
there contains more sandstone than the rock underlying
the complex.

As in the rest of the San Francisco Bay region, the
climate at the Davilla Hill site is characterized by cool
rainy winters and warm dry summers. Temperatures
rarely fall below freezing in the winter; in the summer,
temperatures occasionally exceed 38°C (100°F). Nearly
all precipitation falls during the months October through
April. A private rain gage, maintained by Mr. and Mrs.
A. A. Fields, is located 0.7 km southwest of the Davilla
Hill earth-flow complex. During the period 1967-77, the
average annual precipitation there was 59.9 cm.9 The
minimum annual precipitation during this period was

9Annual precipitation is measured from July 1 of one calendar year to June 30 of the next
year.

 

29.4 cm in 1975-76; the maximum was 94.8 cm in 1972-
73. Figure 25 shows precipitation records for the time
during which field studies were carried out at Davilla
Hill.

MORPHOLOGY AND HISTORY OF MOVEMENT

Morphologic features of the Davilla Hill earth-flow
complex were mapped using a planetable and alidade
(pl. 3). From the planetable map, an interpretative map
was made outlining individual earth-flow deposits within
the complex (pl. 3). The following description refers to
the earth-flow complex as it appeared in June and July
1975.

The earth-flow complex occupies a bowl-shaped
reentrant in the ridge that forms the east wall of Eden
Canyon. The slopes above the earth-flow complex have
an average inclination of 27°. These slopes are cut by
two gullies, one of which contains an earth-flow deposit
(CS, pl. 3). The reentrant narrows in the reach occupied
by the earth-flow complex. The average inclination of
the surface of the complex is 15°. Downslope from the
complex, the reentrant empties out onto a more gently
sloping apron of colluvial material. After major storms,
intermittent streams carry runoff across the surface of
the complex, through a gully in the colluvial apron, and
into the creek in Eden Canyon. A wire fence crosses the
complex about midway along its length (pl. 3). On both
sides of this fence a mixture of grasses and thistles is
present on the earth-flow deposits. Upslope from the
fence, vegetation on slopes undisturbed by mass move-
ment consists mainly of grasses (fig. 26A). Downslope
from the fence, most of the complex is shaded by
eucalyptus, bay laurel, and oak trees (fig. 268).

The Davilla Hill earth-flow complex is approxi—
mately 140 m long and 25 m wide and has 45 m of topog-
raphic relief from the uppermost scarp to the lowermost
earth-flow toe. In plan view, the complex is shaped like
a map of Italy with a broad crown region formed by
earth-slide10 features and a “boot” formed by two earth-
flow toes (pl. 3). The earth-slide features consist of
scarps as high as 1 m, scars as deep as 1 m, and deposits
that form ridges and hummocks as high as 1 m. Im-
mediately downslope from these deposits, the slope in-
clination decreases from 27° to 11°. The gently sloping
area is underlain by a partially preserved earth-flow de-
posit (CZa, pl. 3), the distal boundary of which is a linear
scarp connecting with extension cracks at either end.
Downslope from the scarp, two earth-flow deposits (B7a
and B8, pl. 3) occupy a channel bounded by two inward-

10An earth slide is a landslide involving fine soil that moves rapidly and with much inter-
nal defamation along a planar shear surface.

25

THE DAVILLA HILL EARTH—FLOW COMPLEX

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26 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

 

FIGURE 26.—Davilla Hill earth-flow complex. A, Upslope from wire
fence. Surface of slump deposit is visible in foreground. Arrows
point to area of earth-slide features in crown region and to gully
containing earth-flow deposit C3 (see pl. 3). Much of complex up-
slope from slump deposit was mobilized by earth flows in Janu-
ary 1978. B, Downslope from fence. Prominent lateral ridges as
high as 0.6 m flank this part of complex. In foreground, complex
is 4 m wide. Parts of earth flows 2, 3, and 4, active in 1975, are
outlined by dashed lines.

facing scarps capped by lateral ridges. These lateral
ridges are remnants of an earth flow (Cl, pl. 3), larger
than any of the preserved deposits, that had its source
in the area currently covered by earth-slide scars. The
two earth-flow deposits in the channel spill out onto the
surface of a slump deposit (pl. 3; fig. 26A).

The slump deposit has a surface slope that is rela-
tively gentle and locally reversed. On the left flank11 of
the deposit is an inward-facing scarp covered with stria-
tions plunging downslope at angles of 30° to 36° (pl. 3);
to the right of the deposit is a long narrow depression

”Right and left flanks are relative to an observer facing downslope.

 

containing two earth-flow deposits (B6 and B7b, pl. 3).
The distal boundary of the slump deposit, a jumbled
mass of scarps and small landslide blocks, was the crown
of one earth flow active in 1975; the other three earth
flows active in 1975 mobilized within old earth-flow ma-
terial farther downslope.

The slump deposit and much of the earth-flow com-
plex downslope from it are contained in a channel
bounded by a combination of inward-facing scarps and
lateral ridges (fig. 26A). Flanking the slump deposit it-
self, however, are several earth-flow deposits that lie
outside this main channel (A1, A2, A3, A5, A6, B1, B2a,
B3b, and B4, pl. 3). The two largest of these (A1 and
B4, pl. 3) have piled up behind obstructions consisting of
the wire fence, a large eucalyptus tree, and a clump of
shrubs. All these deposits are older than the inward-fac-
ing scarps bounding the main channel.

Downslope from the slump deposit, the complex is
composed entirely of material transported by earth
flows. The earth-flow deposits that are completely pre-
served are long and narrow; they are bounded by dis-
tinct main scarps, lateral shear surfaces, and bulging
toes. Whereas their boundaries are locally irregular,
most deposits are tongue or teardrop shaped. Many de-
posits, however, are only partially preserved; these re-
mnants are highly irregular in shape.

The earth-flow deposits as well as the active earth
flows observed in 1975 and 1978 contain numerous inter-
nal scarps, desiccation cracks, and cracks caused by dif-
ferential movement; the larger cracks are several cen-
timeters wide and about 1 m deep. Some scarps and
cracks in active earth flows existed before the move-
ment began; other scarps and cracks formed while the
earth flows were moving.

Lateral ridges bordering the lower reaches of the
complex are composite features, formed by numerous
earth flows. These ridges, as high as 0.6 m, with steep
flanks and rounded crests (pl. 3; fig. 26B), were formed
by overflow of earth-flow material out of the main chan-
nel. In one locality (A7, pl. 3), the body of overflowing
material divided and flowed around a tree. In another
locality (intersection of A1 and fence, pl. 3), a bend of
70° in a lateral ridge indicates an abrupt change in the
movement direction of the earth flow that formed the
ridge. Downslope from this bend, a lateral shear surface
was formed only a few centimeters inside the margin of
an older earth-flow deposit (A4, pl. 3). Thus, even
though the earth flow bounded by this lateral shear sur—
face (A6, pl. 3) remobilized nearly all the older deposit,
the boundaries of the younger earth flow do not coincide
with those of the older deposit.

0n the right flank of the complex, three lateral-
ridge segments (A6, A7, and A9, pl. 3) form an echelon.
These segments were formed by a series of three earth

 

THE DAVILLA HILL EARTH—FLOW COMPLEX 29

wetter than at other times. In the active earth flows,
cracks were generally filled or nearly filled with water;
elsewhere in the complex, the surficial material was
wet, although the cracks generally did not contain so
much water. A network of small streams carried runoff
across the ground surface.

Figure 27 plots the displacements of all survey
stakes on the active earth flows as a function of time.
The calculated velocities of the earth flows (slopes of
lines in fig. 27) varied over time and ranged from 1 to 39
cm/d. These velocities, however, were calculated from
measurements taken several days apart and may not
represent the true maximum or minimum velocities.

Surface movements on earth flow 1 were monitored
by stake lines 2 and 2A”. The distribution of displace-
ments indicated that most movement took place by
shear on lateral shear surfaces or by differential dis-
placement near the lateral shear surfaces; less than a
third of the total displacement was due to internal defor-
mation away from these boundary zones (fig. 28A). Sur-
face displacements on earth flow 2 were monitored with
a single survey stake (fig. 27), and so the spatial distri-
bution of displacements was not determined. Although
displacements of stakes were not measured in the verti-
cal plane, observation of displacements of objects near
the lateral shear surfaces indicated that movement of
earth flows 1 and 2 was translational. Movement on
earth flows 1 and 2 ceased between April 23 and June
19, 1975.

Surface displacements on earth flow 3, monitored
by lines 3 and 3A (figs. 27, 28B, 28C), indicated that
somewhat more internal deformation was taking place in
this earth flow than in earth flow 1 (fig. 28), probably
owing in part to differences in boundary conditions.
Earth flow 1 had subparallel flanks where the survey
line crossed it and was completely confined by flanking
scarps; earth flow 3 had flanks that flared outward in
the downslope direction, and material overflowed along
the left flank in the zone of accumulation. Downslope
from the point where this overflow occurred, forward
tilting of the survey stakes during movement indicated
that material was being rolled forward as if it were part
of a caterpillar tractor tread. Two stakes were eventu-
ally washed away by an intermittent stream eroding the
left flank of the earth flow (fig. 28B).

During movement, the main scarp of earth flow 3
grew in height from a few centimeters to slightly more
than 2 m, owing to subsidence of the earth-flow surface
in the zone of depletion. Striations near the top of one
lateral shear surface, however, had plunges parallel to
the surface slope; these features indicated that the in-
itial movement was translational. Earth flow 3 stopped

12Line 2A was installed after the stakes in line 2 were trampled by livestock.

 

moving between March 21 and 31, 1975. Between these
two dates, a gully was cut through the earth flow, and
the drainage it provided probably caused the cessation
of movement.

Material in earth flow 4 softened significantly be-
tween March 8 and 15, 1975; from this softening we in-
ferred that the earth flow had begun to move. No sur-
vey stakes were placed on this earth flow to confirm its
activity, however, and any movement that did occur
could not be detected by visual inspection.

Fresh striations on the scarp on the left flank of the
slump deposit (pl. 3) indicated that this landslide, too,
moved during the winter of 1974-75, although its dis-
placement was not surveyed, and so neither the amount
nor the rate of diSplacement is known.

CORRELATIONS OF PRECIPITATION WITH
MOBILIZATION AND RATE OF MOVEMENT

The interval of earth-flow activity in 1975 coincided
with a period of high precipitation at Davilla Hill (fig.
25). Total precipitation during the winter of 1974-75 was
above average, and more than two-thirds occurred in
February, March, and April of 1975 (fig. 25), the months
when most or all of the earth-flow movement took place.

Earlier in the winter, precipitation totaling 10.9 cm,
between December 28, 1974, and January 9, 1975, did
not cause any earth flows to mobilize. On the first 4 days
of February 1975, an additional 8.8 cm of rain fell; yet no
earth flows were active on February 4. However, 8.5
cm of additional precipitation occurred between Feb-
ruary 4 and 21, and on February 21 three earth flows
were active.

When the earth flows mobilized, the total precipita-
tion after February 1 was between 8.8 and 17.3 cm, and
the total annual precipitation between 29.9 and 38.4 cm.
When earth flow 4 mobilized, the total precipitation
after February 1 was between 22.1 and 24.1 cm, and the
total annual precipitation between 43.2 and 45.2 cm.
Once the earth flows mobilized, their velocities corre-
lated with the amount of daily precipitation; that is,
they moved rapidly during times when precipitation was
high and relatively slowly during times when precipita—
tion was low (fig. 27).

SUBSURFACE DISPLACEMENTS

In earth flow 1, subsurface displacements were
measured by placing a stack of wooden disks in an au-
gered borehole (borehole 10, pl. 3) and, later, excavat-
ing and measuring the displacement of each disk. On
April 9, 1975, the disks, 3.33 cm in diameter and 1.27 cm

30

EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

YEAR 1975 (MONTH—DAY)

 

  
    
 

 

 

 

 

 

 

 

 

 

  
    

 

 

 

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FIGURE 27.—Cumulative displacements of survey stakes on earth flows in Davilla Hill earth
flow complex, January 17, 1975 to June 19, 1975. For stakes in lines 2A and 3A, previous
displacements on surfaces of earth flows 1 and 3, respectively, were averaged to provide
starting points for plots. See plate 3 for locations of stake lines.

 

 

THE DAVILLA HILL EARTH—FLOW COMPLEX 27

flows, each of which came to rest progressively farther
downslope, as illustrated in figure 10. At the downslope
boundary of the earth-flow complex, a completely pre-
served earth-flow deposit—the fourth in this sequence
(A10, p1. 3)—cuts off the third lateral-ridge segment.
The downslope boundary of the earth-flow complex con-
sists of two distinct earth-flow toes, 0.6 to 1.0 m high,
with steep frontal slopes.

Earth-flow movements have been taking place in-
termittently in the Davilla Hill earth-flow complex for
more than a decade (Russell Ferguson, oral commun.,
1975); at least 34 individual earth flows have occurred
(pl. 3), transporting and reworking the soil in the reen-
trant.

We have interpreted the relative ages of the earth-
flow deposits from crosscutting or overlapping relations
and by matching lateral-ridge segments across the com-
plex. The deposits are grouped into three sequences
identified by different capital letters because the ab-
sence of crosscutting or overlapping relations in some
places within the complex precluded the grouping of all
deposits into a single chronologic sequence. Deposits
within each of the three sequences have been dated rela-
tive to each other. Each number or combination of
number and lower-case letter designates a single earth-
flow deposit; numbers increase from the oldest deposit
to the youngest. Lower-case letters refer to earth-flow
deposits that were bracketed in age relative to some de-
posits but which could not be dated relative to other de-
posits with the same number. For example, deposits
C2a, 02b, and 02c are younger than deposit Cl and
older than deposit C3, whereas the ages of deposits 02a,
C2b, and 02c relative to each other are unknown.

MATERIAL

Material in the Davilla Hill earth-flow complex is a
slightly overconsolidated clayey silt of low sensitivity
(table 3), containing varying amounts of organic matter
and minor amounts (less than 5 weight percent) of peb-
ble-size bedrock fragments. The color, consistency, and
organic matter content vary throughout the complex;
however, the grain-size distribution and plasticity of the
material are relatively uniform (Keefer, 1976, 1977a).
The material contains 44 percent clay- (less than 2 pm),
51 percent silt-, and 5 percent sand-size and larger
grains. Atterberg plastic and liquid limits are 30 and 56
percent, respectively. In the Unified Soil Classification
System, the material is a silt with high plasticity (MH).
Predominant colors are gray, yellowish brown, reddish
brown, and black.

The material is pervasively fissured. Open cracks
caused by desiccation and differential movement break
the upper 1 m of soil into blocks smaller than 1 m on a

 

side. Finer, hairline cracks further divide the soil into
lumps a few millimeters on a side.

During most of the study period, the material was
desiccated. Some desiccated material was so stiff that it
could only be penetrated by a sharp blow from a pick;
other desiccated material crumbled readily into pebble-
size lumps along closely spaced fissures. Lumps of desic-
cated material, however, disintegrated within a few
minutes when exposed to water; after even a small
amount of precipitation, a layer a few millimeters thick
on the surface of the earth-flow complex became soft
and sticky, and the material slaked when immersed in
water in the laboratory. In active earth flows, the mate-
rial became so soft throughout that it would not support
a person’s weight.

Peak and remolded undrained strengths of the ma-
terial in active earth flows, measured with. a field vane,
were 11 and 4.9 kN/mz, respectively; the average sen-
sitivity, therefore, is 2.2. In an undrained constant-vol-
ume simple-shear test on this material in the laboratory
(Keefer, 1977a), the sensitivity was determined to be
1.3, in close agreement with the low value determined in
the field. The material has an overconsolidation ratio
(OCR) of 3.1 (Keefer, 1977a); this overconsolidation is
probably due to desiccation.

Clay mineralogy was determined on seven samples
from the site, using the X-ray diffraction techniques de-
scribed by Hein and others (1975). In the six samples of
earth-flow material, the clay contains an average of 49.3
percent illite, 26.2 percent chlorite, 20.1 percent kaoli-
nite, and 4.4 percent montmorillonite (table 4). Clay in
a bedrock sample contains 41.5 percent illite, 33.8 per-
cent chlorite, 21.5 percent kaolinite, and 3.2 percent
montmorillonite (table 4).

SUBSURFACE GEOMETRY

Data from hand-augered borings were used to con-
struct a longitudinal cross section through the reentrant
containing the earth-flow complex (pl. 3). On the steep
slopes at the head of the reentrant, the soil is silty sand;
elsewhere in the reentrant the soil is a clayey silt identi-
cal in appearance to the material in mapped earth-flow
deposits. Soil thickness increases from a few centimet-
ers on the steep slopes above the earth-flow complex to
more than 4 m beneath parts of the complex itself; three
borings drilled into bedrock beneath the complex en—
countered the bedrock surface at depths ranging from
1.7 to 4.2 m (pl. 3). Depths to basal shear surfaces of the
active earth flows, however, were all less than 1.0 m;
these basal shear surfaces were all several meters above
the soil-bedrock contact (pl. 3).

Several auger borings not located on mapped earth-
flow deposits revealed layers containing abundant or-

28 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

TABLE 4.—Clay-mimral content of mtem'als from the Dam'lla Hill
earth-flow complex

 

Location (pl. 3)

Clay-mineral content (percent)

 

 

Sample
Boring Depth (m) Kaolinite Chlorite Illite Montmorillonite

Bedrock ------ XIII 1.2 21.5 33.8 41.5 3.2

Earth—flow VII .2 23.3 36.4 40.3 0
material 1.

Earth-flow IV .3 23.5 27.6 45.5 3.4
material 2.

Earth—flow IV .8 20.3 26.9 51.2 1.6
material 3.

Earth-flow Pit 1 .4 17.3 20.7 52.3 9.7
material 4.

Earth-flow I .3 20.8 20.4 55.3 3.5
material 5.

Earth—flow I 2.4 15.7 25.2 51.4 7.9
material 6.

Average for —— 20.1 26.2 49.3 4.4
earth-flow
materials.

 

ganic material; these layers are interpreted as old top-
soil horizons buried by earth flows. The presence of
these buried topsoil layers, and the similarity between
material outside and inside the mapped earth-flow de-
posits, suggest that most of the material outside the
mapped deposits has also been deposited by earth flows.

STUDIES OF ACTIVE EARTH FLOWS
AND GROUND-WATER CONDITIONS, 1974-78

SURFACE DISPLACEMENTS

Beginning in October 1974, lines of survey stakes
were placed across the Davilla Hill earth-flow complex
to monitor surface displacements (pl. 3). End points of
the lines were placed outside the boundaries of the com-
plex on ground that appeared to be stable. To measure
displacements, a surveyor’s tape was anchored between
the two end stakes in each line, and the position of each
stake in the horizontal plane relative to the tape was
measured with a plumb bob and folding ruler. On the
basis of repeated measurements made before earth-flow
movements began, measurement errors were judged to
be less than 3 cm, except where the distance from the
stakes to the surveyor’s tape exceeded 1 m, where er—

 

rors may have been as large as 15 cm.

Surface displacements on earth flows within the
complex were first recorded on February 21, 1975; on
that date three earth flows (1, 2, 3, pl. 3) were active.
Movement began between February 4 and 21; a survey
on February 4 showed that no earth flows were active.
A fourth earth flow (4, pl. 3) became active between
March 8 and 15.

Neither the locations, shapes, nor boundaries of the
active earth flows could have been predicted in advance,
even from a careful examination of the complex. Only
one earth flow (2, pl. 3) was bounded entirely by exist-
ing scarps and lateral shear surfaces; the other earth
flows were bounded by lateral shear surfaces that cut
across the boundaries of older earth-flow deposits.

On February 21, we noted that material in the ac-
tive earth flows had softened markedly. This material
remained soft as long as the earth flows were moving,
whereas material making up the rest of the earth-flow
complex remained stiff. The boundaries of active earth
flows could, in fact, be mapped in a general way simply
by testing the firmness of the surface. In spite of this
softening, uncased auger holes in the active earth flows
remained open and relatively undeformed while being
transported several centimeters downslope.

The earth flows were active during a period of high
precipitation, when the earth-flow complex was much

THE DAVILLA HILL EARTH—FLOW COMPLEX 31

thick, were placed in the borehole such that each was
free to move independently. The disks were excavated
on July 25, 1975, after earth flow 1 had stopped moving;
figure 29A plots the measured displacements. A slicken-
sided shear surface marked the base of the earth flow

(fig. 293). In all, 94 percent of the displacement observ-
able at the ground surface occurred within 1.27 cm (the
thickness of one disk) of the basal shear surface; little in-
ternal deformation had taken place.

Subsurface displacements were measured less

 

l | I

Lateral shear surface

a.

EXPLANATION

Location of stakes on date measured

W

  

I l l I

Lateral shear su rface

 

 

 

 

    
 
  
  
   
 
 

 

 

 

 

 

 

 

0 3/8/75 E! 4/9/75
4 — A 3/15/75 0 4/14/75 —
D 3/21/75 A 4/23/75
0 3/31/75 I 6/19/75
5 4/2/75 0 Stakes on stable ground. all dates
A
g 6 I l l l | I I
E -2 o 2 4 6 a 10 12 14
LU
E 1
E — I I I I I | I I I I
uI
E
" — O o _
5 ° EXPLANATION O
ff Location of stakes on date measured
Z’ 1 _ . 10/3/74 through 2/4/75 9
3: 2/21/75 Lateral shear surface
9 I End of lateral shear
S 2 _ 0 2/25/75 surface
2 D 3/8/75 Stake lost after
0 t
E - 3/15/75 :1: 371” 5mm"
’2 3 — A 3/21/75 —
E 3 3/31/75 Boundary of toe
2 4 _ 0 4/2/75 _
—1
z v 4/14/75 Ezag/ezlost after measurement
0 5 _ 0 4/23/75 Note: End stakes are not
0 Stakes on stable ground, all dates plotted. The” coordin—
69 ates are (-3.12, 0) and
B . (11,55, OI
l l l l l | | l l l
-1 0 1 2 3 4 5 6 7 B 9 10
‘0 5 I I l I I I I I I I
0 — W *0 —
Lateral shear surface Lateral shear surface
0.5 — _.
EXPLANATION
1-0 — Location of stakes on date measured -
Note: End stakes are not plotted. Their 0 3/21/75
cordinates are (—4.61, 0.91) and D 3/31/75 through 4/23/75
1-5 — (10,05,031) a
2.0 l I I I I I I I l I I
—1.0 —0.5 0 0.5 1.0 1.5 2.0 2,5 3.0 3.5 4.0 4.5 5.0

POSITION ALONG STAKE LINE, IN METERS FROM REFERENCE POINT

FIGURE 28.—Displacement of survey stakes on earth flows of the Davilla Hill earth-flow complex. See plate 3 for

locations of stake lines. A, Line 2A of earth flow 1. B, Line 3 of earth flow 3. C, Line 3A of earth-flow 3.

32 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

 

 

 

 

 

 

l I I I |
Ground surface
0 - _
I I
I I
I
II
0.1 l— II —
I
l
I"
0.2 — I _.
I
II
to I I
s :
I: 0.3 — . _
2 s
2
—. 'I
I l
I; 0.4 _ < _
Lu
2:
I
Basal shear surface 1 '
0 5 T ' —
0.6 m- —
0 7 L _
0.8 l I I l I
0 0.1 0.2 0.3 0.4 0.5
TOTAL DISPLACEMENT 0F DISK FROM
A ORIGINAL POSITION, IN METERS

 

FIGURE 29.—Subsurface displacements measured by placing a stack
of wooden disks in an augered borehole (borehole 10, pl. 3) on
earth flow 1 of Davilla Hill complex. Disks emplaced on April 9,
1975; disks excavated and displacements measured on July 25,
1975. A, Subsurface displacement. Annular space created by dif-
ference between diameters of disks and borehole limited preci-
sion to t 1 cm. B, Slickensided basal shear surface is visible to
left of stack of 3.33-cm-diameter disks. Below basal shear sur-
face, annular space around disks remained open during earth-
flow movement. View obliquely downward.

 

quantitatively in earth flow 3 by driving five stakes to
different depths (15, 30, 46, 71, and 91 cm) in the zone
of accumulation; spacings between stakes ranged from
17 to 56 cm. The stakes were emplaced on February 25,
1975. Between February 25 and March 15, 1975, the
shortest stake toppled owing to softness of the surficial
material, the stakes driven to 30 and 46 cm tilted down-
slope slightly, and the two longest stakes tilted down-
slope until they were buried by earth-flow material and
were nearly horizontal. We concluded that the two
longest stakes had penetrated below the basal shear
surface and that their bases were being held in place
while the stakes were rotated by earth-flow movement.
Thus, the depth of the basal shear surface under the
stake array was between 46 and 71 cm.

WATER LEVELS

During the interval when earth flows were active,
water levels in the complex were measured in uncased
boreholes (fig. 30). With one exception”, water levels
measured on active earth flows were Within 23 cm of the
ground surface. Except for borehole 1, water levels in
inactive parts of the complex were lower than those in
active earth flows. A striking difference between the
water levels in boreholes 4 and 5 was noted on March 8,
1975. On that date, the water level in borehole 5, on an
active earth flow, was at the ground surface, whereas
that in borehole 4, 5 m away on an inactive part of the
complex, was at a depth of 95 cm. The two boreholes
had both been drilled on the same date (February 26,
1975) and to approximately the same depth.

Water-level measurements in open boreholes in fine
soils give only approximate indications of the true
ground-water level (Terzaghi and Peck, 1967; Hanna,
1973). The water-level measurements at Davilla Hill,
however, conform to the more general observations in
that the active earth flows appeared wet throughout and
that water was standing at or near the surface in cracks
on the active earth flows but not on inactive parts of the
earth-flow complex.

WATER CONTENTS

The softening of the material that occurred when
earth flows mobilized was accompanied by an increase in
water content. From December 31, 1974, to April 9,
1975, water-content samples were taken at depths rang-
ing from 0.05 to 2.0 m from 10 auger holes on the earth-
flow complex (boreholes 1-10, pl. 3). Water contents
were measured using standard laboratory techniques

1’I'his exception—borehole 3 on February 26 (fig. 30)—was probably due to lack of suffi-
cient time for equilibration of the water level; the measurement was made 4 days after the
hole wu drilled.

.H

THE DAVILLA HILL EARTH—FLOW COMPLEX 33

(Lambe, 1951); figure 31 shows the water-content pro-
files obtained. Excluding samples with high contents of
organic matter, the average water content for soil in
moving earth flows was 38.4 percent“, the average for
soil in inactive earth-flow deposits was 31.4 percent.
Earth-flow mobilization, therefore, was accompanied by
an average water-content increase of 7.0 percent.

Lateral shear surfaces formed sharp boundaries be-
tween the soft, wet material of the active earth flows
and the stiffer, drier material in inactive earth-flow de-
posits. On February 26, 1975, for example, water-con-
tent samples were taken a few millimeters apart on
either side of a lateral shear surface of earth flow 2 at a
depth of a few centimeters. The soft sample in the active
earth flow had a water content of 38.3 percent; the stif-
fer sample from the immobile material had a water con-
tent of 33.7 percent.

The existence of similar sharp boundaries beneath
earth flows was suggested by distinct increases in resis-
tance to augering felt when auger borings penetrated
basal shear surfaces. Because the augering disturbed
the structure of the earth-flow material, it prevented a
precise determination of water-content distribution near
the basal shear surfaces of active earth flows. However,
measurements in a pit (1, pl. 3) excavated through earth
flow 1 after movement had stopped showed that soil
within a few millimeters of the basal shear surface had
an elevated water content (fig. 32). This result suggests
that thin zones surrounding the basal shear surfaces had
elevated water contents while the earth flows were ac-
tive. Similar basal zones of elevated water content were
described by Hutchinson (1970).

UNIT WEIGHTS, VOID RATIOS,
AND SATURATION

Table 5 lists the unit weights, void ratios, and sat-
urations of one sample from earth flow 2 and of two sam-
ples from earth flow 3. Samples were obtained by push-
ing thin-walled aluminum cylinders, 9.8 cm in diameter
and 10.5 cm high, into the soil and trimming the ends of
the samples flush with the ends of the cylinders. The
samples were placed in airtight containers, and total
weights were measured in the laboratory. Water con-
tents were calculated from measurements made on ma-
terial trimmed off the samples in the field and placed in
separate airtight containers. 15

“Water content wn =100(weight of water)/(weight of dried soil).

lEAlthough the samples from earth flow 3 were obtained several days afler movement
had stopped, water contents of the samples were in the same range as those of material in
active earth flows. This result indicated that only minor desiccation of the material had taken
place after cessation of movement.

 

 

r | l I H l l I II I l I
5
0 _ Borehole 1 +
2 3 9

—o—

0.25 —— I/l_

EXPLANATION

I

0'50 — Borehole on active earth flow 8 _

O

Borehole not on active

0.75 —- earth flow —

1.00 — __

1.25 L ._

11-Dry 0
Total depth =1.35 m

to ‘— I‘m
- N m\
\ \ \

mm M

DATE MEASURED (MONTH/DAYl
YEAR 1975

WATER LEVEL BELOW GROUND SURFACE, lN METERS

 

 

 

1.50 I
m
N
3

1/1 —
m7
2/4 —
4/9 —
4/14

 

 

Date Borehole

Borehole drilled depth
(ml

1 12/31/74 0.75
2 1/17/75 1.60
3 2/21/75 1.30
4 2/26/75 1.42
5 2/26/75 1.24
8 3/31/75 1.12
9 4/09/75 1.12
11 4/14/75 132

 

FIGURE 30.—Water levels in boreholes on Davilla Hill earth-flow com-
plex, January 17, 1975 to April 23, 1975. See plate 3 for locations of
boreholes (numbers 1-5, 8, 9, 11).

LABORATORY SWELL TESTS

Laboratory swell tests were performed to
determine whether the increase in water content that
accompanied earth-flow mobilization was due to an
increase in saturation, in void ratio, or both. To perform
these tests, two naturally desiccated samples of material
from inactive earth-flow deposits were placed in stan-
dard fixed-ring consolidometer cells, 7.6 cm in diameter.
Water was added to the samples, and they were allowed
to swell or consolidate under conditions of low vertical
stress (0.20 kN/mz). These conditions of low vertical
stress simulated the conditions of low overburden stress
in the earth flows. In the consolidation cells, the sam-
ples were prevented from expanding laterally; vertical
expansion or consolidation was measured with dial

34

EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

30 4O

NATURAL WATER CONTENT, IN WEIGHT PERCENT

30 35

30 40

30 40

30 40

 

0.50 -—

0.75

1.00 —-

BOREHOLE 1
12/31/74

1.75 1 '

 

BOREHOLE 2
1/17/75

 

l |

Or
Or

Or
Or

BOREHOLE 3
2/21/75
| l

 

l |

Or

BOREHOLE 4
2/26/75
I l

BOREHOLE 5
2/26/75

 

 

l l

 

30 4O

 

° I

DEPTH, IN METERS

0.25 —

0.50 —

1.00 -

1.50 ’—

1.75 —

 

2.00 1

l

 

 

l l

 

 

 

l l

 

BOREHOLE 6
3/15/75

BOREHOLE 7
3/15/75

BOREHOLE 8
3/31/75

BOREHOLE 9
4/9/75

BOREHOLE 10
4/9/75

FIGURE 31.—Natural water contents in Davilla Hill earth-flow complex, December 31, 1974, to April 9,
1975. Dots indicate determinations in active earth flows; circles, determinations in immobile earth-
flow material; arrow, basal shear surface; Or, sample rich in organic matter. Date below borehole
indicates date drilled and sampled. See plate 3 for locations of boreholes.

THE DAVILLA HILL EARTH—FLOW COMPLEX 35

 

0.00 .

DEPTH, IN METERS

0.50 —

l

 

 

Basal shear surface —>

I

l

   
 

gages. The tests were run for 118 hours (sample 1) and
115 hours (sample 2); however, 90 percent of the
volumetric expansion occurred during the first hour.
Water contents, void ratios, and saturations were calcu-
lated at the beginning and end of each test.

Table 6 lists the results of these swell tests. The av-
‘ erage water-content increase was 7.2 percent, an
amount that corresponded closely to the 7 .O-percent av-
erage increase measured in the field. Thus, addition of
water to immobile earth-flow material was sufficient, by
itself, to bring about the water-content increase that ac-
companied earth-flow mobilization; no remolding or
— other disturbance of the internal structure of the soil
was necessary.

The final void ratios and saturations of the samples
also corresponded closely to values measured in the field
(tables 5, 6). During the tests, the average saturation
increased by 14.5 percent whereas the average volumet-
ric increase was only 1.6 percent. Thus, the increase in

 

 

0.75
0 10

NATURAL WATER CONTENT, IN WEIGHT PERCENT

FIGURE 32.—Natural water contents in pit 1 of Davilla Hill earth-flow
complex, July 25, 1975, showing elevated water content near

basal shear surface.

[Natural water content = 100(weight of water)/(weight of solids).
weight = (total weight)/(total volume).
solids)/(total volume).

20

3° water content was due primarily to an increase in sat-

uration with little accompanying volume change.

 

TABLE 5,—Um't weights, void ratios, and saturation of Davilla Hill

earth-flow material

Total unit
Dry unit weight = (weight of

Void ratio = (volume of pore spaces)/(volume of

 

 

solids). Saturation = 100(volume of water)/(volume of pore spaces)]
Natural Total Dry . .
Earth Sample water unit unit SpelelC Void Saturation
depth . . grav1ty .
flow (m) content weighg weighg of solids ratio (percent)
(percent) (kN/m ) (kN/m )
2 0.18 38.9 17.1 12.3 2.75 1.20 89.3
3 .05 34.1 17.5 13.0 2.75 1.07 87.4
3 .23 37.4 17.7 12.9 2.75 1.10 93.7
Average ------------ 36.8 17.4 12.7 2.75 1.12 90.1

 

TABLE 6.—Laboratory swell tests on Davilla Hill earth-flow material

[Volumetric expansion = 100(final volume - initial volume)/initial volume]

 

 

 

Water content (percent) Saturation (percent) V01d ratio Volumetric
Sample expansion
Initial Final Increase Initial Final Increase Initial Final (percent)
1 32.7 39.5 6.8 76.5 90.0 13.5 1.15 1.18 1.5
2 31.9 39.4 7.5 80.8 96.3 15.5 1.07 1.10 1.6
Average-- 32.3 39.5 7.2 78.7 93.2 14.5 1.11 1.14 1.6

 

36

EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

 

 

 

     

 

 

 

 

 

A
Borehole Depth of Bedrock Depth of
. Date ,
PIezometer , depth basal shear depth pIezometer Remarks
drIIled ,
(m) surface (m) Iml tIp (rnl
I 10/23/75 2.82 0.84 2.24 2.13 - 2,74 ——
II 10/25/75 0.46 — —— 0.32 — 0.37 Destroyed between 4/27/76 and 5/20/76
”I 10/15/75 1.52 — — 1.32—1.37 —
IV 10/22/75 .77 — — .61 — .66 Destroyed between 4/27/76 and 5/20/76
V 10/22/75 2.06 —— — 1.88 - 1.93 Destroyed between 11/16/76 and 12/31/76
Vl 10/20/75 2‘67 —— 2.54 193 - 2‘57 -—
Vll 10/20/75 1.28 .99 — 1.17 — 1.22 —
Vlll 10/16/75 .99 —— — .84 — .89 _—
lX 10/16/75 4.37 — 4.22 3.58—4.19 —
X 11/04/75 2.74 — — 2,01— 2,62 ——
XI 11/04/75 1.75 — 1.73 1.52 — 1,57 Destroyed between 1/6/78 and 1/30/78
XII 11/04/75 2.77 — Surface 2.64 —2.69 —
XIII 10/16/75 1.83 — Surface 1,70 —1,75 —-
IVA 1/14/77 approx .9 — — .79 — .84 —
VA 1/14/77 approx 1.3 — — 1.17 — 1.22 —
VllA 1/14/77 approx 15 — — 1.40 —1.45 —
VIIIA 1/14/77 approx 10 — — .86 — .91 —
B
. XII C | Artesian
t + I Artesia
0.0 —V|IIA—VIIIA —
— VIIIA
0.2 I —
Downslope _ —VI
04 _ Vl - X -
IVA _ IVA’_ NA
0 6 _ IV — |V —V|IA VfiA-VIIA
' —X : —
3; III
Lu IX _ III
E VIII _
2 0‘3 _VIII _VIII 7 ‘
2 _V” — III -VA
;‘ 1 0 VII _V“
- _VII - ‘
E VII VA ' IX
0 12 _VII " -V“ _VA _
— XI — III _VI _IX
14 -
1.6 —
_ v
1.8
S E S 5 °:° ; Q ‘3 3 S 8 8 s 8 z a
52 9 2° 3 34 :7 a E m t. q :I 30‘ :I t. t.
\l \l \l O) 0') 0') 05 U) 0') m C!) a)
U1 L71 0'1 I
| I (A)
_. N 3
s s e
:1 :. :I
\l 01
01

DATE MEASURED

FIGURE 33,—Piezometric measurements on Davilla Hill earth-flow complex. A, Piezometer depths, dates of installation, and dates of de-
struction. 8, Sketch map showing piezometer locations. See plate 3 for detail. C, Water levels in piezometers. D, Dates of piezometer
measurements and precipitation data for observation period. Circles indicate that all piezometers were dry; dots indicate water in one
or more piezometers.

PIEZOMETRIC AND TENSIOME'I‘RIC MEASUREMENTS: 3) had Standard, Casagrande-type porous stone tips, 60

AUGUST 1975-FEBRUARY 1978

cm long and 4 cm in diameter; all other piezometers had
porous ceramic tips, 6 cm long and 2 cm in diameter,

From August 1975 through February 1978, ground- that were specifically designed for use in unsaturated
water levels and pore-water pressures were measured soils. The tensiometers, manufactured by Soilmoisture,

using portable

tensiometers

and open-standpipe Inc., of Santa Barbara, Calif., measured both positive

piezometers (fig. 33). Piezometers 1, VI, IX, and X (pl. and negative pore-water pressures at depths as great as

PRECIPITATION, IN CENTIMETERS

61 cm. The tensiometers and piezometers were installed
according to manufacturers’ instructions and the proce-

THE DAVILLA HILL EARTH—FLOW COMPLEX

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

D
Oct '75 , Nov '75 Dec '75 Jan ’76 Feb ’76 Mar ’76 Apr ’76 June ’76
888 3:2;SSSS::: '3'3'3'3'1’1’2’9399 ‘3‘: 5e 3
ms :sssazsagzs 2%;8339353 ‘3; s3 9
9‘3 1‘ 93 03 b :1 :1 1‘ ‘3 Q 05 :1 :1 ma) \7 :4 :1 O) G) <1 {I m :1 :1 :1 \I
Ox: \1 \1 \I Q m m \l \A \I m m c: o; m 0: a: ch m ca 02
mu! m m m 01 O O m U1 0" O O O OO O O O O O . O O . . . 0
oo o o o o o o 0
Sept ’76 Oct ’76 Nov '76 Dec '76 Jan ’77 Feb ’77 Mar ’77
I “IT IT I | IL
_. _. _\. w
9 7' —- 2.
.A J}
3 e a 3‘.
0: a \‘ \I
O O O 0
Sept '77 Oct '77 Nov '77 Dec '77 Jan ’78 Feb ’78
1 I [l I III lLHlil “I'llT T.
- e N N
m w _; AS
\I 9 P 9’
co \1 {I \l
00 on 00
° 0 o 0
DATE MEASURED

dures described by Keefer (1977a).

No earth flows were active at Davilla Hill during

the period July 1975 through December 1977; the winter

of 1975-76 was the driest ever recorded in the San Fran-

37

cisco Bay area (San Francisco Chronicle, 1976, p. 1), and
the subsequent winter was almost as dry. 16 On most

16Total annual precipitation at Davilla Hill was 29.45 cm in 1975-76 and 31.59 cm in 1976-

77. These values are 49 and 53 percent, respectively, of the 10-year avenge annual precipita—
tion.

38 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

survey dates during this period, all the open—standpipe
piezometers were dry (fig. 33), an observation indicating
that no ground water was present in or adjacent to the
earth-flow material.

The shallow piezometers did contain water during
four brief periods after moderate-size storms (fig. 33);
water levels in these piezometers ranged from 0.57 to
1.66 m below the ground surface (fig. 33). During these
four periods, several other piezometers, including all
those with tips deeper than 1.8 m, remained dry. The
absence of water in the deep piezometers at times when
water was recorded in some of the shallow piezometers
suggests that localized bodies of perched ground water
were formed in parts of the complex after moderate
rainfall. Rapid response of the piezometers after storms
(apparently within 8 hours in the case of the April 8,
1976, measurement) indicated that rainfall infiltrates
rapidly into shallow layers of the earth-flow complex.

Two earth flows became active near the head of the
reentrant between January 6 and January 30, 1978 (pl.
3); activity may have continued into February, although
this observation could not be ascertained because no dis-
placement measurements were made. Piezometric
measurements on January 30, February 14, and Feb-
ruary 23 showed free water in all the piezometers within
the earth-flow complex; water levels ranged from 1.28 m
below the ground surface to slightly artesian (fig. 33).
Water levels in the two active earth flows were not
measured because earth-flow movement destroyed the
one appropriately placed piezometer (XI, pl. 3). The
piezometric measurements, however, confirm observa-
tions in 1975 that ground-water levels in the complex
were near the surface when earth flows were active.
Earth-flow activity in 1978 also occurred during a period
of high precipitation; precipitation in Eden Canyon to-
taled 10.4 cm in December 1977, 25.8 cm in January
1980, and 12.6 cm in February 1980. Total annual preci-
pation when the earth flows began moving was between
26.4 and 46.2 cm.

All but three of the 68 tensiometric measurements
made on 21 different dates from August 1975 through
February 1978 showed that pore-water pressures 61 cm
below the ground surface were negative. This result
conformed to observations and piezometric measure-
ments showing that ground-water levels were generally
deeper than 61 cm during times when earth flows were
inactive (fig. 33). Pore-water pressures were generally
less negative after rainstorms than at other times.
Slightly positive pore-water pressures (at least 4.2 kPa)
were recorded on November 18, 1975, and on February
23, 1978. Owing to instrument failure, no tensiometric
measurements were made on January 30, 1978, or Feb-
ruary 14, 1978, two of the three dates on which the
water levels in piezometers were highest.

 

ANALYSIS OF MOBILIZATION

The earth flows at Davilla Hill moved primarily by
shear along slickensided boundary shear surfaces. The
predominance of boundary shear at Davilla Hill, as well
as the work of previous investigators (Hutchinson, 1970;
Hutchinson and Bhandari, 1971; Chandler, 1972; Hutch-
inson and others, 1974; Dunkerley, 1976; Fussganger
and J adron, 1977; Prior, 1977), indicates that a method
of slope-stability analysis can be applied to the earth-
flow mobilization at Davilla Hill. Slope-stability analysis
combines the principles of equilibrium mechanics, the
Coulomb failure criterion for soils, and Terzaghi’s
theory of effective stress to analyze the potential for in-
itiation of movement along a discrete shear surface in a
soil slope.

Figure 34 illustrates this method of analysis for the
simplified case of a block on a uniformly inclined plane.

According to the principles of equilibrium
mechanics, the block will begin to move when the driv-
ing force L exceeds the resisting force T by an infinitesi-
mal amount. L is the component of the block’s weight
acting in the downslope direction; that is,

L = 'ytVsi'nO,

where y, = the unit weight of material in the block,
V = the total volume of the block,
and O = the angle of slope inclination.

In the case of a soil body, T is the product of the
area of the base of the block and the shear resistance of
the soil, which is defined by the Coulomb failure criteri-
on as modified by Terzaghi’s theory of effective stress.
Thus,

T = IAT'dA

where 'r = the shear resistance of the soil
and A = the area of the base of the block.
The shear resistance is defined by

T = (on—u) tand_> + 6, (1)

where on = the normal stress at a point on the base of
the block,
to = the pore—water pressure at a point on the
base of the block,
(I = the effective angle of internal friction,
and 6 = the effective cohesion.
The quantity (on—u) is the effective stress, as defined
by Terzaghi (1924). The result of a slope-stability analy-
sis is expressed as a factor of safety F, where F = T/L.
A condition F = 1 corresponds to the initiation of move-
ment; if F> 1, no movement can occur.

 

 

THE DAVILLA HILL EARTH—FLOW COMPLEX 39

FIGURE 34.—Stability of block on inclined plane. T, resisting force;
L, driving force; 0, angle of slope inclination. See text for ex-
planation of stability analysis.

Mobilization of earth flows at Davilla Hill was
analyzed using a modified form of the slope-stability
analysis developed by Morgenstern and Price (1965,
1967). This analysis treats the irregular geometries and
nonuniform conditions in earth flows by dividing an
earth flow into a finite number of slices with vertical
sides, determining the differential equations governing
the force and moment equilibrium of each slice, and
numerically integrating the results along the basal shear
surface. The modified Morgenstern-Price method, which
was programmed on a PDP8 minicomputer, was de-
scribed in more detail by Keefer (1977a).17

MEASUREMENT OF SHEAR-STRENGTH PARAMETERS,
USING DIRECT-SHEAR TESTS

Soil shear-strength parameters d) and 6, required
for the analysis, were measured using laboratory direct-
shear tests. The direct-shear test simulates field condi-
tions of boundary shear on earth flows by forcing shear
displacement on a sample to occur in a thin well-defined
zone. A constant vertical load is applied to the sample
during shear, and shear resistance of the soil in the hori-
zontal plane is measured. By performing several tests
under different vertical loads, a Mohr—Coulomb failure
envelope and the shear strength parameters (15 and E can
be determined. By cycling the direction of shear, the di-
rect shear test can also simulate the effects of such large
shearing displacements as occur along natural shear sur-
faces (Skempton and Hutchinson, 1969). A general dis-
cussion of the direct-shear testing method was given by
Lambe (1951).

"A preliminary analysis, described by Keefer (19773), showed that the lateral shear sur-
faces of the earth flows contribute negligibly to the total shear resistance. Therefore, the con-
tribution of the lateral shear surfaces was ignored, and a two—dimensional analysis was used.

 

Two samples from the basal shear surface of earth
flow 1 were tested on a Wykeham—Farrance direct-shear
machine, which had a cylindrical shear box 6.35 cm in di-
ameter. The samples were hand excavated from a pit (2,
pl. 3) and were trimmed in the laboratory to fit the di-
rect-shear apparatus. The basal shear surface was
placed in the plane separating the two halves of the
shear box, and the slickensides were alined parallel to
the direction of shear displacement.

Samples were obtained from a depth of 69 cm,
where the normal stress caused by the weight of mate-
rial above the basal shear surface was on = 11.7 kN/mz.
Assuming that maximum pore-water pressures were as-
sociated with a ground-water level at the ground sur-
face, the minimum effective on site normal stress was on
= 5.1 kN/mz. Each sample was sheared under four dif-
ferent normal (vertical) stresses ranging‘from 5.3 to
19.1 kN/mz. Thus, values of (f) and 6 were measured over
approximately the same range of effective normal stres—
ses as those encountered on site as the ground-water
table rose from the basal shear surface to the ground
surface. Samples were sheared at a displacement rate of
9.75x 103 mm/min, a rate assumed to be slow enough
to allow for complete dissipation of shear-induced pore-
water pressures. “3

Samples were consolidated under each normal load
for 12 to 24 hours before being sheared. Under the first
increment of normal load, each sample was sheared
through a distance of 3.5 mm for each of 12 cycles (6 for-
ward and 6 reverse); under subsequent loads, the sam-
ples were sheared for 8 cycles. Decreases in strength
were observed on successive cycles under many normal
loads, but the strength loss was generally small after
the first two cycles under each load.

Mohr—Coulomb failure envelopes were determined
from straight-line fits of shear resistances measured
during the final forward and final reverse cycles under
each vertical load for each sample (fig. 35). Results from
tests on the two samples agree closely with each other:
for sample 1, d) = 15.8° and E = 0.50 kN/m2; and for
sample 2, 43 = 15.2° and 6 = 0.55 kN/mz. The average
effective shear-strength parameters, therefore, were J)
= 15.5° and E = 0.525 kN/mz.

During shear, the samples generally consolidated
slightly, although on many cycles this general trend was
interrupted by one or more periods of dilatation; dilata-
tion was greatest under the lowest vertical loads. The
shear surfaces in the samples were not perfectly planar,
and so we concluded that this intermittent dilatation
was due to riding up of the top half of the sample over
asperities.

”A later test, in fact, showed that increasing the displacement rate to as high as 1.2 mm/
min causes only minor changes in shear resistance (see fig. 40).

40 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

 

7 I I I I
EXPLANATION
5 — 0

Final forward cycle Sample 1

5’ 0

Final reverse cycle

I I I I T

  
   
    

¢=15.8° -

 

 

Sample 2

SHEAR RESISTANCE, IN KILONEWTONS PER SQUARE METER

 

5:0.55 kN/mzl l

l |

    

¢=15.2°

 

I I I I L

 

0 2 4 6 8

10 12 14 16 18 20

NORMAL STRESS, IN KILONEWTONS PER SQUARE METER

FIGURE 35.—Mohr-Coulomb failure envelopes for direct-shear tests on basal-shear-smface material in earth flow 1 of Davilla Hill earth—
flow complex. (1), effective angle of internal friction; c, effective cohesion.

For most soils containing a large portion of clay, (I)
and 6 decrease with increasing shear displacements. At
very large displacements, values of (l and E approach
lower bounds known as residual values, (I), and Er
(Skempton and Hutchinson, 1969). The direct-shear
tests were performed on samples containing parts of the
basal shear surface along which several meters of move-
ment had already occurred; thus, the values of J) and E
measured in the tests were probably approximately
equal to the residual values for this earth-flow material.

SLOPE-STABILITY ANALYSIS OF EARTH FLOWS

Slope-stability analysis was applied to the three
earth flows shown by direct measurement of displace-
ments to be active (1, 2, 3, pl. 3). Required_for the anal-
ysis are: the shear-strength parameters 4) and E, the
configurations of the ground surface and basal shear
surface of each earth flow (pl. 3), the total and dry unit
weights of the soil (table 5), and the pore-water pres-
sure at specified points on the basal shear surfaces.

 

Pore-water pressures were known only rather im-
precisely. However, field observations and measure-
ments indicated that the earth flows had mobilized when
ground-water levels within them rose to levels inter-
mediate between the basal shear surfaces and the
ground surface (figs. 30, 33). A rise in the ground-water
level in an earth flow would cause the shear resistance
in the soil, as defined by equation 1, to decrease owing
to increases in pore-water pressure. We hypothesized
that increased pore-water pressure is the mechanism
that mobilizes the earth flows.

To test our hypothesis, we analyzed earth flows 1,
2, and 3, and treated the pore-water pressure as vari-
able. An initial analysis was performed on each earth
flow for conditions of no pore-water pressure (u=0) on
the basal shear surface. Several subsequent analyses
were then performed for various elevations of the water
table in each earth flow. In all cases, pore-water pres-
sures were computed on the assumptions that the slope
of the water table was uniform and parallel to the aver-
age slope of the ground surface and that permeability

;

 

 

 

A MODEL FOR EARTH FLOW 41

and ground-water flow in the soil were uniform. The fac-
tor of safety F for each earth flow was then plotted as
a function of the average depth of the water table below
the ground surface (fig. 36).

All three earth flows were immobile (F> 1) for con-
ditions of zero pore-water pressure on the basal shear
surfaces; all three earth flows reached the condition of
mobilization (F = 1) when the water table was at some
depth intermediate between the basal shear surface and
the ground surface (fig. 36). Results of these analyses
thus agree well with observational data on water levels
(figs. 30, 33) and indicate that pore-water pressures in-
duced by rising water tables did, indeed, mobilize the
earth flows.

DISCUSSION

Field observations show that the earth flows at
Davilla Hill are mobilized by addition of water to the
soil; our combined field, laboratory, and analytical
studies indicate the mechanism by which this mobiliza-
tion takes place. During seasons when little rain falls at
Davilla Hill, potential earth-flow material is unsatu-
rated; the material is disrupted by a network of cracks
and is subjected to negative pore-pressures. During
periods of high precipitation, water infiltrates into the
earth-flow complex, locally saturates the fissured near-
surface material, and dissipates these negative pore-
pressures. After moderate precipitation, local bodies of
perched ground water form; with additional precipita-
tion, a continuous body of ground water forms. As the
water table associated with this body rises to near the
ground surface, pore-water pressures increase on poten-
tial failure surfaces within the earth-flow material.
When the pore-water pressures locally exceed threshold
values, earth flows are mobilized.

Most movement on the earth flows at Davilla Hill
takes place by boundary shear, and the water-content
changes that accompany mobilization can be accounted
for by addition of water without any internal distur-
bance or remolding. Internal disturbance and remold-
ing, therefore, play only a minor role in the earth-flow
process at Davilla Hill; the material is mobilized by the
addition of water alone. The main effect of this addition
of water is to increase the pore-water pressures.

Earth flows at Davilla Hill recur year after year in
the same materials. This recurrence is due, in part, to
the moderate velocities at which the earth flows move;
because of these moderate velocities, the earth flows
move only a few meters during a single winter. Thus,
many years and many remobilizations are required be-
fore the earth-flow material reaches the gentle slopes
near the present downslope boundary of the complex.

Slope-stability analysis indicates that rising pore-
water pressure is the mechanism mobilizing earth flows

 

out of the older earth-flow materials. The question re-
mains, how were earth flows originally mobilized at
Davilla Hill? Using upper-bound, peak values of shear-
strength parameters measured in a Geonor simple-shear
apparatus (a), = 25.4°, a, = 5.0 kN/mz), Keefer (1977a) re-
ported that earth flows cannot mobilize out of soil at
peak shear resistance on the slopes presently existing at
Davilla Hill, even with a water table at the ground sur-
face. The near-surface soil at Davilla Hill, however, is
pervasively fissured, and this fissuring reduces the
shear resistance of the soil mass below its peak value. If
the soil were sufficiently weakened by these fissures,
then pore-water pressures caused by a water table at
the ground surface could mobilize earth flows out of pre-
viously undisturbed material in the same way that earth
flows were mobilized out of older earth-flow material
during the study period.

Two field observations indicated that fissures sig-
nificantly reduce the shear resistance of the soil mass.
First, the maximum depth of the earth flows that
mobilized in 1975 and 1978 coincides with the depths of
the deepest extension and desiccation fissures. This
coincidence, due partly to the way in which the network
of fissures facilitates the infiltration of water, is also
probably due partly to a reduction in soil-mass shear
strength. Second, boundaries of active earth flows con-
sistently cut across boundaries of older earth-flow de-
posits, rather than utilizing existing lateral shear sur-
faces. This observation suggests that the shear resis-
tance of the fissured soil mass is not significantly higher
than that along the shear surfaces.

Even without weakening of the soil by fissuring,
earth flows could be mobilized at Davilla Hill out of ma-
terial previously undisturbed by mass movement, if
artesian conditions were to raise pore-water pressures
sufficiently above those due to a water table at the
ground surface. Artesian conditions do sometimes exist
in some parts of the complex (fig. 33).

Because the geometries of and ground-water condi-
tions within the earth flows that originally formed at
Davilla Hill are unknown, a rigorous slope—stability anal-
ysis is not possible. However, the weakening of the
near—surface soil by fissures and the presence of local
artesian ground-water levels suggest that earth flows
can be mobilized out of previously undisturbed material
in the same way that earth flows are mobilized out of
older earth-flow deposits.

A MODEL FOR EARTH FLOW

Analysis of the earth flows at Davilla Hill and at
sites in England (Hutchinson, 1970; Hutchinson and
Bhandari, 1971), northern Ireland (Hutchinson and

42

FACTOR OF SAFETY

EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

 

2.00

1.50 '—

1.00 -

I

I

Earth flow 1

I

 
   

F

Pore pressure = 0

 

 

0.40

0.60

0.80

1.00 1.20 1.40

1.60

 

1.00 -

0.80 -

 

Earth flow 2

l

Pore pressure = O

 

 

0.60
0.00

1.10

0.10

0.20

0.30

0.40

0.50

0.60

 

1.00 —

 

Earth flow 3

l

Pore pressure = 0

 

 

0.50
0.00

0.10

0.20

0.30

0.40

0,50

DEPTH OF WATER TABLE, IN METERS

060

FIGURE 36.—Mobilization of earth flows in Davilla Hill earth-flow complex, showing factor of safety F as a function of water-
table depth. Mobilization occurs when F = 1.

1...‘

 

 

A MODEL FOR EARTH FLOW 43

others, 1974), and Czechoslovakia (Fussganger and J ad—
ron, 1977) indicate that earth-flow mobilization, in gen-
eral, can be explained by a slope-stability model of soil
mechanics that incorporates the effects of pore-water
pressure. This model is also useful as a starting point for
analyzing how the velocity of a moving earth flow is gov-
erned; this aspect of earth-flow behavior, however, re-
quires additional analysis. The main purposes of this
section are to discuss mobilization and to propose a
model to explain the behavior of moving earth flows.

EARTH-FLOW MOBILIZATION

The slope-stability model explains earth-flow
mobilization and several other characteristics of earth
flows. In every active earth flow to which this analysis
has been applied, rising pore-water pressure was shown
to be capable of causing mobilization (Hutchinson, 1970;
Hutchinson and Bhandari, 1971; Hutchinson and others,
1974; Fussganger and Jadron, 1977; Keefer, 1977a, b, c;
Keefer and Johnson, 1978). At Davilla Hill and at most
other earth-flow sites, pore-water pressures necessary
for mobilization were generated by infiltration of water.
In a few earth flows, necessary pore-water pressures
were created by rapid loading of the zones of depletion
by material from other earth flows. This rapid loading
created pore—water pressures in excess of those that
would result from a water table at the ground surface.
At sites where these excess pore-water pressures were
generated, earth flows were observed to move on slopes
as gentle as 4° (Hutchinson and Bhandari, 1971; Hutch-
inson and others, 1974), and Hutchinson and Bhandari
calculated that this mechanism could allow earth flows
to move on even gentler slopes.

Two more general observations support the conclu-
sion that rising pore-water pressure is, in general, the
cause of earth-flow mobilization. First, earth flows
throughout the world commonly mobilize during times
of high precipitation or seasonal snowmelt (Rapp, 1960;
Radbruch and Weiler, 1963; Campbell, 1966; Prior and
others, 1968, 1971; Hutchinson, 1970; Hutchinson and
others, 1974; Nilsen and Turner, 1975; Oberste—Lehn,

. 1976; Keefer, 1976, 1977a; Kelsey, 1977, 1978; Swanson

and Swanston, 1977; Harden and others, 1978; Nolan
and others, 1979; J anda and others, 1980). During such
times, infiltration of water into hillside materials causes
pore-water pressures to rise. Second, many earth flows
occur on hillsides underlain by impermeable rock or soil
layers that have dips parallel to the ground surface
(Sharpe and Dosch, 1942; Kachadoorian, 1956; Prior and
others, 1968; Zéruba and Mencl, 1969); such geologic
settings are conducive to the formation of perched water
tables with resulting high pore-water pressures.
Nothing fundamental in the slope-stability model
requires a minimum, threshold slope for earth-flow oc-

 

currence, and, where excess pore-water pressures are
generated, earth flows can move on very gentle slopes
(Hutchinson and Bhandari, 1971; Hutchinson and
others, 1974). However, in most areas Where systematic
studies have been carried out, earth flows originate only
on slopes with inclinations greater than some threshold
value (Radbruch and Weiler, 1963; Dunkerley, 1976;
Turnbull, 1976; Harden and others, 1978; fig. 3). The
minimum slope inclination on which an earth flow can
mobilize, according to the slope-stability model, depends
on the unit weight and shear-strength parameters of the
soil, the maximum pore-water pressure generated, and
the geometry of the mass that is mobilized. Earth flows
exhibit a limited range of shapes, and the unit weights
of soil materials, in general, vary only within a limited
range. Thus, the threshold slope angle in a given area is
primarily determined by a combination of the minimum
soil shear resistance and maximum pore-water pressure
generated at sites in that area.

The slope-stability model also explains what causes
movement on an earth flow to cease. All the parameters
in the model (unit weight, configurations of basal shear
surface and ground surface, shear-strength parameters,
and pore-water pressure) are interrelated; changes in
one parameter affect the values of the other parameters
that are necessary for continued movement. Thus, if all
the other parameters remain constant, an earth flow
will stop moving if the slope inclination decreases. This
interdependency explains why few earth flows move
beyond the bases of the hillsides on which they origi-
nate. The model also explains why many earth flows
stop moving part way down apparently uniform slopes;
if pore-water pressures decrease below the threshold
value, owing to drainage or evaporation, an earth flow
will stop moving even if the slope is uniform.

MODEL OF A MOVING EARTH FLOW

Whereas the slope—stability model explains how
earth flows mobilize and some other aspects of earth-
flow behavior, an explanation of what controls the veloc-
ity of an earth flow once it has been mobilized requires
the consideration of several additional factors. In par-
ticular, a satisfactory model of a moving earth flow must
be consistent with measurements of earth-flow veloci-
ties, observations of the morphology of shear surfaces
and of associated zones of disturbance, and measure-
ments of the distribution of displacements within earth
flows.

VELOCITY MEASUREMENTS
Earth flows exhibit two basically different patterns

of velocity behavior. The more common pattern is one of
slow movement that persists for several days, months,

44 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

TABLE 7.—Earth-flow velocities

 

 

 

Aver ' d
surfgg: Period of Ve1oc1ty (m/ )
Site s1ope measurement Reference
(degrees) (method) Average Maximum
Panama Cana1 Zone
Cucaracha 1 -------------------- 2 weeks 4.2 -- Cross (1924).
1 week(?) 1.2 —-
6 months .76 --
Cucaracha 2 -------------------- 10 days 4.2 -- Do.
Cucaracha 3 -------------------- 20 minutes -- 1.6x103 Do.
Ames, Colo ------------ 25-30 1 month .3 -- Varnes (1949).
Slumgu11ion, Co1o ----- 7.5 13 years (trees) 4.9)(10'3-1.3><10‘2 —- Crande11 and Varnes (1961).
2 years (survey stakes) 2.1x10'3-1.7x10'2 —-
6 months (camera) 1.2x10'2 -—
Kirkwood, Mont -------- 6-10 1 month .9 -- Had1ey (1964).
Haerenga-o—kuri, —- 26 months 8.8x10'3 8.2x10'3 Campoen (1966).
New Zea1and.
Stoss, Switzer1and---- 7.5 1 year 1.9x10’2 -- Skempton and Hutchinson
(1969).
Mount Chausu, Japan--- 8.5 20 years 6.7x10'2 .24 Do.
Stonebarrow Down, 9.0 2.5 years 4.9)(10‘2 .46 Do.
Eng1and.
Handlova, Czecho- 9.0 5.25 months .40-1.5 6.4 Z5ruba and Menci (1969).
s1ovakia.
Be1tinge, Eng1and ————— 7.0 51 months 4.0)(10'2 2.0 Hutchinson (1970).
Minnis North, northern -— 11 months 3.6x10'2 .13 Prior and others (1968).
Ire1and. 20 -- -- 67 Prior and Stephens (1972).
-- 5 minutes(?) -- 1.2x104 Hutchinson and others (1974).
Davi11a Hi1]
Earth f10w 1 ————— 13 40 days 6.8x10'2 .16 Keefer (1977a).
Earth How 2 ----- 17 47 days 3.2x10'2 7.0x10‘2 Do.
Earth How 3 ————— 20 , 39 days 12x10-2 .39 Do.
Cycle Parkl ----------- 12 26 days 5.9x10'2 .19 Do.
Me1endy Ranch1
Stake ”line B -------------------- 43 days 3.4x10-2 4.9x1o-2 Do.
Stake 1ine c -------------------- 43 days 3.7x1o-2 4.9)(10‘2 Do.
Sochi, B1ack Sea
coast, U.S.S.R.
Earth f1ow P ----- 14 7 years 3.6x10'5 -— Ter-Stepanian and Ter—'
‘ Stepanian (1971).
Earth Flow 0 -------------------- 7 years 4.2x10'5 -- Do.
"Y" Earth flow, —- 85 days .50 2.07 Oberste—Lehn (1976).
LomeEias Muertas (Jan.-Apr. 1974)
area . -- 23 days .33 -—
(Feb. 26—Mar. 20, 1975)
-— 7 days 1.49 2.26
(Mar. 20-27, 1975)
Van Duzen River basin,
Calif.
Ha110ween earth -— 3 years 6.5x10'2 -- Ke1sey (1977, 1978).
flow.

 

A MODEL FOR EARTH FLOW 45

TABLE 7,—Earth-flow velocities—Continued

 

Average

Velocity (m/d)

 

 

, Period of
Site 53;;Sgc measurement Reference
(degrees) (method) Average Maximum
Donaker earth flow -------------- 1 year 1.6x10'3 -- Kelsey (1977, 1978).
Cashla Pooda
earth flow 3
(west tongue) ----------------- 1 year 4.7x10' -- Do.
(east tongue) ----------------- 1 year 4.7x10'3 --
Broken Road
earth flow
(west tongue) ----------------- 1 year 4.7x10'3 —— Do.
(east tongue) ----------------- 1 year 3.3x10'3 --
Chimney Rock -- 1 year 1.5x10‘2 —- Do.
earth flow.
Falling Tree —- 1 year 7.4x10'3 -- Do.
earth flow.
Lookout Creek earth 14 1.4 years 3.8x10“4 -- Swanson and Swanston (1977).
flow, H. J. Andrews (Dec. 12, 1974-
Experimental Forest, May 18, 1976)
Dreg. -- 5 years (1974-79) 2.4x10'4 -- Swanson and others (1980).
Redwood Ersek basin,
Calif. ’
Poison Oak Prairie -- 2.4 years 9.1x10'5—1.2x10'Z -- Harden and others (1979).
earth flow.
Minor Creek earth -- 2.4 years 7.4)(10‘5-1.5x10'Z -- Do.
flow.
Rain Gage earth —- 1.7 years 5.9x10‘5-8.4x10'3 -- Do.
flow.
Devils Creek —— 2.1 years 8.4x10'4-2.3x10‘3 -- Do.
earth flow.
Counts Hill -- 3.1 years 3.5x10-5-4.9x10-3 —- Do.

Prairie earth
flow.

 

1See figure 1 for location.

2Velocities are based on survey measurements several months a
so actual velocities may be greater than those calculated from sur

3Given range of velocities was derived from Harden and others

of survey stakes between each of two successive survey dates.

or years; the less common pattern is a surge of rapid
movement that lasts a few minutes. Surges have been
observed only on few earth flows, and it is not known
whether all earth flows are capable of surging. How-
ever, the available data show that all earth flows, in-
cluding those that surge, move in a slow, persistent
manner most of the time. Measured velocities during
slow, persistent movement range from less than 1 mm/d
to several meters per day (table 7). Velocities vary from
day to day, minute to minute, or second to second, but
continuous acceleration leading to high velocities does
not generally occur.

One of the earliest and clearest accounts of slow,
persistent earth-flow movement was provided by

 

part; these intervals may include some periods of inactivity, and
vey measurements.

(1978, table 8) by determining the average velocity of each line

Blackwelder (1912, p. 489-490), who described move-
ments within the Gros Ventre earth-flow complex in
northwestern Wyoming as follows:

“According to residents of the district, the [earth—flow complex]
first came into action in May, 1908. So far as I am able to learn, no one
actually saw it beg-in; but it is believed by some that the initial move-
ment was fairly rapid if not indeed precipitate. When first observed,
the disturbance was manifested only at the head of the gulch, where
large masses of the slippery Morrison and Sundance (Jurassic) clays
had slumped down along the steeper slopes, overturning trees and
leaving a general wreck. Either quickly or slowly, the impulse from
this upper mass was then communicated to the old landslide debris
farther down the valley, and that in turn began to press forward,
bulge, and crack. The novel thing about this case is that the movement
of at least the lower part was very slow and yet continuous, like that
of a glacier * * *.

46 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

“Whether or not motion continued during the winter of 1908-1909
I have not learned; but in the spring of 1909 the material in the lower
part of the valley slowly pushed forward and its surface bulged into
low irregular domes fretted with open crevasses, many of which were
several feet wide * * *.

“So far as observed, the motion of the slide was not at any time
rapid enough to be actually seen. The evidence of it, however, was
plain enough. It was found impossible to keep the Forest Service tele-
phone line in repair more than a few days, for the poles would slowly
move down hill or be overturned and thus snap the wire. The wagon
road soon became so hopelessly twisted and broken that it was almost
impossible for wagons to follow it without capsizing, and it was no
easy task to cross it even on a saddle horse. Attempts to repair the
damage were almost futile, because in a few days the road would be
rendered again impassable by folds of earth several yards in height or
by gaping crevasses with vertical walls * * *.

“According to members of the United States Forest Service, the
slide did not move as one mass, but rather in sections; the disturbance
began on the east side and manifested itself week by week at new
places. Changes progressed most rapidly in the wet spring months and
declined noticeably toward autumn. The slow but apparently incessant
movement continued through the years 1908, 1909, and 1910, but in
1911 had practically ceased.”

Significant aspects of this description that illustrate
the behavior of earth flows in general are: the slow
movement at varying velocities over a period of several
years, the acceleration during wet periods, and the pro-
gressive mode of failure in which failure began near the
crown of the complex and proceeded downslope.

Blackwelder also noted that an initial surge of rapid
movement may have occurred on the Gros Ventre earth—
flow complex. Such initial surges of rapid movement
have been observed on other earth flows; for example,
Cross (1924, p. 25) stated: “The motion of [an earth
flow] is apt to be rapid at first, but it soon becomes slow
and resembles that of a glacier.” The initial surge of
rapid movement of an earth flow in the San Francisco
Bay region was described in a similar way by Krauskopf
and others (1939, p. 630-631): “The principal movement
* * * was sufficiently rapid and noisy to cause a near-
panic in the surrounding countryside. During several
days the motion subsided * * *.” Study of aerial photo-
graphs flown in 1974 showed that movements were, in
fact, still taking place 35 years after the original descrip-
tion.

Most earth flows move faster during periods of high
precipitation or snowmelt than during drier periods
(Campbell, 1966; Prior and others, 1968, 1971; Hutchin-
son, 1970; Hutchinson and others, 1974; Oberste-Lehn,
1976; Keefer, 1977a; Kelsey, 1977, 1978; Swanson and
Swanston, 1977; Harden and others, 1978; Nolan and
others, 1979; J anda and others, 1980; Swanson and
others, 1980).19 Correlations between precipitation and
velocity are complex, however, and the velocity of an

H‘()ne exception to this generalimtion was observed on the Slumgullion earth flow in the
San J uan Mountains of Colorado, which moved at a nearly constant velocity during an obser-
vation period of13 years (table 7).

 

earth flow also appears to depend on several other fac-
tors.

A few earth flows have been monitored with de-
vices that record movement continuously; velocities of
these earth flows vary according to three distinct pat-
terns (fig. 37). In the first and most common pattern,
earth flows move at constant velocities for several days
at a time (figs. 37A, 373); these periods are commonly
interrupted by shorter periods of acceleration 0r decel-
eration (fig. 37A). The second movement pattern is one
of stick- slip, in which earth flows advance at relatively
low velocities for several hours and then abruptly surge
forward a few millimeters or centimeters at relatively
high velocities (fig. 370). The average velocities of earth
flows exhibiting stick-slip behavior are determined by
the number of small surges within a given period. Both
the first and second patterns result in generally slow
and persistent movement.

The continuous—recording devices also registered a
third and basically different movement pattern—that of
a major surge, in which an earth flow advances several
meters at a velocity of several meters per minute (fig.
37D). At the Minnis North site in northern Ireland,
where such surges were recorded, they are most com-
mon on earth flows being rapidly loaded under condi-
tions generating excess pore-water pressures. The
surges commonly occur during or shortly after
rainstorms (Prior and Stephens, 1971, 1972; Hutchinson
and others, 1974); most surges begin suddenly and are
followed by periods of smooth, gradual deceleration.
One surge was described by an eyewitness as follows
(Hutchinson and others, 1974, p. 371):

“(3) At 4.15 pm there was heavy rainfall while surveying was in prog-
ress, and considerable noise from surface water which ran down the
feeder slides [earth flows with relatively steep surface slopes] and
across the accumulation slide [earth flow with relatively gentle surface
slope] * * * There was no sign of any movement in the accumulation
slide at this time.* * *

(b) At 4.25 pm a sudden cessation of the noise of running water was
observed even though it was still raining heavily * * *. As the noise
ceased, a layer of mud was observed to move rapidly down the feeder
slides, where previously nothing moved except water. This mud over-
rode the two electrical piezometer positions and deep—seated move-
ments of the accumulation slide commenced simultaneously. These
movements were largely confined to a well-defined channel, bounded
by shear surfaces, except where overflow produced miniature levees.
The material accelerated perceptibly on the accumulation slope and
again on the steeper front slope, eventually spilling across the road as
a great lobe of wet mud.

(c) By 4.30 pm all movements of mud had ceased and a lobe 0.5 metres
thick and about 400 square metres in area was completely blocking the
road. The noise of running water was heard again, and the water was
seen to be following the much deepened [earth flow] channel.”

SHEAR SURFACES AND ADJACENT
ZONES OF DISTURBANCE

All earth flows are bounded by slickensided shear
surfaces (figs. 15, 17, 19, 29B); their ubiquity indicates

 

A MODEL FOR EARTH'FLOW 47

\I

 

IIIIII
GI—A

IIIIIIIIIIIIIIITIIII

   
 

 

 

 

CUMULATIVE DISPLACEMENT, IN METERS

 

 

 

 

0 I l I I l | l I I I l I | I I l I I I L I I I I I J
JASONDIJ FMAMJ JASONDIJ FMAMJJA s
1963 1964 1965
22° I I I I I I I
180— B —
140 — —
100— —
60 - _
20’— _.
4 I l I I I
NOV18 19 20 21 22 23 24 25

 

CUMULATIVE DISPLACEMENT, IN CENTIMETERS

 

 

 

 

0
NOV18 19 20 21 22 23 24 25

 

3/10/71 3/31//71 7/28/71 11/4/71 _

Surge

 

CUMULATIVE DISPLACEMENT, IN METERS

 

 

 

HOURS

FIGURE 37.—Patterns of movement of earth flows monitored with con-
tinuous—recording devices. A, Example of movement at relatively
constant velocity with short periods of acceleration and decelera-
tion. Waerenga-o—kuri earth flow, New Zealand (from Campbell,
1966). Reprinted with permission from the National Water and Soil
Conservation Authority, New Zealand. 8, Movement at relatively
constant velocity. Earth flow 1, Minnis North, northern Ireland
(from Prior and Stephens, 1972). Reprinted with permission from
the Geological Society of America and D. B. Prior. C, Slip-stick
movement. Earth flow 2, Minnis North, northern Ireland (from
Prior and Stephens, 1972). Reprinted with permission from the
Geological Society of America and D. B. Prior. D, Major surges.
Earth flow 1, Minnis North, northern Ireland (from Hutchinson and
others, 1974). Reprinted with permission from the Geological Soci-
ety of London and J. N. Hutchinson.

 

that boundary shear is an important component of
earth-flow movement. In overall form, shear surfaces
are planar or gently curved, although they also contain
local asperities (fig. 17). Adjacent to some basal shear
surfaces, zones of disturbed material are also present
(fig. 18); disturbed zones examined by us ranged from
0.2 to 30 cm in thickness (figs. 17, 19). Zones of dis-
turbed material, a few centimeters or tens of centimet-
ers Wide, also adjoin many lateral shear surfaces. The
disturbance is manifested by sets of echelon cracks
trending oblique to the shear surfaces (fig. 143).

DISTRIBUTION OF DISPLACEMENTS

Measurements of surface and subsurface displace-
ments confirm that most movement on earth flows takes
place by boundary shear or by displacement in thin
zones adjacent to the boundary. The predominance of
displacements in boundary zones is particularly striking
in subsurface-displacement profiles. In Davilla Hill
earth flow 1, 94 percent of the total displacement meas-
ured at the ground surface took place on or within 1.3
cm of the basal shear surface (fig. 29A). Similarly, in an
earth flow at Beltinge, England, 89 to 95 percent of the
displacement measured at the ground surface took place
on or within 20 cm of the basal shear surface (fig. 38).
Analogous results were obtained by Prior and Stephens
(1972), Who measured the deformation of flexible plastic
tubes in an earth flow at Minnis North, northern Ire-
land, and concluded that most displacement took place
by slip on the basal shear surface.

Measurements of surface displacements also show
that, whereas some internal deformation occurs, dis-
placements on and adjacent to lateral shear surfaces
predominate (fig. 39). Many surface-displacement pro-
files show little internal deformation (figs. 39A-39C);
other profiles are irregular and show abrupt discon-
tinuities in displacement due to shear on discrete inter-
nal shear surfaces (figs. 39C, 39D); in still other profiles,
more uniform distribution of differential displacements
across the earth flows indicates that some distributed
internal shear or flow is taking place (figs. 39E, 39F).

SURGES

The surges of rapid movement that occur on some
earth flows can be explained by using the above data
and slope-stability analysis. According to slope-stability
analysis, an earth flow is mobilized when pore-water
pressures on the shear surfaces increase to values at
which the resisting shear force is smaller than the driv-
ing force. If pore-water pressures were to rise even
higher, as they commonly do because of additional pre-
cipitation and infiltration, then the resisting shear force

48 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

would decrease still further. According to Newton’s
third law, as long as the driving force exceeds the resist-
ing force, the earth flow will accelerate; this acceleration

Ground surface

 

 

 

 

 

Basal shear surface

0 0.5 1.0 1.5 METERS
| l | I

FIGURE 38.—Subsurface-displacement profile of earth flow at Bel—
tinge, England. From 89 to 95 percent of total displacement took
place on or within 20 cm of basal shear surface. Measurements were
made with borehole inclinometer (from Hutchinson, 1970). 1, initial
position of inclinometer tube, on April 15, 1964; 2, position on April
22, 1964; 3, position on April 28, 1964; 4, position on June 11, 1964.
Lines 1 to 4 are corrected for initial deviations from vertical. Re-
printed with permission from J. N. Hutchinson.

 

would, if continued, lead to movement at high velocity,
which is apparently what happens in a surge.

In a published eyewitness account (Hutchinson and
others, 1974) of a major surge, the beginning of the
surge was preceded by rapid loading in the zone of ac—
cumulation that, in turn, caused an increase in pore—
water pressure. The surge began abruptly, and acceler-
ation led to high velocities within a few minutes. The
surge continued until the earth flow reached the base of
the hill on which it originated and encountered a gentler
slope; this decrease in slope restored the force balance
by decreasing the driving force. Association of other
surges on this same group of earth flows with periods of
high precipitation (Prior and Stephens, 1971, 1972;
Hutchinson and others, 1974), indicated that increased
pore-water pressure, in general, caused the surges.

Slope-stability analysis can be similarly used to ex-
plain the small surges that are part of the stick-slip
movement of some earth flows (fig. 370). These small
surges may occur when pore-water pressure briefly ex—
ceeds the threshold value required for mobilization.

VELOCITY, BOUNDARY SLIP,
AND BOUNDARY ROUGHNESS

Surges occur relatively infrequently and have been

recorded on only a few earth flows. More commonly,
earth flows move at relatively constant velocities for
long periods. Although these periods of constant veloci-
ty are interrupted by periods of acceleration or deceler-
ation, sustained acceleration leading to very high veloci-
ties does not generally occur. Slope-stability analysis,
however, predicts that an earth flow will accelerate as
long as the pore-water pressure exceeds the threshold
value required for mobilization.

To explain these long periods of constant velocity,
using an unmodified slope-stability analysis, it is neces-
sary to assume that pore-water pressures are main-
tained at, but not above, threshold values over long
periods. We infer that such maintenance of pore-water
pressures is unlikely because of the variations in precipi-
tation, infiltration, drainage, and evaporation to which
all active earth flows are subjected. Even if the pore—
water pressure were to remain constant over time in
some cases, there is no apparent physical reason to ex-
pect that the constant pore-water pressure would
exactly equal the threshold value. Hence, slope-stability
analysis must be modified to explain the absence of sus-
tained acceleration. This modified analysis must then
address conditions on and adjacent to the shear sur-
faces, because the largest component of movement oc-
curs there.

4

i
1

 

'
C

V

*a,~

 

l

‘V ‘ . ‘
V .

V o

A MODEL FOR EARTH FLOW 49

One possible mechanism for preventing sustained
acceleration when pore—water pressures rise above
threshold values assumes the existence of a velocity-de-
pendent component of shear resistance. If, for example,
the shear resistance of the material along the shear sur—
faces were given by the equations

[T — (on—u) tan 3) — 6]"=AES; 'r>(on—u)tan $+6 (2a)

and 0 =AE‘S; Tswn— u)tan J> + 6 (2b)
where E's is the deformation rate, A and n are constants,
and T, on, u, (I), and 6 are defined as in equation 1,
then the shear resistance would increase as the veloci-
ty20 increased, and this increased resistance would re-
tard the movement. This rheologic model'is a combina-
tion of the Mohr—Coulomb and the power-law_models,
and equation 2a reduces to equation 1 when ES=O. A
form of equations 2 with n: 1 was suggested as a
rheologic model for soils by several investigators (Ter-
zaghi, 1931; Stroganov, 1961; Ter—Stepanian, 1963; Yen,
1969; Johnson, 1970).

To determine whether shear resistance along earth-
flow shear surfaces does, in fact, increase with increas-
ing velocity, we performed an experiment, using mate-
rial from Davilla Hill earth flow 1. In this experiment,
a sample containing part of the basal shear surface was
placed in the direct-shear apparatus, consolidated, and
then sheared at several velocities ranging from
1.5x 10‘3 mm/min (0.21 cm/d) to 1.2 mm/min (180 cm/
d). The range of velocities used in the experiment in-
cluded the full range of velocities measured on the earth
flow in the field. At the lower velocities, conditions were
probably drained; at the higher velocities, conditions
probably were partially drained or undrained. The sam-
ple was initially sheared for 12 cycles at a velocity of
37x 10‘3 mm/min (5.3 cm/d); the decrease in shear re-
sistance associated with increasing displacement was
thus significantly diminished before velocity of shear
was varied. The sample was consolidated between tests
at different velocities. The vertical load remained con-
stant during the experiment (on = 5.3 kN/mz); total
shear resistance in the horizontal plane was measured
during one forward and one reverse cycle at each veloci-
ty, and the two values of shear resistance obtained were
averaged to plot the results in figure 40.

A linear-regression-line fit to the data in figure 40
yields the relation:

1' = 1.800 + 0.046log10 v,

2"The velocity is obtained by integrating the deformation rate E, over the thickness of
the zone in which deformation occurs.

 

where T = the shear resistance (in kN/mz)
and v = the velocity (in cm/d).

The near-constancy of shear resistance over the
range of velocities tested indicates that material along
the shear surfaces does not exhibit any significant veloc-
ity-dependent component of shear resistance. This re-
sult conforms to results from direct-shear tests on other
fine soils (Kenney, 1967; Ramiah and Purushothamaraj,
1971).

Another possible mechanism for retarding move-
ment is boundary roughness. This mechanism, believed
to be important in retarding basal slip in temperate
glaciers (Kamb and LaChapelle, 1964; Weertman, 1964;
Paterson, 1969), may also operate on earth flows.
Though not verified by experiment, this mechanism con-
forms to the data on earth-flow velocity, displacement
distribution, and shear-surface morphology.

We hypothesize that boundary roughness retards
movement as follows. Asperities on the boundaries of
earth flows force material adjacent to the shear surfaces
to deform around the asperities or to shear through
them. The asperities themselves contribute a compo-
nent of shear resistance that varies over time as the
earth flow deforms and advances over different parts of
the bounding channel. In addition, the properties of the
material thus forced to deform around the asperities
may differ significantly from those of the material along
the shear surfaces. In particular, soil that does not con-
tain discrete shear surfaces commonly exhibits a veloci-
ty—dependent component of shear resistance (Fleming
and Johnson, 1975; Mitchell, 1976). Therefore, deforma-
tions in this material could retard earth-flow movement.

Field observations show that asperities occur on
earth-flow shear surfaces (fig. 17) and that the material
in zones adjacent to shear surfaces has been deformed
(figs. 18, 19). Even parts of basal shear surfaces that ap-
pear planar, such as the basal shear surface of Davilla
Hill earth flow 1, contain small asperities, a few mil-
limeters high. Traces of most lateral surfaces are also ir-
regular, and internal shear adjacent to lateral shear sur—
faces has been directly observed (Hutchinson, 1970).

These observations show that, in detail, earth-flow
boundaries are rough and that material adjacent to the
shear surfaces deforms and shears. In glaciers that slip
on their beds, the velocity of basal slip depends on the
shape, size, and spacing of asperities and on the
rheologic properties of the ice (Paterson, 1969). Similar
factors may also control the velocity of slip in earth
flows, although neither the mechanism by which defor-
mation near asperities takes place nor the rheologic
properties of the earth-flow material are yet well
enough known to analyze the boundary shear and as—
sociated deformations quantitatively.

50 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

  
  
  
   
 
   

 

 

  

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 
     

 

 

 

 

   

 

 

 

 

 

f v
, Stake g ‘0 5 I I I I I I I I I I
line D uI—J "I
Lateral 2 a _
shear z 0 _ ‘V
surface _- GUHV
Lu
2
Stake j 0 5 _ _
line K 3 ' EXPLANATION
<
5 Location of stakes on date ‘
4 _I 1.0 — measured ‘
4
z 7
3 5 0 8/19/74
0) _
1:: I _ __
E 2 o 1.5 v 8/6/75 y
w 2 I [I
E 1 2 V 4/22 76
LL ._ _.
0 l l l I I— 2'0
0 1 2 3 4 E ‘
METERS E 2 5 _ 4
it ' B .9
i
m
5 30 | I I I | I I I I I | I | ’
0 10 20 30 40 50 60 70 80 90 100 110 120 130 140
POSITION ALONG STAKE LINE, IN METERS
w‘ 0 II I T I I I I T II I r I I 7 I I I I
E October 1958 Stakes on stable
m I I I Stakes on stable ground ground
5‘ 5 ‘ June 1959 — (g (g
E ' ‘ w Lateral w 0 _ W ‘
m I shear 5
-’ S t mbe 1959 f e
g .0 a Iw W
_ (I! 0.5 —- —
g I I Internal shear surface/I E
1“ ~ Lateral shear surface
2 E 15 — I\Lateral shear surface — 112.!
o E I | 3 1.0 — —
0: E w
"L z
I. 20 — I — <
5 October 1960 I7)
5 I I I _, 1.5 ~ _
<(
§ 25 — I I I — z
—I C ‘2
f), I: 210 — —
5 30 I | I I I I I I I I 2 /Lateral
0 20 40 60 80 100 120 140 160 180 200 220 O shear
u: f
POSITION ALONG STAKE LINE, IN METERS I; 2.5 — 5‘" ace —
E
. '0-5 I I I I I I I 2
w w 3.0 — —
E 2 Internal shear surface D
4 .1
LU 0 _, A a,
i g 35 I I I I I
:7, . ;- o 5 1o 15 20 25 30 35 '
_I v V v
g 0.5 — EXPLANATION — POSITION ALONG STAKE LINE, IN METERS ‘
C3 gt) Location of stakes
g E on date measured EXPLANATION ' ’
u: 1 O — 1 I, [I A
g 2 A 0 25 73 Location of stakes on date measured
E .2. o 1/15/74 ‘
,_ 1.5 _ 0 3/23/74 _ 11/4/74 through 212 75
E I 7/23/74 3 19 75 i
E D 4/1/76 A 32775
3 2'0 — v 119/76 — . 4/6175 ’
5;, v 4/5/75 5 I 41375
E
25 I I I I I ' I I ‘ 28 76 5 a-
0 10 20 30 40 50 60 70 BO

POSITION ALONG STAKE LINE, IN METERS

FIGURE 39.—Surface-displacement profiles of earth flows as indicated

by displacements of lines of survey stakes. A, Most movement oc-
curs on or adjacent to lateral shear surfaces. Earth flow at Bel—
tinge, England (from Hutchinson, 1970). Dates of measurement:
1, November 14, 1962; 2, March 27, 1963; 3, April 16, 1963; 4,
June 7, 1963. Reprinted with permission from J. N. Hutchinson.
8, Most movement occurs on or adjacent to lateral shear surfaces.
Stake line 3, Devil’s Creek earth flow, Redwood Creek basin,
Calif. (from Harden and others, 1978). C Large component of

movement at boundary and small component of differential move-
ment on a discrete internal shear surface. Earth flow at Slumgull-
ion, Colo. (from Crandell and Varnes, 1961). D, Large compo-
nents due to movement at boundary and to differential movement
on a discrete internal shear surface. Earth flow at Cycle Park,
Calif. (from Keefer, 19773). E, Significant component of distri-
buted internal shear or flow. Stake line 1, Poison Oak Prairie
earth flow, Redwood Creek basin, Calif. (from Harden and
others, 1978). F, Significant component of distributed internal

a, 1"

 

(

:1

b

 

 

A MODEL FOR EARTH FLOW 51

 

 

 

 

N S
I l f
o UDD-oo-oooooflfl —
20 _ _ EXPLANATION
Location of stakes on date measured
0 1/18/74
I 3/3/74
0 3/17/74
(I)
E A 4/13/74
5
5 D Stakes on stable ground,
Z all dates
u?
z
:1
I.“
X
<(
p.
U)
_l
<
Z 40 — _
9
K
O
E
O
0:
LL
I—
2
LL!
2
Lu
0
<(
_l
Q.
‘9
o
60 — -
F
80 1 | l
0 30 60

POSITION ALONG STAKE LINE, IN METERS

shear or flow; boundary slip, however, is largest component of
movement. “Y-slide” earth flow, Lomerias Muertas area, Calif.
(after Oberste-Lehn, 1976). Reprinted with permission from
Deane Oberste-Lehn.

 

INTERNAL DEFORMATION

Surveyed displacement profiles indicate that some
internal defamation also occurs in earth flows (fig. 39).
Some of this deformation is due to shear on discrete in-
ternal shear surfaces; the mechanism controlling this
shear probably resembles that controlling shear at the
boundaries. However, some distributed internal defor-
mation, which may be due to flow within the material,
also occurs.

Displacement profiles of material flowing in a chan-
nel have been calculated by several previous inves-
tigators using various rheologic models for soils (Yano
and Daido, 1965; Paterson, 1969; Johnson, 1970; Cunnin-
gham, 1972; Keefer, 1977a). For channels of uniform
cross section and for uniform material properties, these
calculations show that the material moves faster near
the centers of channels than near the boundaries, a re»
sult in general agreement with velocity profiles meas-
ured on earth flows (fig. 39). The exact shapes of these
theoretical velocity profiles depend on the rheologic
model used, the properties of the material, and the
geometry of the channel.

Assuming that equations 2 apply to material in the
interior of an earth flow, even though they do not apply
to material disrupted by discrete shear surfaces, veloci-
ty profiles due to internal flow in infinitely wide and
semicircular channels were derived by Keefer (1977a).
According to these derivations, a rigid plug with no in-
ternal deformation exists in the center of the channel.
Between this plug and the channel boundary is a zone
where internal deformation takes place; the velocity
profile in this zone is curved. The thickness of this zone
of deformation and its degree of curvature depend on
the shear-strength parameters and unit weight of the
material, the surface slope, the values of constants A
and n in equation 2a, and the pore—water pressure.

The detailed field and laboratory measurements
needed to quantitatively compare theoretical profiles
with profiles surveyed in earth flows have not been
made. However, the shape of the theoretical velocity
profile generally agrees with the shapes of profiles sur-
veyed on earth flows that exhibit significant distributed
internal deformation (figs. 39E, 39F). In addition, the
analysis for an infinitely wide channel, in which pore-
water pressure was treated as a variable, shows that
the maximum velocity within an earth flow increases
with the pore-water pressure. This result also agrees
with field observations.

COMPOSITE MODEL

Earth flows move primarily by boundary shear, and
the velocity appears to be controlled primarily by a

01
N

EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

 

2-2 | IIIIIIII I II

2.0 — 1
O

llllll I ll

llllrr I II

M

 

 

SHEAR RESISTANCE, IN KILONEWTONS
PER SQUARE METER

llllll 1 ll

llllll l I |

wCDar
L"C
C

 

llllll I llllll

 

10 100 1000

DISPLACEMENT RATE, IN CENTIMETERS PER DAY

FIGURE 40.—Shear resistance as a function of velocity in direct-shear tests on basal-shear-surface material in earth flow 1 of Davilla Hill earth-
flow complex. Numbers indicate order in which tests were run. Line indicates linear-regression fit of data.

mechanism involving boundary roughness and complex
deformations in thin zones near the boundary shear sur-
faces. Added to this major boundary—shear component of
displacement are smaller contributions from distributed
internal deformation and from shear on discrete internal
shear surfaces.

CONCLUSIONS

Earth flows differ from other types of mass move-
ment on the bases of morphology, materials, and rate
and kind of movement. Characteristic morphologic fea-
tures are: a tongue or teardrop shape; a rounded, bulg-
ing toe; a sinusoidal longitudinal profile, concave upward
near the head and convex upward near the toe; a pair of
flanking lateral ridges; and a set of discrete boundary
shear surfaces.

Many earth flows occur in large complexes made up
of myriad individual earth-flow deposits. Within these
complexes, earth flows mobilize from material in older
earth-flow deposits and from material in other types of
mass-movement deposits, as well as from material not
previously involved in mass movement. The boundaries
of earth flows that remobilize older earth-flow material
commonly cut across those of older earth-flow deposits.
The common remobilization of earth—flow deposits indi-
cates that no significant irreversible changes take place
in the soil when an earth flow stops moving. Therefore,
material in earth-flow deposits is highly susceptible to
renewed movement.

In many zones of depletion are deposits of slumps
or of other types of mass movement. This relation has
led many previous investigators to suggest that earth
flows are derived from slumps or other types of mass
movement by progressive disruption of internal struc-
ture. Observations during this study, however, indicate

 

that this mechanism is not generally valid. In many
zones of depletion, mass-movement deposits are derived
from failure of the steep scarps left behind by earth-flow
movement; thus, these mass-movement deposits are
younger than the associated earth flows, and the earth
flows are not derived from them. In addition, some
earth flows mobilize out of materials not previously
transported by mass movements, and other earth flows
mobilize directly out of older earth-flow material in the
absence of any other mass-movement process.

Material in earth flows consists of fine soil contain-
ing a significant amount of entrained water. The propor-
tions of clay, silt, sand, and coarser material vary from
earth flow to earth flow, although silt- and (or) clay-size
grains predominate in most earth flows. Clay-mineral
composition also varies from site to site; earth flows
form in clays containing abundant kaolinite, illite, or
chlorite as well as in clays containing abundant mont-
morillonite. Earth-flow materials are pervasively fis-
sured; these fissures create planes of weakness and fa-
cilitate infiltration of water.

At two sites where appropriate measurements have
been made, earth-flow materials have low sensitivities.
A low sensitivity may distinguish earth-flow materials
from materials in other types of mass movement—par-
ticularly, “rapid earth flow” or “quick clay flow”—al—
though more measurements at other sites are needed to
confirm or disprove this hypothesis.

Mobilization of earth flows is accompanied by an in-
crease in water content. Measurements at one site indi-
cate that this water-content increase is due to in-place
saturation of the material without significant volume
change, remolding, or other disturbance.

Although some internal defamation occurs within
earth flows, most movement takes place on or im-
mediately adjacent to their boundaries. Conditions at
the boundaries, therefore, are of primary importance in

 

 

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. O

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REFERENCES CITED 53

determining how earth flows are mobilized and how
they move. Slope-stability analysis of soil mechanics ex-
plains how earth flows are mobilized. In this analysis,
the shear resistance of the soil is given by the Coulomb
failure criterion modified by Terzaghi’s theory of effec-
tive stress, which accounts for the effects of pore-water
pressure. Earth flows are mobilized by increases in
pore-water pressure along shear surfaces. At most
sites, the pore-water pressures necessary for
mobilization are generated by infiltration of water into
the soil. At a few sites, necessary pore-water pressures
are created by rapid loading in zones of depletion. This
rapid loading, which creates an undrained condition,
generates high pore-water pressures that allow earth
flows to move on slopes with inclinations of only a few
degrees. Movement of earth flowsl commonly ceases
when they encounter decreases in slope or when drain-
age or evaporation decreases pore-water pressure.

Earth flows are characterized by slow, persistent
movement; they generally move with velocities of a few
meters per day or less, and they remain active for sev-
eral days, months, or years. During this slow, persis-
tent movement, velocities vary over time; many earth
flows move faster during periods of high precipitation
than during drier periods. However, correlations be-
tween precipitation and velocity are complex, and earth-
flow velocity depends on several other factors, including
local boundary conditions.

In addition to slow, persistent movement, surges
occur on some earth flows. During these surges, which
last for a few minutes, velocities as high as 8 m/min have
been measured. The surges are caused by increases in
pore-water pressure that lower shear resistance below
the point at which an earth flow originally mobilizes.
The resulting imbalance between the driving force and
the available shear resistance causes an earth flow to ac-
celerate and leads to a high-velocity surge. Surges are
relatively infrequent, however, and most earth flows do
not accelerate continuously, even when pore-water
pressures rise above those required for mobilization.
Thus, in the general case, some mechanism acts to re-
tard earth-flow movement.

A retarding mechanism consistent with the avail-
able data involves boundary roughness and deformation
in thin zones adjacent to the boundaries. We

‘ hypothesize this mechanism to work as follows. As—

perities on boundary shear surfaces force material in

‘ zones adjacent to the shear surfaces to deform; these as-

perities provide a shear-resistance component that var-

‘ ies over time. In addition, material deforming around

the asperities may exhibit a component of shear
resistance that increases as the velocity increases. Thus,
both the asperities themselves and deformations in the

 

adjacent material retard the movement of an earth flow.
According to this boundary-roughness model, the
velocity of an earth flow is complexly controlled by sev-
eral factors, including the number, shape, and spacing of
asperities on the shear surfaces, the rheologic proper-
ties of the material that deforms around asperities, and
the shear-strength parameters and pore-water pressure
on shear surfaces.

A retarding mechanism consistent with the avail-
able data involves boundary roughness and deformation
in thin zones adjacent to the boundaries. We
hypothesize this mechanism to work as follows. As-
perities on boundary shear surfaces force material in
zones adjacent to the shear surfaces to deform; these as-
perities provide a shear-resistance component that var-
ies over time. In addition, material deforming around
the asperities may exhibit a component of shear
resistance that increases as the velocity increases. Thus,
both the asperities themselves and deformations in the
adjacent material retard the movement of an earth flow.
According to this boundary-roughness model, the
velocity of an earth flow is complexly controlled by sev—
eral factors, including the number, shape, and spacing of
asperities on the shear surfaces, the rheologic proper-
ties of the material that deforms around asperities, and
the shear-strength parameters and pore-water pressure
on shear surfaces.

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'8

A-.“..

t

‘6

REFERENCES CITED 55

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56 EARTH FLOWS: MORPHOLOGY, MOBILIZATION, AND MOVEMENT

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GPO 787—042/112

I”

 
 

  

GEOLOGICAL SURVEY

UNITED STATES DEPARTMENT OF THE INTERIOR

é'iiY‘EfrV“\
_ v? my“ ) (”Al/kg) _
’ it

v. w

122°00' \__‘ l A“

 

' '1 r: O a, ) PROFESSIONAL PAPER 1264

 

PLATE 1

 

 

  
  
   

 

t

yields barn
earth-flow

X complex

    

’1: fl.“ ‘~)\
I w m ow ~

 

brb

  

Davilla Hill
80 earth—flow
complex

    
   

 

Base from US. Geological Survey, 1:24 000
Dublin, Hayward, 1973

  

EXPLANATION

QUATERNARY DEPOSITS

 

af Artificial fill

 

ter Alluvial terrace deposits

Boundaries of earth—flow complexes

 

 

 

ea Active within past 3 years
er Recently active within past 3 to 15 years
60 Active more than 15 years ago or earth—flow scar

 

Deposits of other types of landslides

 

 

 

Ia Active
Ir Recent
lo Old

 

TERTIARY FORMATIONS—Generally covered by 2 to 4 feet of slopewash

 

bro Orinda Formation

 

 

brb Briones Formation

 

 

OTHER SYMBOLS

—— — Contact—Dashed where approximately located

—-|—1— —L .L Scarp—Solid where fresh; dashed where subdued. Hachures point downslope

g Regraded earth—flow material

/'\/'\/'\

Slumping

AAA

CEDDIEED) Gullying
O O O 0

Pi in and inci ient ull in
O O O P Q P 9 V 9

Soil creep

-—‘-—‘—~‘— Manmade cut

  

_‘

37°43’

 

 

 

 

 

~17: l22°00'

 
     

SCALE TIBOOO ‘1‘! lNTERlOR~GEOLOGlCAL SURVEY, RESTON, VlRGlN|A7T9837682622

 

I E . . _

E g; 0 500 1000 METERS . Geology and surfICIal deposns

g E, l l l Area of map mapped by R, W. Turnbull

g (22“ | | | l | (See Turnbull, 1976)
E O 500 1000 2000 3000 FEET

APPROXIMATE MEAN

DECLINATION, 1983

MAP SHOWING EARTH—FLOW COMPLEXES IN EDEN CANYON AND VICINITY, ALAMEDA COUNTY, CALIFORNIA

CONTOUR INTERVAL 20 AND 40 FEET
NATIONAL GEODETIC VERTICAL DATUM OF1929

  

   
  

UNITED STATES DEPARTMENT OF THE INTERIOR
GEOLOGICAL SURVEY

PROFESSIONAL PAPER 1264
PLATE 2

 

Ridgecrest

 

 

 

 
  

SCALE 1:750 000
10 20 30 4O 50 METERS
2—;

   
    

 

 
 

CONTOUR INTERVAL I METER
SPATIAL ACCURACY: HORIZONTAL 1 I METER
VERTICAL : 2 METERS

 
 
  

J300m

 

Base by Carlos Aboim 7 Costa and R. S. Stein, 1976

Geology from Aboim 7 Costa and Stein (1976)

 

EXPLANATION
Scarps—Hachures point downslope H—H Fenceline + Control point
‘H-v-I- Major bounding scarp 4*“ Crack, <0.5 m deep unless noted may Shrub
Trm'm' Steep scarp 73¢ Open crack or graben Wet, clayey soil
m Gentle scarp n74 Toe of earth—flow deposit @ Ponded water
ITIT Deep scarp M Trace of basal shear surface; hachures on upthrust r--J Ephemeral stream
block
33° .,, a. .
W Shallow scarp LL Striations, showing amount and direction of dip Hummocky topography

 

 

 

if INTERIORiGEOLOGICAL SURVEY, RESTON, VIRGIN|A71983 (382622

MAP AND CROSS SECTION OF THE HIDDEN VALLEY EARTH—FLOW COMPLEX,
ALAMEDA COUNTY, CALIFORNIA

 

 

PLATE

 
 

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1
Du
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P
A
P
L
A
N
m
S
S
E
F
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., . M .. .... w M .., e w
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. .. M ., . ., , w ., .n H
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OLOGICAL SURVEY

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$7 lNTERlORAGEOLOGICAL SURVEY, RESTON, VlRGlN|A719837682622

 

Geology by DK Keefer, 1975

Base by David K Keefer and Karen H. Keefer, 1975

HISTORY OF MOVEMENT

SCALE 1 :200

MORPHOLOGIC FEATURES AND LOCATIONS OF SURVEY—STAKE LINES, BOREHOLES, AND PITS

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32 METERS

12 14 16 18 20 22 24 26 28 30

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AREAOF MAP

 

CONTOUR lNTERVAL I METER
DATUM IS BASE OF LOWEST RECOGNIZABLE EARTH—FLOW TOE

MAPS SHOWING THE DAVILLA HILL EARTH—FLOW COMPLEX, ALAMEDA COUNTY, CALIFORNIA

 

 

 

 

Physiographic Divisions and
Differential Uplift in the
Piedmont and Blue Ridge

By JOHN T. HACK

 

GEOLOGICAL SURVEY PROFESSIONAL PAPER 1265

A .rtudy of the topography and other surface features
of the Piedmont and Blue Ridge, as related
to late tectonic history

 

 

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON:1982

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UNITED STA Es DEPARTMENT OF THE INTERIOR * ‘
‘ jAMES G. WATT, Secretary i .

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GEOLOGICAL SURVEY

 

 

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Dallas L. Peck, Director
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Library of Congress Cataloging in Publication Data ‘
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Hack, John Tilton, 1913—
Physiog'raphic divisions and differential uplift in the Piedmont and Blue Ridge. .V
(Geological Survey professional paper ; 1265) »
Bibliography: p.
Supt. of Docs. no. I 19.16:1265 *
1. Physical geography—Piedmont (U.S. : Region) 2. Physical geography—Blue Ridge Mountains. 3. Isostasy. I. Title. L y
11. Series. l
GB124.P53H33 551.4’3’0975 82—600079 )-
For sale by the Superintendent of Documents, U.S. Government Printing Office A v
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CONTENTS
Page Page
Abstract 1 Southern Blue Ridge province ___________________________ 27
Introduction 1 Chilhowee—Walden Creek belt ________________________ 28
Plan of report 2 Mountain highlands 28
Regional setting 6 Brevard zone and Hendersonville bulge _____________ 30
Piedmont Lowlands 10 Intermontane basins 32
Stream patterns and evidence for sedimentary overlap ____ 11 Mount Rogers area 34
Quantitatlve analySIS of drainage patterns __-____ ________ 12 New River plateau 3 4
Regional physiographic features related to the Pine Blue Ridge escarpment 35
Mountain belt 14 D R' tur 38
Differential movement along the Fall Zone _____________ 14 an _ iver cap e
Northeastern Highlands 15 Lmvxlle River 40
Southwestern Highlands 20 , Evolutlon of the escarpment ______________________ 44
Foothill zone 23 Time frame of late tectonism 45
Northern Blue Ridge Mountains _________________________ 25 References cited 46
ILLUSTRATIONS
Page
FIGURE 1. Map of the Southeastern United States, showing subdivisions of the Piedmont and Blue Ridge _________________________ 3
2. Generalized lithologic map of the Piedmont and Blue Ridge provinces 4
3. Averaged topographic map of part of the Southeastern United States 5
4. Relief map of the Piedmont and Blue Ridge provinces 7
5. Drainage map of part of the Appalachian Highlands 8
6. Topographic profile across the Blue Ridge and Piedmont from Salem, Va., to the Durham Triassic and Jurassic basin ...... 9
7. Topographic profile across the Blue Ridge from Johnson City, Tenn., to South Mountains, N.C _________________________ 10
8. Profiles of 14 streams in the Piedmont Lowlands ‘ 15
9. Profiles of five streams in the Northeastern Highlands 16
10. Map of the Patapsco River basin, Maryland 17
11. Logarithmic profile of the Patapsco River, Md 18
12. Logarithmic profile of the Appomattox River, Va 19
13. Transverse profiles of two valleys in the Patapsco River basin _ 20
14. Profiles of some first-order tributaries in the Patapsco River basin 20
15. Averaged topographic map of the Southwestern Highlands 21
16. Simplified topographic and geologic maps of the Southwestern Highlands 22
17. Longitudinal profiles of the Tallapoosa and Coosa Rivers 23
18. Logarithmic profile of the Tallapoosa River 24
19. Profile and geologic map along the crest of the Northern Blue Ridge 26
. 20. Map of the Southern Blue Ridge province 28
i 21. Profiles along two trench valleys 30
22. Map showing large trench valleys and basins in part of the Mountain highlands 31
23. Geologic map of the basin of the French Broad River 33
4 24. Map showing the asymmetry of the drainage system of the Southern Blue Ridge province 36
25. Strip maps showing the southeastern margin of the Southern Blue Ridge province 37
’4 26. Map of the headwaters area of the Dan River 39
i A 27. Profile of the headwaters of the Dan River 40
28. Map of the Linville River and surrounding areas 42
v 29. Profile of the Linville River 43
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, TABLES
L a. Page
TABLE 1. Mensuration data for drainage basins of the Piedmont and three selected basins in the Coastal Plain _____________________ 13
* 2. Comparison of topography at four locations in the Asheville basin 34

 

, III

 

 

PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIF T IN
THE PIEDMONT AND BLUE RIDGE

By JOHN T. HACK

ABSTRACT

The Piedmont and Blue Ridge are dynamic landscapes that have
undergone substantial change since the orogenies that ended in late
Paleozoic or, as some believe, early Mesozoic time. The southern Blue
Ridge region south of Roanoke, Va., lies on the crest of a topographic
uplift that corresponds to the eastern continental drainage divide. To
the north, this uplift and divide cross the Appalachian Valley and form
the crest of the Appalachian Plateaus as far north as central Penn-
sylvania. The northern Blue Ridge Mountains as well as parts of the
Piedmont are on the eastern part of the uplift area. The southeastern
margin of the uplift corresponds to a line within the Piedmont
physiographic province that extends northeastward from the
Tallapoosa River at the Fall Zone and crosses the Rappahannock River
at the Fall Zone. The differential elevation on either side of this line is
sharp in some places, as, for example, northeast of Atlanta, Ga. In
other places, the difference in elevation is difficult to detect, and, in ef-
fect, the line becomes a broad monoclinal slope.

The region as a whole can be divided into at least six broad
subregions that have somewhat different histories in late geologic
time.

The Piedmont Lowlands subprovince, southeast of the uplifted area,
is dominated by a monotonous topography of low rounded ridges and
ravines largely underlain by saprolite on crystalline rocks. Isolated
ranges of hills of greater relief are scattered across the region; those
investigated are directly related to the presence of erosionally resist-
ant rocks. Stream patterns as well as broad topographic forms indicate
that although the southern part of the Piedmont Lowlands was prob
ably once covered by younger sediments, this area has been exposed to
erosion for a long time. In North Carolina, the inner part of the Pied-
mont Lowlands has strongly trellised stream patterns, which suggest
that subaerial erosion was active for an even longer time period,
perhaps since the latest orogeny. North of the Cape Fear River, the
outer part of the Piedmont Lowlands was covered by either fluvial or
marine sediments or both, probably during Miocene time. Tectonic
activity has affected the Piedmont Lowlands in late geologic time. The
Fall Zone that forms the southeast border is, at least in places, con-
trolled by faults active in Tertiary time. Late faults have also been
found in the Pine Mountain area of Georgia. Minor differences in relief
affecting large regions within the Piedmont Lowlands may be related
to different rates of uplift in addition to rock resistance, either past or
present.

The Piedmont northeast of the Potomac River (Northeastern
Highlands) rises to more than 300 m in altitude. The major streams
have convex profiles that steepen as they near the Coastal Plain.
Unusually narrow valleys and broad upland surfaces indicate an in-
creased rate of erosion and show that the relief is now or recently has
been increasing because of uplift or tilting.

West of the southern end of the Piedmont Lowlands is an area herein
called the Southwestern Highlands that in some respects is similar.
The area is crossed by two large streams that have convex profiles.
The highest mountain ranges in the area rise to altitudes greater than
600 m.

 

Northwest of the Piedmont Lowlands, the topography and relief are
higher, and in some places, the rise is gradual, forming a Foothill zone
between the Piedmont Lowlands and the high Blue Ridge. This zone is
morphologically mare complex than the Piedmont Lowlands. North of
the Roanoke River, the foothills are commonly chains of isolated hills
and ridges generally underlain by resistant rocks. The hills increase in
height near the Blue Ridge, an indication that they owe their height to
tectonism of late geologic age. South of the Yadkin River, the hills are
believed to be residual, the remnants of a larger highland that has been
only partially reduced to the lower relief of the general Piedmont sur-
face.

The Blue Ridge physiographic province has two subdivisions. The
Northern Blue Ridge Mountains are underlain by resistant rocks and
in effect form the easternmost range of the Valley and Ridge province.
Three different rock types form the crest in various places; the dif-
ferences in their resistance as well as in the width of outcrop belts
determine the altitude of the range. North of Chester Gap in northern
Virginia, the lower altitude of the range may be related in part to dif-
ferences in uplift history over a broad region.

The Southern Blue Ridge province is a complex area. Part of it, the
New River basin, is essentially a plateau that owes its relative height
to resistant rocks crossed by the New River downstream from the Blue
Ridge itself. The southern part is underlain by massive and resistant
gneiss and metasedimentary rock and has relief exceeding 1,100 m.
Rock control has a strong influence on forms locally. Deep trenchlike
valleys follow weak-rock outcrop areas like the carbonate rock of the
Murphy marble belt as well as fracture zones or shear zones along
faults. The Asheville basin may be partly controlled by the outcrop of
weak rocks but also by intersecting shear zones.

The Southern Blue Ridge province is bounded on its southeast
margin by a bold escarpment that corresponds to the major drainage
divide. As shown by a major piracy by the Dan River, at least the
northeastern part of this escarpment has retreated toward the north-
west. At the Grandfather Mountain window and south of it, the escarp
ment may be stabilized along a zone of resistant rocks. The origin of
the escarpment is probably a monoclinal flexure or a movement along a
series of undetected faults as much as 30 km or more southeast of the
present position of the escarpment.

The time frame in which the uplift in the region took place can be
estimated from data on erosion rates in areas of different relief, as sug-
gested by Ahnert’s concept relating rates of relief reduction, rates of
uplift, and rates of erosion. His analysis shows that if uplift is not con-
tinuous, any area such as the Blue Ridge would be reduced to 10 per—
cent of its initial relief in a period of only about 20 million years. This
means either that uplift of the areas northwest of the Piedmont
Lowlands has taken place since Miocene time, or, more likely, that it
has gone on continuously at different rates, the highest rate being in
Miocene time.

INTRODUCTION

Geologists have long thought that the late tectonic
history of the Appalachian Mountains involved vertical

1

2 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

uplift. It is difficult to escape this conclusion because the
mountain system declines in both altitude and relief
near its southwestern terminus to become buried
beneath a cover of Cretaceous and Tertiary sediments.
Many studies of the forms of individual areas within the
mountain chain have been based on the assumption that
the hilltops were relict features inherited from surfaces
having much lower relief, called “peneplains.” The
degree of dissection of these peneplains was considered
a measure of the uplift. The times and average rates of
successive periods of uplift could be deciphered if the
age of the peneplains could be established. This age was
rarely possible to determine by stratigraphic methods in
the Appalachians because of the scarcity of fossiliferous
postorogenic sediments in the erosional landscape.

The existence of relict forms inherited from older
landscapes of the kind postulated by W. M. Davis and his
followers is questionable. The ubiquitous pattern of in-
terconnecting valleys separated by branching ridge
systems throughout the region is an indication that this
topographic form prevails and probably always has.
Regional differences in topographic relief are signifi-
cant. Many can be explained by differences in erosional
resistance of the bedrock. Some can be explained only by
differences in rates or history of uplift (Hack, 1960).
This study attempts to identify different erosional
systems within the crystalline rock area of the
unglaciated Appalachians and their tectonic relation.

My earlier work in the Appalachian region has con-
vinced me that the present landscape is dynamic and
that most features can be explained by a balance of
forces involving uplift, erosion, and rock resistance. The
landscape chosen for study at that time was mostly in
the Ridge and Valley province, where extreme dif-
ferences in rock resistance are obvious (for example,
limestone or shale as compared with quartzite). More
subtle tectonic effects were not detected, and this aspect
of the problem was not addressed except to acknowl-
edge that it existed.

Ahnert (1970) has critically evaluated the relation of
erosion and uplift in areas of differing rainfall, average
slope, and relief. He has shown that topographic relief in
a variety of areas having complex geology has a close
linear relation to rate of denudation as measured by
present rates of debris transport in streams. Differences
in climate between the areas apparently have little ef-
fect on erosion rates. Average slope is related to
denudation rates but not as closely as relief. Ahnert
pointed out that although erosion rates and relief are
probably not evenly balanced so as to achieve a steady
state, a tendency toward such a balance exists. His
analysis also indicates that, in part because of isostatic
compensation, it would require periods on the order of
18 million to 20 million years to reduce a landscape to 10

 

percent of its initial relief if no tectonism were involved.
If his analysis is valid, high relief should be correlated
with high uplift rates at present or at some time in the
not-too-distant past.

In the discussion of the Piedmont and Blue Ridge that
follows, I assume that differences in relief from place to
place may be the result of differences in rock resistance,
which I refer to as rock control (Y atsu, 1966). On the
other hand, differences in relief may be due to different
rates or different histories of tectonic uplift. However, it
is assumed that all high-relief features must be
postorogenic, that is, Mesozoic or Cenozoic. Further-
more, areas having differences in relief must have had
different tectonic histories if it can be shown that the
differences were not caused by rock control.

Evidence other than relief can be used to analyze tec-
tonic effects and histories. Stream patterns may reflect
stages during downwasting. Stream-profile analysis
proves to be a powerful tool with which to identify possi-
ble tectonic movement, because the steepness of the
channel slope for a given discharge is a measure of
stream power that may, like relief, be related to either
the resistance of the rock through which the stream
flows or to local uplift or tilting. Faults, folds, and other
local features are, of course, also manifestations of tee-
tonism.

The writer is indebted to many colleagues in the US.
Geological Survey for advice on the geology of many
areas, and especially to Milan Pavich, Helaine
Markewich, and Douglas Rankin for their reviews of the
manuscript.

PLAN OF REPORT

The Piedmont and Blue Ridge provinces, together,
can be divided into subprovinces or areas having dis—
tinctly different topographic forms. On the basis of the
geology, topography, and stream systems, six major
subprovinces (fig. 1) have been identified.

The discussion is illustrated by four maps of the
region. Figure 2 is a highly generalized lithologic map
adapted from the “Tectonic Lithofacies Map of the Ap-
palachian Orogen,” by Williams (1978). This map was
chosen because it is probably as close to a lithologic map
as can be constructed at so small a scale.

Figure 3 is an averaged topographic map that shows
the topographic setting of the Piedmont and Blue Ridge
provinces and their relation to the adjoining topography
of the folded Appalachians and the Appalachian
Plateaus. Contours on the basement rocks beneath the
Coastal Plain are also shown. The subaerial altitudes are
adapted from an averaged contour map having 50-m
contours that show mean altitudes. The map shows the
mean altitude within 3-minute rectangles (approximate-

 

PLAN OF REPORT

  
 
  
  
 
 
   
 
  
  
 
 

TRIASSIC and JURASSIC BASINS
(dotted areas)

Ne Newark basin
Ge Gettysburg basin
Cu Culpeper basin
Ri Richmond basin
Fa Farmville basin
Da Danville basin
DS Durham—Sanford basin
Wa Wadesboro basin

 

 

 

\
\ o 100 200 300 KILOMETEHS
l l 1 L I
l\..

EXPLA NATION

 

 

 

 

 

 

 

TALLAPOOSA-RAPPAHANNOCK LINE
SUBDIVISIONS OF PIEDMONT LOWLANDS

CA Carolina slate belt OP Outer Piedmont of North Carolina and Virginia
UW Uwharrie Mountains KM Kings Mountain belt
PM Pine Mountain belt

FIGURE 1.—Ma.p of Southeastern United States showing the subprovinces within the Piedmont and Blue Ridge and the Tallapoosa-
Rappahannock Line.

PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

EXPLANATION

{13* Crystalline basement

  
   

2 Clastic and volcanic rocks of low to medium
metamorphic grade

Q
E] Carbonate rocks
[1

Mostly highly deformed metasedimentary
and metavolcanic rpcks

 

 

 

vv Metasedimentary rocks, chlorite-schist, 82°

slate, and volcanic rocks 38°

 

 

 

 

\ 0 100 ZIIJD Blifl KILOMETEHS
|

 

32°;

 

 

 

 

 

 

Metasedimentary rocks, migmatites, and El Mainly quartz-feldspar gneiss and schist
quartz-kyanite rocks

Relatively undeformed sedimentary and Cataclastic rocks
volcanic rocks (Carolina volcanic slate belt)

\I ”l,

Plutonic rocks m Triassic and Jurassic sedimentary and
‘ " volcanic rocks

FIGURE 2. —Generalized lithologic map of the Piedmont and Blue Ridge provinces. Data from Williams (1978).

 

PLAN OF REPORT

 

74°
4’

76°

78°

 

800

s.
AS? I l l
sér;
\AQV A,
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. Vi mas}.
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400 KILOMETERS

l

 

Contout interval is 150 m between 0 and 450
Contour interval is 152 m between 0 and —1520

EXPLANATION

Areas In the Coastal Plain underlain

by Miocene and younger deposits

Over 1250 m

0—1250 m

\\\\“

\A

s\\\‘
450—850 m

NEGATIVE CONTOURS NE OF SAVANNAH RIVER ARE ON BASEMENT. NEGATIVE CONTOURS SW OF SAVANNAH

RIVER ARE ON BASE OF BEDS OF AUSTIN AGE

 

d data (original contour map furnished by M. F.

, and the Appalachian Plateaus. The contours on the land surface

f the Southeastern United States, showing the crystalline rock area of the Piedmont and

Blue Ridge in relation to the Coastal Plain, the Valley and Ridge
show mean altitudes within 3-minute rectangles and were prepared from digitize

0
m.
f
0
w
m
.w
.m.
m
m
t
.w
m.
a
v
A
—
om
E
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U
m
F

Kane). Generalized contours on the basement are from the “Tectonic map of the United States" (US. Geological Survey and American

Association of Petroleum Geologists, 1961).

6 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

1y 6x 5 km in size). The means are computed from
thousands of points within each rectangle. This kind of
map, of course, does not show the true topography. It
leaves out many details that, if shown, would tend to
obscure the general topographic form. Such local details
commonly are related to rock control of the topography.
Thus, a generalized map may be emphasizing forms in
part inherited from or caused by tectonic movements.

Figure 4 is a relief map made visually by counting the
numbers of 100-ft contours within each 100-km2 unit of
the Universal Transverse Mercator grid, as shown on
the 1:250,000-scale 1° to 2° maps of the U.S. Army Map
Service and the U.S. Geological Survey. A relief map is
quite different from a topographic map in that the
values shown vary depending on the definition of the
units used. Probably the best method of making a relief
map would be to measure or estimate the difference be-
tween ridge crests and adjacent valley bottoms. Such
measurements could be made by making a series of
detailed topographic profiles across an area and
subtracting the altitudes of the low points from those of
the high points along the profiles (see fig. 6 for an exam-
ple of such a profile). However, the work involved for so
large an area is prohibitive. The method used here gives
higher relief values, but the purpose of the map is to
distinguish large areas of different relief, particularly
within the monotonously low Piedmont province.
Relative values can serve this purpose.

Figure 5 is a map of the drainage system of the Pied-
mont and Blue Ridge that includes all the headwater
areas of drainage flowing across the region. The data
are from the outline maps of the States (U.S. Geological
Survey 121,000,000 series).

Figures 6 and 7 are detailed profiles across two areas
in the Blue Ridge and Piedmont. These profiles were
made using large-scale (1:24,000 and 1:62,500)
quadrangle maps by the U.S. Geological Survey. Their
locations are shown in figure 4. Other maps of smaller
areas are used to illustrate specific areas discussed in
the text.

REGIONAL SETTING

The Piedmont and Blue Ridge form an elongate area
in the Southeastern United States, 100 to 300 km wide
and about 1,375 km long. As the averaged topographic
map (fig. 3) shows, the region is asymmetric in the ver-
tical dimension, the high area known as the Blue Ridge
being along the northeast margin. A close examination
of the map shows that the crest of this range does not
correspond to the highest average topography. The
topographic high is roughly parallel to the present
coastline, whereas the Blue Ridge trends in a more
northeasterly direction. Judson (1975) has noted that
the general topographic high south of central Penn-

 

sylvania may have been formed in early Mesozoic time,
as it is related to a low-gravity anomaly and also cor-
responds to the drainage divide between the Atlantic
and Gulf of Mexico drainage. I have suggested that the
major Tertiary uplift of the Appalachians also crested
along this drainage divide. The details of the forms of in-
dividual narrow ranges were determined by rock control
(Hack, 1980, p. B14—B15).

The line separating the Piedmont Lowlands from the
other divisions of the crystalline rock area (fig. 1) ap-
pears to be a major dividing line in the region. It is ex-
pressed in the topography by a slight steepening of the
regional slope. At the southwestern end of the region,
the line terminates approximately Where the Tallapoosa
River leaves the Piedmont. On the northeast, it inter-
sects the Fall Zone where the Rappahannock River
leaves the Piedmont. In Georgia, the altitude of this line
is about 300 m. It declines to 180 m in southern Virginia
and to sea level at its northern end. This line roughly
forms the southeastern boundary of regions having high
relief, although it becomes a little vague in central
Virginia. The Tallapoosa—Rappahannock line also cor-
responds in places to a major line on Williams’ (1978)
tectonic lithofacies map of the Appalachian orogen. It is
near the south boundary of his units 4 and 5 (unit 4, fig.
2, in this paper), except in southern North Carolina,
where it is in Williams’ unit 23 (unit 9 in fig. 2).

The break in slope of the basement rocks at the inner
edge of the Coastal Plain is a striking feature (fig. 3).
These basement rocks are an eastward extension of the
rocks of the Piedmont. The sharp increase in slope in
places is the same order of magnitude as the increase in
slope of the inner Piedmont and the rise to the Blue
Ridge crest. Northwest of the deep Salisbury embay-
ment, however, the rise of the subaerial topographic sur-
face is steep only at the edge of the Coastal Plain. A
short distance inland, the rise to the Applachian crest in
Pennsylvania is gradual, and the general topographic
surface is lower than it is to the south. The Salisbury em-
bayment collects the drainage of some of the largest
rivers in the Appalachians, including the Delaware, Sus-
quehanna, Potomac, James, and Roanoke. This embay-
ment and its extension to the south in North Carolina
received great thicknesses of sediment, particularly in
Cretaceous and Miocene time (Owens, 1970). In late Ter-
tiary time, it also received large quantities of sediment,
including sediment from the Hudson River system
(Owens and Minard, 1979).

The decline in altitude of the Appalachian crest in
Maryland and Pennsylvania may be related to the tec-
tonic system that formed the Salisbury embayment. The
embayment and a possible troughlike area to landward
may have been responsible for gathering a fanlike net-
work of large rivers.

 

REGIONAL SETTING

 

 

 

34°

 

 

 

 

\ I
32° I’ l J 0

 

Steep hills and mountains
(relief more than 50 m in
6-km2 area)

 

Moderate relief
( relief averages 150 m in
100-km2 area)

 

 

 

 

 

 

 

200 300 KlLOMETERS
l___l____L.__l
EXPLANATION
'- Low relief

 

 

 

 

 

 

 

( relief averages 60 m in
100—km2 area)

'- '_ Low relief

( relief averages 30 m in
100-km2 area)

FIGURE 4.—Relief map of the Piedmont and Blue Ridge provinces (based on the number of 100-ft contours within 100-km2 areas, as
shown on topographic maps at a scale of 1:250,000). Lines A—A' and B-B’ refer to topographic profiles of figures 6 and 7.

PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

a 74°
78° 7? \

 

_~____

34°—

32°—

   

-o

100 ZI'JU 300 400 KILDMETERS
I I I

 

EXPLANATION

.- ..~..’--'—~v

ATLANTIC-GULF DlVlDE

TALLAPOOSA-RAPPAHANNOCK LINE

FIGURES 5. —Drainage map of part of the Appalachian highlands (shaded area), including the Piedmont and Blue Ridge. Rivers are iden-
tified by numbers: 1, Delaware; 2, Schuylkill; 3, Susquehanna; 4, Patapsco; 5, Potomac; 6, Rappahannock; 7, James; 8, Appomattox; 9,
Meherrin; 10, Roanoke; 10a, Dan; 11, Neuse; 12, Cape Fear; 12a, Haw; 12b, Deep; 13, Yadkin; 14, Catawba; 15, Broad; 16, Saluda; 17,
Savannah; 18, Oconee; 19, 0cmulgee; 20, Flint; 21, Chattahoochee; 22, Tallapoosa; 23, Coosa; 24, New; 25, French Broad; 26,
Tuckasegee; 27, Hiwassee; 28, Ebowah; 29, Tennessee.

 

REGIONAL SETTING

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10 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

B

METERS
2000 -

Basement
rocks

H—i

1800 - Appalachian

Valley

1600 -

Chilhowee
Group

1400 -

1200 '-

1000 —

800 -

600 -

400 --

200 -

 

 

 

Ashe Formation F

Basement rocks

Grandfather Mountain window
Morga nton basin

South Mountains "'

 

 

 

 

l_ l I

50 KILOMETERS
I l J

Vertical exaggeration X386

FIGURE 7.—Topographic profile across the Blue Ridge from the Appalachian Valley near Johnson City, Tenn., to South Mountains south of
Morganton, NC. See figure 4 for location (B—B’).

A similar but less impressive coalescence of streams
occurs landward of the embayment southwest of the
Cape Fear arch (fig. 3), where several large Piedmont
streams join to form the Peedee and Santee Rivers.
These two rivers are less than 10 km apart at the Atlan-
tic coast (see rivers 13 to 16, fig. 5). If the basins of these
rivers were at one time gathered together in a trough-
like area, that trough is not defined in the present
topography and has been destroyed by erosion.

PIEDMONT LOWLANDS

The Piedmont Lowlands comprise most of the area
southeast of Williams’ (1978) unit 8, which consists of
polydeformed and metamorphosed sedimentary rocks
that he suggested may have marked the eastern border
of the North American continent in early Paleozoic time.

 

The rocks in the Piedmont Lowlands are predominantly
feldspathic gneiss and schist intruded by plutons, most
of which are granitic, but the area also includes sedimen-
tary and volcanic rocks of lower metamorphic grade.
The topographic expression of these rocks is different
from that of the gneiss and schist. The largest area of
low-grade sedimentary and volcanic rock is the Carolina
volcanic slate belt (figs. 1 and 2), which is mostly a low-
relief area underlain by argillite, slate, and tuffaceous
rocks. In the northern part, feldspathic intrusive and
volcanic rocks are more abundant. These rocks are quite
resistant and generally underlie hilly topography. The
Uwharrie Mountains (fig. 1) exceed 300 m in altitude
and have sharp relief. Many hills rise 90 m above their
bases. The importance of rock control in shaping hilly
areas as well as individual hills is quite evident and is
well shown on the geologic map of the Denton

p

PIEDMONT LOWLANDS 1 1

quadrangle (Stromquist and others, 1971) and in the
field guide by Seiders and Wright (1977).

The Kings Mountain belt is an elongate area 15 to 30
km wide. It consists mostly of metasedimentary rocks,
but it also contains granite gneiss, biotite gneiss,
metamorphosed quartz diorite, and intrusive granitic
bodies. The metasedimentary rocks are primarily
nonresistant rocks, such as metasiltstone, phyllite, and
marble, but resistant rocks, such as quartzite, kyanite
quartz rock, and conglomerate, also occur in the belt.
The resistant rocks form chains of low hills. Crowders
Mountain near Gastonia, N.C., has a relief of 183 m and
rises to an altitude of 460 m. The crest of this mountain,
as well as that of Kings Mountain southwest of it, is held
up by narrow outcrops of quartzite and conglomerate.
Clasts of these rocks cover parts of the lower slopes. As
the correlation between local relief and rock type is pro-
nounced in the Kings Mountain belt, I conclude that the
unusual topography in this belt is probably controlled by
rock resistance.

The Pine Mountain belt of Georgia and Alabama is
considered by some geologists as part of the Kings
Mountain belt. Although the two are not contiguous,
they may be connected by the Towaliga fault (Hatcher,
1972). The rocks in the Pine Mountain belt are similar to
those in the Kings Mountain belt as they include
micaschist, gneiss, and extensive, although thin, beds of
quartzite. An important feature of the area is a
downfaulted block containing unconsolidated fine-
grained fluvial deposits of Paleocene age on the north
border of the belt. These deposits constitute a basin
more than 3 km long and 0.8 km wide. They are bounded
by a major fault (Towaliga fault) on the north, and by a
local fault (Warm Springs fault) on the south. The
Paleocene sediments are overlain by sand and gravel
washed from Pine Mountain. These deposits are also af-
fected by the Warm Springs fault (White, 1965;
Christopher and others, 1980). This occurrence of fine-
grained Paleocene sediments shows that part of the
Georgia Piedmont 40 km north of the Fall Zone had a
cover of Tertiary sediments and also that tectonic defor-
mation took place within the Piedmont in Tertiary time.
On the other hand, the high topographic areas within the
belt, such as Pine Mountain and Oak Mountain, probably
do not owe their altitudes directly to the faults. The high
areas are localized along outcrop belts of quartzite and
alluvial fans of quartzite Clasts. The maximum relief is
150 m, equivalent to the relief at Crowders Mountain in
the Kings Mountain belt.

A few small hilly areas other than those in the
Carolina slate belt, the Pine Mountain belt, and the
Kings Mountain belt also occur, commonly related to
small lenticular zones of quartz-kyanite rock. Willis
Mountain in central Virginia is perhaps the most spec-

 

tacular; it consists of a narrow range of hills more than
40 m high that rises to an altitude of 345 m above sea
level (Espenshade and Potter, 1960).

The Piedmont Lowlands as defined here (fig. 1) con—
tain several Triassic and Jurassic basins. Except for the
Danville basin, the relief of these basins is generally
lower than that of the crystalline rock areas surrounding
them. The Danville basin, a long narrow basin at the
western margin of the area, contains 4,500 m of con-
tinental sedimentary rocks, much of which are arkose
and graywacke. Because of the arenaceous nature of the
rocks, the relief is high, ranging from 45 to 140 m.
Whiteoak Mountain, the highest range of hills, is higher
than the adjacent Piedmont, but, in general, the relief is
comparable with that in the Piedmont (as shown in the
cross section, fig. 6). The Farmville and Richmond
basins are underlain by shale, thin sandstone, and coal.
The relief within areas of 100 km2 averages about 55 m,
in some places reaching 85 m, but differs little from that
of the surrounding Piedmont.

The Durham-Sanford basin has generally low relief,
somewhat lower than that of the Piedmont to the west
(fig. 6.) The rocks range from claystone and shale to
fanglomerate. They include carbonaceous clay and
mineable coal. The Wadesboro basin is separated from
the Sanford basin by overlapping sediments of the
Coastal Plain. Its relief is also somewhat lower than
than of the adjacent Piedmont and can be accounted for
by the lower resistance of the rocks.

STREAM PATTERNS AND EVIDENCE
FOR SEDIMENTARY OVERLAP

The river patterns of the Piedmont Lowlands (fig. 5)
support the idea that parts of the region were covered
by overlapping sediments in late geologic time. In con-
trast to much of the drainage northwest of the lowlands,
the major streams flow directly down the slope of the
Piedmont toward the Coastal Plain; two distinct pat-
terns can be clearly distinguished, one southwest of the
Cape Fear River (fig. 5, no. 12) and the other north of it.

Southwest of the Cape Fear River, almost the entire
Piedmont Lowlands area is drained by large streams
that head at the Atlantic—Gulf of Mexico divide. The ma-
jor trunk streams of highest order flow in rather
straight courses down the regional slope; no structural
control is apparent. Many of the lower order streams
join them at fairly wide angles, and first- and second-
order streams commonly are parallel to the prevailing
northeast structural trend. This kind of adjustment is
especially common in the Slate belt. Only major streams
drain into the Coastal Plain, for the inner edge of the
plain is a cuesta that is higher than the Piedmont north-
west of it. The cuesta is generally penetrated only by the

12 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

larger streams, and its northwest slope drains into long
tributaries of the major Piedmont streams.

Staheli (1976) has suggested that the drainage of the
outer Piedmont of Georgia was superimposed from a
Coastal Plain cover and extended inland about 190 km,
almost to the Blue Ridge front, a concept that explains
the lack of structural control of the master streams. The
work of Christopher and others (1980) at Warm Springs
supports Staheli’s concept, because, as they state (citing
Cramer, 1979), updip projections of the Paleocene
Nanafalia Formation in Georgia suggest that a blanket
of sediment did exist in this area. Furthermore, the
crossbeds in the Paleocene fluviatile sediments north of
Pine Mountain indicate a flow to the south and the
absence of a topographic high at that time.

The superposition hypothesis can probably be ex-
tended through South Carolina because the outer Pied-
mont has a drainage pattern similar to that found in
Georgia northward as far as the Cape Fear River. The
major streams of this drainage system, as will be shown,
have basins that are broader than those of most conse-
quent streams of similar length. This width probably is
related to the fact that the drainage is perpendicular to
the structural trend of the rocks across which they flow.
If deep erosion, several hundred meters or more, took
place after the original drainage was superimposed on
the Piedmont rocks, there would be a strong tendency
for tributaries to form parallel to rock structure and at
large angles to the main stream at the expense of
streams that were perpendicular to the structure. The
process would begin as soon as the hard rocks were
penetrated by the trunk streams.

A different drainage pattern exists northeast of the
Cape Fear River, at least as far as the Rappahannock
River (fig. 5, no. 6). In this region, the belt of northeast-
trending streams along the inner Piedmont (fig. 5)
widens. However, the major rivers that head at the
Atlantic—Gulf of Mexico divide, such as the Roanoke and
the James Rivers, flow across the outer Piedmont in
rather direct courses. Between these streams are many
smaller streams that head in the Piedmont but that flow
directly into the Coastal Plain. The Piedmont—Coastal
Plain boundary is closer to the ocean in this region than
it is farther south, and the Coastal Plain sediments
overlap the Piedmont without any apparent cuesta. The
protrusion of the Piedmont margin, or Fall Zone, to the
east corresponds to the inner edge of the Cape Fear arch
(fig. 3.) North of the protrusion, patches of fine-grained
sediments and river gravels are found some distance
west of the Coastal Plain margin (for example, Goodwin,
1970). These features suggest a Coastal Plain overlap of
the outer part of the Piedmont that is much younger
than the overlap to the south, as explained below. A nar-
rower overlap in the Washington, D.C., area, that is

 

probably of the same age (described by Darton, 1951,
and discussed by Hack, 1975), is believed to be of late
Tertiary age.

QUANTITATIVE ANALYSIS OF DRAINAGE PATTERNS

The differences in the two drainage patterns de-
scribed above can be further defined by an analysis of
river-basin shapes. One way to compare the shape of
drainage basins is simply to measure their lengths along
the main stream and then to calculate the average width
by dividing the drainage area by the measured length.
The ratio of width to length would be a measure of basin
shape. These calculations have been done for 11 basins
of the outer Piedmont in Virginia and North Carolina
(table 1) as well as for 9 of the larger basins north of the
Cape Fear River. As expected, the average width of the
southern basins is greater than that of the northern
basins. The average width-length ratio is 0.28 south of
and including the Cape Fear basin and 0.19 north of the
Cape Fear basin. This difference is sizable, but, as the
data show, the scatter between individual basins is
large. One reason for the wide range is the fact that
basin shape is partly a function of basin size. Con-
siderable research has shown that, as basins enlarge
downstream, they tend to develop lower width-length
ratios. In other words, large basins are commonly more
elongate than small ones. If there are no geologic re-
strictions on basin shape, basins 100 km2 in area on the
average have a width-length ratio of 0.33, whereas
basins 10,000 km? in area on the average have a width-
length ratio of only 0.16.

This characteristic of drainage basins was studied by
Hack (1957) using data on drainage areas and stream
lengths as measured along the stream, including bends
and meanders. The data included 92 measurements in
the Shenandoah Valley as well as data obtained by
Langbein (1947) for 400 localities in the Northeastern
United States that included streams whose basins were
more than 2,000 mi2 (5,180 km?) in area; stream length
was found on the average to approximate the function
L = 1.4 A”, where L is length along the stream in miles
as measured from the head of the longest stream in the
basin and A is the drainage area of the basin in square
miles. As this equation is not dimensionally balanced,
the constant of proportionality (in this case 1.4) varies
with the units of measurement. In kilometers and
square kilometers, the equation is L = 1.27 A“. Leopold
and Langbein (1962) showed that drainage systems
simulated by constructing random walks become more
elongate as they enlarge downstream. Their example
suggested that stream length was proportional to the
0.64 power of the basin area. Hack (1965), in a study of a
stream system in grooved glacial and lacustrine deposits
9,000 years old and younger, showed that restraints on

,A‘I ‘ :-

 

PIEDMONT LOWLANDS 13

TABLE 1.—Mensurat'ion data for drainage basins of the Piedmont and three selected basins in the Coastal Plain
[Piedmont basins are assumed to terminate at the Fall Zone]

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

Average Predicted
Basin Length Area width length-law Difference Width-length
(km) (km‘) (km) width (percent) ratio
Piedmont south of Cape Fear River
Deep-Cape Fear, N.C. 136 9,680 71 40 + 77 0.52
Yadkin, N.C. 317 18,663 58 53 +9 .19
Catawba, N.C.—SC. 279 11,691 42 43 —2 .15
Broad, SC. 241 13,263 55 46 + 19 .23
Saluda, SC. 226 6,612 29 34 — 14 .13
Savannah, Ga. 233 18,277 78 53 + 47 .54
Oconee, Ga. 178 7,337 41 36 +14 .23
Ocmulgee, Ga. 142 6,128 43 33 + 30 .30
Flint, Ga. 137 5,239 38 31 + 23 .28
Outer Piedmont north of Cape Fear River
Po River, Va. 36 297 8.2 9.2 - 11 0.23
North Anna, Va. 71 1,052 15 16 —6 .21
Little River, Va. 38 281 7.3 8.9 — 18 .19
South Anna, Va. 90 1,113 12 16 — 25 .13
Appomattox, Va. 162 3,387 21 26 — 19 .13
Nottoway, Va. 70 1,193 17 16 + 6 .24
Meherrin, Va. 86 1,797 21 20 +5 .24
Fishing Creek, N.C. 64 1,161 7.6 11 —30 .12
Swift Creek, N.C. 61 464 18 16 +12 .28
Tar, N.C. 114 1,793 16 20 — 20 .14
Neuse, N.C. 142 4,430 31 29 + 7 .22
Coastal Plain

Edisto, SC. 211 7,457 35 36 - 2 0.17
Salkahatchie, 8.0. 129 3,129 24 25 — 4 .19
Ogeechee, Ga. 267 7,684 29 36 — 19 .11

 

 

the lateral development of stream systems changed the
value of the constant of proportionality and, to a smaller
extent, changed the value of the exponent. Also many
large basins, especially very large ones, tend to enlarge
in area relative to length as they grow, and thus may
have exponents less than 0.5 (Mueller, 1972, 1973; Mose-
ly and Parker, 1973). Very large basins are more likely
to be restricted by tectonic, geologic, and other factors
than are small basins, but, of course, even small basins
can be influenced by various geologic factors. Recently,
a study of 155 drainage basins in Hokkaido, Japan
(Shimano, 1975), showed that an empirical equation
relating stream length and basin area L = 1.413 A051, ex-
pressed in kilometers and square kilometers, fit the data
closely. Shimano’s work, however, indicates that errors
in measuring meanders of different sizes do affect the
relationship by increasing the value of the exponent
slightly. Shimano also constructed an equation for basin
length and area L = 1.224 A0-576 in kilometers and square
kilometers. This equation avoids the problems involved
in measuring meandering streams and is used in the
analysis that follows.

In table 1, the average width of the drainage basins

 

south of and including the Cape Fear River basin and 11
of the basins north of the Cape Fear River are compared

with the basin width that could be predicted if their
dimensions followed the length law as determined by
Shimano (1975). Three large rivers of the Coastal Plain
are also included in the table. The equation for normal or
predictable basin width can be determined directly from
the basin area by using Shimano’s length-width equation
and is W: 0.82A0-424.

As the table shows, the streams south of and including
the Cape Fear basin are mostly much wider than the
predictable width based on the length law as determined
by Shimano. This is interpreted to mean that the basins
have a tendency to form long tributaries because of
some structural control of all but the major streams.
Two exceptions are the Saluda River basin (14 percent
narrower than the length-law width) and the Catawba
River basin (2 percent narrower than the length-law
width). The shape of the Saluda River basin is readily ex-
plained, for that river flows in and parallel to a swarm of
diabase dikes of Triassic age shown on the geologic map
of the United States (King and Beikman, 1974). These
dikes can be expected to have strongly restricted the
lateral growth of the basin. The Catawba River basin is
exceptional also in that it has a long reach upstream in
the inner Piedmont in which the river course is strongly
controlled by rock structure. This region is presumably

14 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

upstream and northwest of the area of superposition, so
that the main stream itself is structurally controlled.

The 11 smaller river basins north of the Cape Fear
River that originate on the Piedmont are considerably
narrower. They average about 8 percent narrower than
their length-law widths, and they resemble closely the
three rivers of the Coastal Plain that are also included in
the table. Presumably, these streams have patterns
similar to the normal patterns of consequent streams
whose basins have formed unrestricted by geologic or
tectonic factors. This assumption is consistent with the
concept of a late overlap of the outer Piedmont north of
the Cape Fear River by either marine or fluvial deposits.
As the drainage patterns have not been much affected
by the trends of the rocks in the Piedmont, I assume that
the overlap could probably have taken place in late Ter-
tiary time, as studies in the Washington, DC, area sug-
gest (Barton, 1951). The area of overlap by upper Ter-
tiary Coastal Plain sediments corresponds to a steep
gradient of the submerged and buried crystalline base-
ment and is associated with the deep Salisbury embay-
ment (fig. 3).

The analysis of basin width in the Piedmont supports
the idea that the large southern drainage basins have
been affected by geologic restrictions imposed during a
long period of erosion occupying most of the Cenozoic.
The shorter streams of the outer Piedmont to the north
are probably much younger, and their original pattern,
developed on a sedimentary cover, has not been substan-
tially altered.

REGIONAL PHYSIOGRAPHIC FEATURES RELATED
TO THE PINE MOUNTAIN BELT

The Chattahoochee and Flint Rivers (20 and 21, fig. 5)
have anomalous profiles in the outer part of the Pied—
mont. Both rivers descend abruptly about 90 m between
the northern part of the Pine Mountain belt and the Fall
Zone, a distance of about 40 km. The average gradients
in these rocks are 0.0024 for the Chattahoochee River
and 0.0021 for the Flint River, greater than the gra-
dients upstream by factors of 5.3 and 4.2, respectively.
Rock control is a possible explanation, as the two
streams must cross several outcropping layers of resist-
ant quartz rock within these steep reaches that would in-
crease the caliber and resistance of the bedload. On the
other hand, the Chattahoochee River is much larger
than the Flint, having a drainage area of 11,450 km? at
the Fall Zone, compared with 5,240 km2 for the Flint.
One would expect that the gradient of the Chatta-
hoochee River would be gentler than than of the Flint if
both rivers had adjusted their profiles. Upstream for a
distance of 96 km from the Pine Mountain Belt, the
average gradient of the Chattahoochee is 0.00045,

 

whereas that of the Flint is only slightly steeper,
0.00050.

When the profiles of these two rivers are compared
with those of the Ocmulgee and Oconee, the neighbbring
streams to the east, one sees little similarity. The Oc-
mulgee and Oconee show a concave profile as they de-
scend the Piedmont, whereas the Chattahoochee and
Flint Rivers have a sharp change in gradient just above
the Pine Mountain belt that makes their overall gradient
convex. The low gradient of the Flint River upstream is
reflected in the low relief of the region north of the Pine
Mountain belt (fig. 4), an indication that the Pine Moun-
tain belt has acted as a barrier to slow the erosion of the
river basin upstream relative to the bordering areas.
The steep reaches in and below the Pine Mountain belt
may be related entirely to quartzite in the bedload of
these rivers. However, one can speculate that the initial
convex profiles of these two rivers were formed while
faulting was in progress, that both streams were then
affected to the same degree, and that time has not been
sufficient since faulting ceased for the profiles to
become adjusted in relation to the relative sizes of the
rivers. On the Chattahoochee River, the sharp break in
slope that presumably should locate the most recent
fault is below the Goat Rock fault at or close to the inner
edge of the Coastal Plain. On the Flint River, the sharp
break in slope is within the Piedmont at or near the Goat
Rock fault.

DIFFERENTIAL MOVEMENT ALONG THE FALL ZONE

South of the Roanoke River, most of the streams of
the Piedmont enter the Coastal Plain more or less at
grade, that is, without a sharp irregularity in the profile
(fig. 8). The Chattahoochee and Flint Rivers are excep—
tions, as explained above. The Savannah River has a
short steep reach at the Fall Zone in Augusta, Ga., in
which the river drops 11.7 m in 5.5 km. This slope is
about 0.002, compared with 0.00056 in a more typical
16-km reach upstream (data from U.S. Army Corps of
Engineers). The steep reach crosses the Belair fault
believed to be active in the Tertiary (Prowell and O’Con-
nor, 1978).

The Cape Fear River (fig. 8) has a prominent steep
reach at the Fall Zone. The steep reach is just above
North Carolina Highway 217 near the town of Erwin.
The river drops 12 m in 8 km, a slope of 0.0015, com-
pared with 0.0003 in a 24-km reach upstream (data from
U.S. Army Corps of Engineers). No late fault has been
reported in this area. Other more pronounced steep
reaches occur at the Fall Zone on the Roanoke, Ap-
pomattox, James, and Rappahannock Rivers. On the
James River, a fault zone of Tertiary age is reported just
downstream from the head of tidewater at Hopewell,
Va. (Dischinger, 1979). The Rappahannock River

 

NORTHEASTERN HIGHLANDS l5

 

 

 

 

  

 

 

RAPPAHANNOCK RIVER
100 —
_ F2
0 /
_ JAMES RIVER
1oo -
_ z
o /F
200 APPOMATTOX RIVER
100 —
- /F2
0
- ROANOKE RIVER
200 —
m 100—
I _
Lu /FZ
I— o
“" 200
E _ DEEP-CAPE FEAR RIVER SYSTEM
E 100—
U-l " /F2
0
2 °-
— NEUSE RIVER
I; zoo—
<
100-
.. F2
0

 

YADKlN—PEEDEE RIVER SYSTEM

  

300 -

 

  

 

300 — SALUDA-CONGAREE RIVER SYSTEM

200

100

 

 

200 - SAVANNAH RIVER

    

100—

 

OCONEE RIVER
300

200

100

 

 

OCMULGEE RIVER
300

200

100

 

 

300 — FLINT RIVER

    
 
 
 
 

200 Apalachicola River

  
 

 

 

 

 

 

 

 

/FZ
200 : 100 _
100 - 0 _
0 ' _ CHATTAHOOCHEE RIVER
300— CATAWBA-WATEREE-SANTEE RIVER SYSTEM 300 ~
_ _ Apalachicola River
200 - 200 -
100 - 100 :
0 I I I | L I I I 0 r I T I L I I I I
0 200 400 600 800 0 200 400 600 800

DISTANCE FROM SOURCE, IN KILOMETERS

DISTANCE FROM SOURCE, IN KILOMETERS

FIGURE 8. — Profiles of 14 streams in the Piedmont Lowlands. Only those parts of the profiles within the Piedmont are shown. Distances along the
profiles are measured from the stream heads. Fz, Fall Zone.

crosses a fault zone of Tertiary age at Fredericksburg,
Va., in the Fall Zone (Mixon and N ewell, 1977).

NORTHEASTERN HIGHLANDS

The Northeastern Highlands is an area of substantial-
ly higher relief than the Piedmont Lowlands. It extends
from northern Virginia to the Delaware River and is
bounded on the southeast by the Fall Zone and on the
west and north by the Blue Ridge Mountains and the Ap-

palachian Valley. The highlands include a strip of inter-
connected Triassic and Jurassic basins on the northwest
margin, in general lower than the highlands proper. The
area southeast of the Triassic and Jurassic basins in
Maryland and southeastern Pennsylvania is com—
paratively high. It was recognized by Campbell (1929,
1933) as distinctly different in relief from the surround-
ing Piedmont; he referred to it as the Westminster an-
ticline, an upwarped peneplain. Cleaves and Costa
(1979) have cited evidence indicating late geologic uplift

16 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL. UPLIFT, PIEDMONT AND BLUE RIDGE

of the Maryland Piedmont. I described the area briefly
in an earlier paper (Hack, 1980) as an example of an area
uplifted in late geologic time relative to other parts of
the Piedmont. The evidence includes (1) higher relief, (2)
a steep gradient in the basement beneath the Coastal
Plain, (3) differences in the gradients of sheets of river
gravel of Cretaceous and Miocene age, and (4) the Staf-
ford fault zone at the Coastal Plain boundary described
by Mixon and Newell (1977). In this paper, additional
evidence based on stream profiles and valley forms is
discussed.

The relief of the Northeastern Highlands is in itself a
distinctive feature suggesting an uplift rate higher than
that in other Piedmont areas. In the: center of the
crystalline rock area, the relief exceeds 125': m within
many areas of 100 kmz, whereas, in much of the Pied-
mont, relief is less than 90 m in areas of that size. It is
also of interest that, although the Triassic and Jurassic
basins included in this area (the Gettysburg and Newark
basins) contain large areas of low relief, generally of
shaly rocks, the more resistant rocks within them

generally have relief comparable with that in the edger» ‘

cent highlands. For example, in the Frenchtown-
quadrangle of the Newark basin (Drake and others,
1961) a 100-er12 area of fanglomerates has a total relief
of 220 m, equivalent to the relief in the crystalline rocks
of the adjacent Reading Prong to the north. The
Reading Prong is a part of the New England
physiographic province, not included in this report.
Areas of similar and even higher relief of as much as 270
In occur on diabase sills and quartzose fanglomerates in
the Gettysburg basin. None of the Triassic and Jurassic
basins in the Piedmont Lowlands has comparably high
relief, indicating that the areas of Triassic and Jurassic
rock within the Northeastern Highlands probably are in
a different tectonic setting and are within the uplift
area.

Stream profiles and valley forms, not previously
analyzed, are useful features for analysis. As the
topographic map (fig. 3) shows, the highlands are cut by
two deep and narrow valleys, those of the Potomac and
the Susquehanna Rivers. These long streams originate
in the Appalachian Plateaus. They are referred to here
as exterior streams because they originate outside the
Piedmont and Blue Ridge. No similar deep valleys exist
in the Piedmont Lowlands. Figure 9 shows the longi-
tudinal profiles of these rivers and other rivers of the
subprovince. Distances on the horizontal scale are
measured from the source of the longest tributary, but,
except for the Patapsco River, the profiles include only
the lower reaches. Consider first the Susquehanna
River. This stream is the largest crossing the Piedmont
in terms of drainage area and discharge, yet its profile is
strongly convex. The convexity begins near Selinsgrove,

 

30° DELAWARE RIVER

200;

100

llllll

all Zone
0 /F

 

SCHUYLKILL RIVER

 

 

w _
5 100— \\
E 0 ' /Fall Zone
_ SUSQUEHANNA RIVER

Z 100 - \
g 0 /Fall che
g PATAPSCO RIVER
5 200
<

100 —

/Fal| Zone

 

POTOMAC RIVER

100
“ Fall Zone
0 I I I I I I I I I

0 200 400 600 800
DISTANCE FROM SOURCE, IN KILOMETERS

 

 

 

1

FIGURE 9. —Profiles of five streams in the Northeastern Highlands. In
all the streams except the Patapsco River, the upper reaches are not
shown, but distances are measured from the stream heads (along the
longest tributary). All except the Patapsco head in the Valley and
Ridge or Appalachian Plateaus provinces.

Pa., 7% km upstream from the Piedmont. The channel
slope continues to steepen through the Triassic and
Jurassic belt and in the crystalline rocks. Although it is
not shown in the figure, the steep slope continues below
sea level, and at the river mouth, the drowned
Pleistocene channel is bordered by steep walls now
covered by sediment. Its altitude at the point where it
enters Chesapeake Bay is -45 m.

The Potomac River is similar, but the convexity is not
so regular. The river has a steep concave profile in the
Valley and Ridge province. At Harpers Ferry, W.Va., it
drops more steeply as it passes through the Blue Ridge.
In the Triassic and Jurassic rocks, it again assumes a
concave profile. However, 29 km above tidewater, it
enters the crystalline rock area and plunges through a
deep gorge to the Coastal Plain. The average slope in
the gorge is 0.009. Below the head of tidewater, the
Pleistocene channel is buried in sediments, but 19 km
downstream at Alexandria, Va., borings show the
bedrock channel at -30 m. This is an average slope of
about 0.0016, much less than the slope in the Piedmont.

The Delaware River (fig. 9) has a more gentle slope in
the lower reaches, averaging about 0.0006 in the Pied-
mont. At the Fall Zone at Trenton, N.J., the river turns
sharply south. Sixteen kilometers downstream, the
elevation of the bottom of the buried Pleistocene chan-

 

I

r ‘1'
7

y

r‘ “
?
* r

r
?

NORTHEASTERN HIGHLANDS 17

nel is — 15 m. The average slope in this reach is 0.00095,
somewhat steeper than the slope upstream. The op-
posite situation occurs in the Potomac, where the buried
channel is much gentler than the subaerial channel in
the Piedmont. The Schuylkill River (fig. 9) has a concave
profile in the Piedmont, but it is notched at the Fall
Zone.

The Patapsco River is the only one of the group shown
in figure 9 that originates within the highlands. As can
be clearly seen, it steepens as it approaches the Coastal
Plain. Note that its profile is similar to that of the Rap-
pahannock River (fig. 8). These data seem to indicate

   

 

 

that the Susquehanna, Patapsco, Potomac, and Rap
pahannock Rivers were all affected by differential move-
ment at the Fall Zone in late geologic time when the
Piedmont was uplifted with respect to the Coastal Plain.
The characteristics of the Schuylkill and Delaware River
profiles are different. These rivers may be northeast of
the zone of latest disturbance.

The conclusion that uplift took place in late geologic
time within the Northeastern Highlands can be further
supported by evidence of an evolutionary development
of the topographic forms. The Patapsco River basin (fig.
10) is in the central part of the highlands on the east side

l
BALTIMORE !
l

FIGURE 10. — Map of the Patapsco River basin, Maryland. Numbers in circles identify first-order tributaries shown in figure 14. Numbers in
rectangles are gradient index values, in meters, of the reaches shown. The gradient index is essentially the slope of the logarithmic
profile along a river reach and is a crude index of stream power. It is measured by obtaining the ratio of the fall within the reach, in
meters, to the difference between the natural logarithms of the river lengths at each end of the reach, as measured from the source of

the stream.

18 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

of the high divide, known as Parrs Ridge, that separates
the drainage toward Chesapeake Bay from that toward
the Triassic and Jurassic basins. In the lower eastern
part of the basin, slopes are steeper and power is
greater than they are in streams near the divide, as
might be expected in a landscape that was uplifted
relative to the Coastal Plain in late geologic time. The
process that has produced these features resembles the
classic concept of landscape rejuvenation by headwater
migration of steepened slopes.

Consider first the profile of the Patapsco River itself.
The profile is plotted on a logarithmic scale (fig. 11),
starting at the stream head farthest from the river
mouth, on this river, the Cranberry Branch north of
Westminster, Md. Streams having normal profiles
generally plot on a straight line on such a graph. A nor-
mal stream in this context is one in which discharge and
slope are inversely related so as to conserve the power
of the stream to transport a load of a certain caliber at
approximately the same rate along its course. Figure 12
is the logarithmic profile of the Appomattox River as an
example for comparison. The Appomattox River crosses
a part of the Piedmont covered by thick saprolite and
weak sediments of a Triassic and Jurassic basin. It has a
smooth profile and flows through banks of silty material.

 

The river bars are composed of sand and fine gravel. Ex-
cept for the first 4 km below the source and the 30 km
above the Fall Zone, the Appomattox River has a
remarkably smooth profile that follows fairly closely the
same parabolic curve.

In contrast, the Patapsco River (fig. 11) has a
logarithmic profile that becomes increasingly steep as it
approaches the river mouth at Baltimore Harbor. In
general, this profile seems not to be related to the rocks
in the basin. Locally, however, changes in slope in places
correspond to geologic contacts. Such a change is most
obvious in the lower part of the river where it cuts
through the gabbro of the Baltimore Complex.

The bed material in the Patapsco River basin becomes
coarser in a downstream direction. Estimates of the me-
dian size of bed material were made at several localities,
and data for Gillis Falls (fig. 10), a headwater tributary
of the South Branch, have been published (Hack, 1957).
Medianlsize of bed material in this tributary increases
from 7 mm at a distance of 1 km from the source to 40
mm at a distance of 13 km. At Hoods Mill on the South
Branch, the median size has increased to 80 mm, and at
Hollofield on the main stem of the Patapsco River, the
median size is 170 mm. This locality is in the Baltimore
Gneiss. At extreme low water, the riverbed resembles a

 

Cranberry Branch

 

200

PATAPSCO RIVER '

ALTITUDE, IN M ETERS

 

l I lllllll

West Branch

 

North Branch

       

 
   

\\ \ Liberty Reservoir

Baltimore
Complex

 

l I l |lll|

 

(in ‘0

DISTANCE FROM SOURCE, IN KILOMETERS

FIGURE 11. —Logarithmic profile of the Patapsco River, Maryland.

NORTHEASTERN HIGHLANDS 19

 

200

APPOMATTOX RIVER

100‘

ALTITUDE, IN METERS

 

    
        

 

 

 

 

 

l llllllll l

 

1.0 10

100

DISTANCE FROM SOURCE, IN KILOMETERS

FIGURE 12. -Logarithmic profile of the Appomattox River, Virginia.

boulder field. The channel at this locality is 50 m wide
and 1.6 m in average depth. These data are dramatic
evidence that the power of this river system increases
substantially toward the mouth.

The valley forms also change downstream; the side
slopes become steeper and higher (fig. 13), and tributary
streams increase in slope in the lower part of the basin
(fig. 14). The streams shown in figure 14 were selected
because they were approximately the same length from
source to junction with the master stream.

Hypsometric analysis of quadrangles in this area
shows that the topography is plateaulike in the sense
that the largest part is upland. For example, a 100-km2
area of the Winfield Quadrangle that lies on the
drainage divide at the crest of Parrs Ridge has an
elevation—relief ratio (Pike and Wilson, 1971) of 0.6. In
other words, the mean relief is 60 percent of the total
relief. A normal ridge-and-ravine landscape charac-
teristic of most of the Piedmont where isolated high hills
are lacking would have a ratio of approximately 0.5.
Hammond (1964), in his study of the landforms of the
United States, classified the entire crystalline rock area
of the Northeastern Highlands as tablelands having
moderate relief. In this terrain category of Hammond’s,

 

50 to 80 percent of the area is gently sloping, local relief
is 300 to 500 ft (within areas 6 mi across), and 50 to 75
percent of the gentle slope is on the upland.

The topographic features of the crystalline rock area
in the Northeastern Highlands seem to point to uplift in
late geologic time as the explanation for the systematic
change toward the interior of the area and away from
the Coastal Plain. They suggest uplift at the Coastal
Plain boundary or possibly upward tilting toward the
northwest. The limestone valley and Triassic and
Jurassic basins behind the crystalline rock, on the other
hand, have large areas at lower altitudes, particularly in
the southern part of the Gettysburg basin, probably
because this area is underlain almost entirely by
nonresistant rocks such as limestone, siltstone, and
shale. Where crystalline rocks such as diabase and re-
sistant sedimentary rock such as fanglomerate do occur,
relief is equivalent to that of the highlands on the
southeast.

Probably the uplifted area extends far to the north-
west, and the Northeastern Highlands are simply the
southeastern margin of the Appalachian Mountain
chain. They are not an anticlinal area, as Campbell
(1929) believed, but a monoclinal area.

PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

36m

122 m
GILLIS FALLS ABOUT 3.4 KILOMETERS FROM SOURCE

VI

I91 m
PATAPSCO RIVER NEAR HOLLOFIELD, MD.

FIGURE 13. —Transverse profiles of two valleys in the Patapsco River
basin.

SOUTHWESTERN HIGHLANDS

The Southwestern Highlands is similar in some ways
to the Northeastern Highlands. Both areas are under-
lain in large part by polydeformed metasedimentary and
volcanic rocks of Williams’ (1978) map units 8 and 2. In
the Southwestern Highlands, these rocks are bordered
on the northwest by a band of elastic rocks of lower
metamorphic grade, many of them fine grained and in-
terbedded with quartzite. As in the Northeastern
Highlands, the relief is moderate, although in places, it
is much higher than that in the Piedmont Lowlands. The
highest parts of the area are along the northwest
margin, where Cheaha Mountain (730 m) and Talladega
Mountain form a curving ridge held up by quartzite in
the phyllitic rocks of the Talladega Slate. Figure 15 is an
averaged topographic map showing the setting of the
Southwestern Highlands in relation to the Blue Ridge
and Piedmont. Figure 16 shows generalized topography
in relation to rock types.

The physiography of the area has been described as a
series of plateaus. The northern, or Georgia, part was
divided into three areas by La Forge and others (1925).

 

— 260

®—24o

‘220

 

260 ~

240 *

220 — ©

200 '—

 

“ 260
‘. 240
® ~ 220

4200

j180

 

260 7

240 T

220 I-

...I @

180L - 180
<f160

ALTITUDE, IN METERS
ALTITUDE, IN METERS

 

 

160
140
120

100 ®

80

 

500 1000 1500

60L
0
l I I I I I | I I I I I I I I I I I I

DISTANCE FROM TRUNK STREAM, IN METERS

FIGURE 14.—Profiles of some first-order tributaries in the Patapsco
River basin. For locations, see figure 10.

p.
l

 

SOUTHWE STERN HIGHLANDS 21

p r
Qfifl

 

Atlanta

50 100 KILOMETERS
l J

 

CONTOUR INTERVAL
30 m below 450
150 m above 450

FIGURE 15. — Averaged topographic map of the Southwestern Highlands and surrounding area (shaded). Altitudes are based on the mean altitude
of points within 3-minute rectangles.

The Dahlonega Plateau at the north is a dissected
upland. The lower Atlanta Plateau extends westward in-
to the area, and the Tallapoosa Plateau constitutes the
higher area to the southwest. The Alabama section was
called the Ashland Plateau by Johnston (1930). These
areas are not true plateaus, as the landscape forms are
little different from those of the rest of the Piedmont but
are higher in altitude and relief and have steeper slopes.

A prominent feature of the area is an escarpment
more than 30 m high that separates the Tallapoosa
Plateau or upland from the Atlanta Plateau to the north-

 

east (A, fig. 16). This escarpment corresponds to the
drainage divide separating the headwaters of the
Tallapoosa River from small tributaries of the Chat—
tahoochee River on the east and from the Etowah River
on the northeast. The Etowah River drains into the
Coosa system in the limestone and shale valley to the
northwest at a much lower level. The Tallapoosa River
system is higher than the other two systems, probably
because of rock control. Steep reaches occur in the
volcaniclastic rocks as well as in the feldspathic rocks
above the junction with the Little Tallapoosa River.

22

34°

PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

87° 35°

85° 84°

 

 

 

 

  

‘2
‘8

§

 

 

 

 

a} Anmsto
<15“
00 0;?
A M“
2/

cm
5/5.?

 

   
   

100 KILOMETERS
CONTOUR INTERVAL 122 m

 

 

34°

ms Metasedimentary rocks
v Volcanic rocks

9 Gtaphitlc schist and phyllite

98 Garnet and graphific schist

f Feldspathlc schist and gneiss

 
 
 
 
 

C‘ EXPLANATION
04
Sr ‘v/L» -
4L PLAl N Areas hlgher than 488 m
87° 86° 8 ° 34°
EXPLANATION

m9

 

33°

 

 

  

 

 

 

 

 

 

Granitic rocks
by Biotite gnelss

 

mg Metagraywacke and quartzite
bv Brevard zone

100 KILOMETERS

 

 

 

FIGURE 16.—Simplified topographic and geologic maps of the Southwestern Highlands. Geologic data from Neathery and others

A

44

SOUTHWESTERN HIGHLANDS 23

Many local features of the topography of this region
appear to be related to rock control. On the other hand,
the topography and relief on all the rock types become
lower as the Coastal Plain is approached on the
southwest. The general altitude of the landscape also
declines to the southeast, and relief becomes distinctly
lower near the Brevard zone (fig. 15).

The profiles of the Coosa and Tallapoosa Rivers (fig.
17) suggest substantial uplift of the area relative to the
Coastal Plain in late geologic time. The Coosa River
passes through the area in a broad low valley, but its
gradient is steep (0.00095) for a river of such large size.
As river profiles prepared by the US Army Corps of
Engineers show, in the lower 60 km of the Piedmont,
the river was intrenched as deeply as 130 m before im-
poundment by a series of dams. The area of the drainage
basin of this river is about 2,600 kmz. In terms of
average discharge, it is the second largest river crossing
the Piedmont, exceeded only by the Susquehanna River.
Figure 17 shows the extreme convexity of the profile,
beginning in the Appalachian Valley.

The Tallapoosa River, like the Patapsco River in the
Northeastern Highlands, is not an exterior stream. Its
source lies within the crystalline rocks. Its profile is
much like that of the Patapsco, and its power increases
downstream toward the Fall Zone (fig. 18). A 32-km
reach immediately above the Fall Zone has an average
slope of 1.9 m/km or 0.0019. The gradient index (Hack,
1973a) is 490 gradient m (1,600 gradient ft), equivalent
to that of many powerful streams in the Blue Ridge,
such as the French Broad River below Asheville, NC.
Unfortunately, this steep lower reach of the Tallapoosa
River is now drowned by the waters of Martin Lake so
that it cannot be seen. However, at Martin Dam, the
river was intrenched into the upland approximately 84
m before impoundment (US. Army Corps of Engineers

 

   
 
 

 

 

 

 

400
TALLAPOOSA RIVER

2 30° ' F H z
E a one
m 200 - Joins Coosa
g _
Z 100 —
u; o '
0 Enter Piedmont
a 200 _ COOSA RIVER Fall Zone
; Joins
:(J 100 _ Tallapoosj

)—

0 1 1 I ' l l '
o 200 400 600 800

DISTANCE FROM SOURCE, |N KILOMETERS

FIGURE 17. —Longitudinal profiles of the Tallapoosa and Coosa
Rivers. Distances are measured from the heads of the streams.

 

data). Upstream from the lake, the average gradient in a
long reach is 165 gradient meters (870 gradient ft). The
river in this reach has an average discharge of 225 cms
(2,426 cfs) and is a powerful river, though it flows
through a topography of low relief.

The general configuration of the topography of the
Southwestern Highlands of Alabama indicates a sharp
uplift at the Fall Zone and a more gradual monoclinal
rise from the Brevard zone northwest toward the Ap-
palachian Valley (fig. 15). The topography in the Ap-
palachian Valley is strongly controlled by the bedrock.
Nevertheless, the resistant formations underlie plateaus
and mountains 300 to 500 m in altitude, comparable with
the general level of the Alabama part of the highlands,
though not with the highest ridges such as Cheaha
Mountain (730 m). In the extreme northeastern part of
the highlands, the monoclinal rise to the Blue Ridge
province is steeper.

FOOTHILL ZONE

The Foothill zone, as defined here, extends from the
Potomac River southwestward through the inner Pied-
mont to near the South Carolina—North Carolina bound-
ary. Its boundary on the northwest is defined by the
southeast margin of the Blue Ridge physiographic prov-
ince (Fenneman and Johnson, 1946). The southeastern
boundary is defined approximately by the outer limit of
hilly areas and isolated hills within the inner Piedmont
and by a slight increase in slope (fig. 3).

The Foothill zone can be divided into three areas. The
part north of the Roanoke River consists mostly of small
ranges of hills and mountains underlain by resistant
rocks and separated by extensive lowlands having much
lower but varied relief. The Catoctin Formation, con-
sisting of mafic lava flows, forms many ranges of high
hills and must be one of the most resistant rock types,
though the relief varies within local areas. The
metasedimentary rocks of the Evington Group exposed
in the long northwesterly reach of the James River con-
tain resistant beds that underlie ranges of hills. These
rocks are commonly quartzites of the Mount Athos For-
mation and of the Candler Formation, which also con-
tains phyllite, muscovite schist, and locally marble
(Espenshade, 1954, p. 15). Ranges of steep but low hills
occur on the Lynchburg Formation; in Madison County,
Va., hills are apparently on beds of conglomerate ex—
posed in a syncline (Allen, 1963, map cross section
A-A’). A broad belt of coarse-grained granitoid rocks of
Proterozoic Y age forms the core of the Blue Ridge anti-
clinorium just southeast of the Blue Ridge Mountains.
The rocks have considerable variety; some are sheared

24 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

 

 

    
   

300—
" Leave Talladega belt
(I)
n: T
Lu
I;
E ‘ TALLAPOOSA RIVER
E 200—
u; n
D
I) _
|:
’—
_, _
<
100—
_ Join Coosa River
0 I I I I I I l I | I I | I I I
1.0 10 100

 

 

DISTANCE FROM SOURCE, IN KILOMETERS

FIGURE 18.—Logarithmic profile of the Tallapoosa River.

and fragmented by cataclasis. Within this belt are many
ranges of hills. The nature of the rock in places locally
controls the topography. For example, Big Cobbler
Mountain in the Orlean Quadrangle, Virginia, is within
the local outcrop area of the magnetite-rich granite of
the Robertson River Formation of Allen (1963) (James
Clarke, oral commun., 1978). Thus, in this area, the ef-
fect of rock control is evident, and no proof exists that
the local relief is related to uplift or tilting of the area.
Nevertheless, the height of the hills is generally greater
to the northwest close to the Blue Ridge Mountains.

Between the Roanoke River and the Yadkin River, the
Foothill zone is more clearly defined and takes the form
of an upland that has a relief uniformly greater than
that of the Piedmont to the east. In much of this area,
the zone is defined by two geologic provinces, the Smith
River allochthon and the Sauratown Mountain window.
Locally, quartzite forms mountains higher than their
surroundings (for example, see Espenshade and others,
1975), but in general, the relief of the Foothill zone is
markedly greater than that of the Piedmont Lowlands
to the east.

 

Between the Yadkin River and the South Mountains,
where the Foothill zone ends, the boundary is drawn
across the lowlands of the inner Piedmont, that is, the
lowlands between the Brevard zone and the Kings
Mountain belt. The boundary shown on figure 1 follows
a faint lineament that can be seen clearly on topographic
maps that show contours only (US Geological Survey
maps of Winston-Salem, Charlotte, and Knoxville
Quadrangles, scale 1:250,000). Whether this boundary
itself has real meaning is uncertain, but it does serve as
well as or better than an arbitrary boundary because
most hilly areas of the inner Piedmont are northeast of
the line. Figure 6 shows the contrast in relief. In this
part of the Foothill zone, as is generally true, rock type
does control the topography to some extent, although to
a lesser degree than it does in the Virginia part of the
Foothill zone. For example, the Brushy Mountains and
South Mountains are both areas that have many peaks
with altitudes of 750 to 900 m. These mountains are
underlain by the more resistant rocks of the region, as
shown on the geologic map of the Charlotte 2°
Quadrangle (Goldsmith and others, 1978). The rocks in-

 

FOOTHILL ZONE 25

clude (1) a biotite gneiss unit containing hornblende
gneiss, garnetiferous gneiss, and felsic volcanic rocks;
(2) an augen gneiss, known as the Henderson Gneiss; (3)
muscovite and sillimanite schist and minor quartzite and
sandstone; and (4) a biotite gneiss unit that contains
layered quartzofeldspathic gneiss, metasandstone, and
siltstone.

Although the conspicuously high areas do contain
what appear to be resistant units, all these resistant
units also occur in low relief areas, where a variety of
other rock units include granite gneiss, migmatite, horn-
blende gneiss with amphibolite, and schist. Even the
sillimanite-rich rocks are not restricted to the moun-
tains. This relation between topographic forms and rock
types suggests that both tectonic forces and rock control
have operated to produce the present forms, tectonic
processes playing a major role. Figure 7 shows that the
local contrasts in relief are impressive. The South Moun—
tains rise 450 m above the Morganton Basin. The local
relief within the mountains averages about 100 m, com-
pared with less than 30 m in the basin. Many of these
hills and small mountain ranges may be residuals formed
by slope retreat as described by Kesel (1974).

Four large rivers cross the outer border of the Foothill
zone. Three of them have a reach that has a steep gra-
dient at the crossing. However, the steep reach on each
river corresponds to what is probably a resistant rock
type. The gradient index or logarithmic slope (see fig. 10
for definition) can be used to compare the profiles of
rivers having different discharges or lengths. Gradient
index values greater than 300 gradient m are charac-
teristic of mountain streams, whereas values in the
Piedmont are commonly less than 150 gradient m.

The James River has a gradient index of 365 gradient
m in a reach 20 km long in the area crossed by the outer
boundary of the Foothill zone. Downstream and
upstream values average about 150 gradient m. The
steep reach occurs in rocks of the Evington Group and in
metavolcanic rocks of the James River syncline, both of
which are probably resistant. The Roanoke River has a
steep reach 15 km long just below the Danville Triassic
and Jurassic basin, which forms the boundary of the
Foothill zone. Values range from 275 to 560 gradient m.
The rock units are the “Melrose granite” and “Shelton
granite gneiss.” Immediately below the steep reach,
values drop to 75 gradient m.

In North Carolina, the Yadkin River crosses the
boundary of the Foothill zone without a break in its pro-
file. Twenty kilometers upstream is a short reach of 290
gradient m in a resistant plutonic rock, the Cranberry
Gneiss (Espenshade and others, 1975). No reason exists
to believe that the Foothill zone boundary has any effect
on the profile of this river. The fourth river, the
Catawba, crosses the Foothill zone boundary at the

 

Lookout Shoals Reservoir site. The gradient index of the
channel here is 300 gradient m, much higher than values
for some distance above or below. The rock is mapped as
part of a large area of biotite schist, which would not be
expected to steepen the profile.

The evidence for an active uplift localized along the
boundary of the Foothill zone is not clearly supported by
the stream-profile data, at least on the large streams.
The Yadkin River shows no steep reach along the line.
The Roanoke has a steep reach a few miles downstream
at the border of a Triassic and Jurassic basin. The other
two rivers, the James and the Catawba, do have steep
reaches on the boundary, but the reach on the James
could be related to rock resistance rather than to a tee-
tonic break. These facts may mean either that the outer
boundary of the Foothill zone is a broad zone of uplift
that is not faulted at its border or that the large river
profiles have become adjusted through time to the
resistance of the rocks that they cross. Nevertheless, the
Foothill zone must be an area that has at some time been
tilted upward, and even though many of the hill areas
correspond to resistant rock, the hills are generally
higher near the Blue Ridge Mountains than they are far-
ther out in the Piedmont.

NORTHERN BLUE RIDGE MOUNTAINS

The term “Blue Ridge” has acquired several meanings.
As a geographic feature recognized by nongeologists, it
is simply the range that forms the southeastern border
of the Appalachian Mountains between northern
Georgia and Pennsylvania. South of Roanoke, it in-
cludes only the escarpment and mountains behind it that
rise directly above the Piedmont. This common defini-
tion excludes, for example, the Great Smoky Mountains
as well as some other ranges southeast of the Ap-
palachian Valley. Physiographers have long used the
term “Blue Ridge” as a name for a physiographic prov-
ince. The boundaries are shown on the map of the
“Physical Divisions of the United States,” published by
the U.S. Geological Survey (Fenneman and Johnson,
1946). The province includes all the high mountains
underlain by crystalline rocks from northwest Georgia
to Pennsylvania. Structural geologists now use the term
“Blue Ridge belt” to describe a geologic province (King,
1955, p. 358). In their definition, the Blue Ridge belt is
an anticlinorium cored typically by crystalline rocks of
Precambrian age and flanked by stratified rock of prob-
able Precambrian and Early Cambrian age. It includes
all the area northwest of the Brevard zone in Alabama
and Georgia. It thus includes the Southwestern
Highlands of this report, an area that is part of the Pied-
mont of the physiographers. The geologic definition
broadens the Blue Ridge in northern Virginia to include

26 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

all the rocks northwest of the Culpeper Triassic and
Jurassic basin, that is, the Blue Ridge-Catoctin Moun-
tain anticlinorium. However, from the north end of the
Brevard zone near Winston-Salem to Charlottesville,
Va., the southeastern boundary is not clearly defined.

In this paper, I have followed the usage of the
physiographers but have divided the Blue Ridge
physiographic province of Fenneman and Johnson into
two parts. The part north of the Roanoke River, which is
essentially a narrow range of high mountains, is re-
ferred to as the Northern Blue Ridge Mountains. The
area south of Roanoke is referred to as the Southern
Blue Ridge province. I have separated the two because I
believe that the origins of the topography are not the
same.

The Northern Blue Ridge Mountains are essentially
the frontal mountain range of the Valley and Ridge
province. They are underlain by a sequence of Precam-
brian and Cambrian rocks that form the northwest limb
of the Blue Ridge anticlinorium of the geologists (E spen-
shade, 1970). The sequence is composed of rocks that are
resistant to erosion.

Only three rock units support the crest of the range
along its entire length, a distance of 395 km (fig. 19): (1)
a complex assemblage of medium to coarse granitoid

p v
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THORNTON POTO RIVER
MANASSAS (9:32:
GAP

PEAKS OF
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GAP GAP

 

 

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E 600
E 200

 
 
  

 

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ROCKFISH SWIFT RUN

STONY CHESTER
MAN

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rocks, some of which contain large crystals of potassic
feldspar and include hypersthene granodiorite; (2) a
volcanic unit known as the Catoctin Formation; and (3) a
thick series of quartzite, arkose, and phyllite called the
Chilhowee Group that forms the base of the Paleozoic
sedimentary column exposed in the Appalachian Valley.

The hypersthene granodiorite unit extends most of the
length of the Northern Blue Ridge Mountains and is the
northwestern rock unit within the large area of Precam-
brian granitoid rocks that make up the core of the Blue
Ridge anticlinorium and that are referred to as the
Virginia Blue Ridge Complex of Brown (1958). The
hypersthene granodiorite is medium to coarse grained
and massive. It contains andesine or calcic oligoclase,
potassic feldspar, pyroxene (both hypersthene and
augite), biotite, and blue quartz. In many places, the
feldspar crystals are more than 8 cm long. In places, the
hypersthene granodiorite is altered to unakite, a rock
composed of feldspar, epidote, and blue quartz. The two
types grade into each other (Jonas, 1935). As the
hypersthene granodiorite is locally variable in texture
and composition, it has been commonly referred to as
the “Pedlar Formation” (Bloomer and Werner, 1955;
Allen, 1963). The “Pedlar Formation” is generally more
resistant than other rock units in Brown’s Virginia Blue

pCv pCv Ycr
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0 50 100 150 200 KILDMETEBS
I 1 I 1 J
EXPLANATION
“-5 KY, Em? E. 9&3 3m
Chilhowee Group Catoctin Formation Metarhyolite of the Swift Run Formation Virginia Blue Ridge
(clastic rocks) (metabasalt) Catoctin Formation (phyllite) Complex of Brown (1958)

(granitic rocks)

FIGURE 19. — Profile along the crest, and geologic map, of the Northern Blue Ridge from the Roanoke River to southern Pennsylvania. On the
geologic map, the crest of the range is shown by a line of dashes and dots.

44

A

NORTHERN BLUE RIDGE MOUNTAINS 27

Ridge Complex. This resistance can be demonstrated in
the area mapped by Bloomer and Werner (1955), which
includes 1,600 km2 in the Blue Ridge and Foothill zone
between the James River and Rockfish Gap (fig. 17).
Comparison of the mapped geologic units with
topographic maps shows that the “Pedlar Formation”
occupies the highest summits in both the Blue Ridge and
the Foothill zone. The highest peak in the area is the
Priest (1,236 m), about 11 km southeast of the range
crest. In the Foothill zone, these rocks underlie the
Tobacco Row Mountains that contain a peak 891 m in
altitude. Other rocks of Brown’s Virginia Blue Ridge
Complex in the Foothill zone occupy areas of lower
altitude and relief. The relief of the Lovingston Granite
Gneiss, for example, is somewhat less than that of the
Tobacco Row Mountains. The most extensive rock unit,
the “Marshall Formation,” is a medium-grained
granitoid rock that occupies valleys and low hilly areas.
The Roseland anorthosite body occupies a pedimentlike
area of about 17 km2 in the Foothill zone that has a total
relief of only 75 mm. Figure 19 shows that the “Pedlar
Formation” underlies some of the highest areas along
the Blue Ridge crest.

The Catoctin Formation, another unit having con-
siderable resistance to erosion, contains a variety of
rock types but consists mostly of massive green
metamorphosed basalt composed of albite, chlorite,
epidote, and actinolite. Agglomerate and other
pyroclastic rocks occur between the lava flows. The
thickness of the formation is variable and reaches 550 m
south of Thornton Gap (Reed, 1955). In this area, the
Catoctin lavas are massive and have many features that
are seen in recent lava flows, such as columnar jointing
and amygdules filled with secondary minerals. Seven
lava flows, having a total thickness of 460 m, can be
identified. The massive character of these rocks and the
gentle dips to the east account for the fact that this area
is the highest part of the Northern Blue Ridge Moun-
tains (fig. 19.) The Catoctin Formation is not
everywhere so massive and in places is sheared and
schistose. Nevertheless, as the differential topographic
relief in the Foothill zone shows, it is a resistant rock
and a ridge maker along much of its outcrop. For most
of the extent of the Blue Ridge anticlinorium north of
the Potomac River, the Catoctin Formation forms the
core, flanked on both sides by the Chilhowee Group. In
Pennsylvania, the principal rock type is metarhyolite. In
that region, however, the divide that forms the crest of
the range follows the Catoctin Formation for only a
short distance (fig. 19).

The Chilhowee Group is a clastic sequence of Pro-
terozoic Z and Early Cambrian age. The basal part is
generally a metamorphosed sandstone known in the
north as the Weverton Formation, and in the south as

 

the Unicoi Formation. In some places, a thin red slate
(former Loudoun Formation, now abandoned and
replaced by Catoctin) forms the base of the group. The
Weverton or Unicoi, however, is 150 to 500 m thick and
is mostly arkose but also contains shale and quartzite. It
is overlain by the Harpers Formation (Hampton Forma-
tion in the south), a sequence of metamorphosed shale
and siltstone and minor sandstone and quartzite;
thickness ranges from 300 to 600 m. The topmost forma-
tion in the group is the Antietam Formation (Erwin For-
mation in the south), composed of siliceous sandstone
and vitreous quartzite, 125 to 250 m thick. These forma-
tions generally dip steeply to the northwest and form
the northwest flank of the Northern Blue Ridge Moun-
tains. The Weverton Formation underlies the crest of
the range at many places for about 45 percent of the
crest length, a combined distance of 176 km. Over most
of this distance the divide is lower than it is on the other
rocks. The entire Chilhowee Group forms a wide belt of
mountains on the northwest margin of the Blue Ridge
that is generally lower in altitude than the crest of the
range and that is also lower than many hills southeast of
the crest. Thus, rocks of the Chilhowee Group, including
the Weverton Formation, are probably somewhat less
resistant than those of the Catoctin and “Pedlar” Forma-
tions.

Rock control does not explain the entire profile of the
range. Fiqure 19 shows that from Chester Gap to the
northern end in Pennsylvania, the range is lower than it
is to the south. North of the Potomac River, the range
does not exceed 620 m in altitude, and its average
altitude is little more than half that of the southeastern
part. On the other hand, the outcrop of resistant rocks in
places is as wide or wider than it is to the southeast. This
difference is perhaps best explained as the result of tee-
tonic causes.

The Northern Blue Ridge Mountains appear to be part
of an erosional system in an uplifted area that includes
the mountain ranges of the Valley and Ridge province
that rise to approximately the same altitudes. The forms
and relief in the Valley and Ridge province are con-

trolled largely by rock resistance and width of outcrop.

The decline in altitude north of the Potomac River may
be due simply to a lesser vertical component of uplift in
that region either currently or at some time in the past.

SOUTHERN BLUE RIDGE PROVINCE

The Southern Blue Ridge province extends from the
Roanoke River southwestward into north Georgia,
where the mountain chain on the crystalline rocks is in-
terrupted by lowlands along the lower Etowah River. As
defined here (fig. 1), the Blue Ridge province ends at
Pine Log Mountain north of Cartersville, Ga. The north-

28 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

west border of the Southern Blue Ridge province follows
the outcrop of the uppermost quartzite beds in the
Chilhowee Group of Early Cambrian age. The southeast
border is an erosional escarpment related only locally to
the underlying rocks. From the Georgia~South Carolina
border northward, the crest of the escarpment forms
the drainage divide between streams draining to the
Atlantic Ocean and those draining to the Gulf of Mexico.
Viewed as a whole, the escarpment cuts diagonally
across the entire Blue Ridge anticlinorium.

The Southern Blue Ridge province can be divided into
several areas on the basis of topography. Some of these
areas correspond to distinctive geologic areas, but some
do not (fig. 20). Rock control, however, is responsible for
many landscape features.

CHILHOWEE—WALDEN CREEK BELT

The landforms on the northwestern margin of the
province are closely related to the geologic pattern.
They form a zone of elongate mountain ridges and
valleys 5 to 50 km wide, underlain by sedimentary and
metasedimentary rocks of low metamorphic grade. Con-
trasts in rock resistance are great, as the rocks range in
composition from carbonate and shale to massive
quartzite. Long, steep-sided ridges separated by parallel
valleys are characteristic of the belt. Northeast of the
French Broad River, the quartzite of the Chilhowee
Group forms most of the ranges, whereas the valleys are
underlain by infolded or faulted carbonate and shale of
Cambrian age typical of the Appalachian Valley, such as
the Shady and Honaker Dolomites and the Rome For-
mation. Southwest of the French Broad River,
metasedimentary rocks of Proterozoic Z age, referred to
as the Walden Creek Group, are involved, forming a belt
of intermediate relief between the Chilhowee ridges and

35° 84° 35°

  

34°

 

 

 
 

. . cH
. K N ILHOWEE WALDEN
_ ,::E..MD Y fawning {31‘ .

  

the mountain highlands (Swingle and others, 1966;
Hadley and Nelson, 1971).

The mountains in the Chilhowee—Walden Creek belt
are not as high as many of the peaks of the mountain
highlands. Nevertheless, some ridges underlain by
quartzite of the Chilhowee Group exceed 1,200 m in
altitude, especially in the central part of the area. Local
relief is also high, in places more than 750 In. The many
longitudinal valleys, especially where they are underlain
by carbonate rocks, are bordered by large fans and
floored by residuum.

MOUNTAIN HIGHLANDS

The largest part of the Southern Blue Ridge province
consists of ranges of high mountains, many of which ex-
ceed 1,500 m in altitude. Mount Mitchell, the highest
mountain in the Eastern United States, is 2,040 m in
altitude. Local relief is generally high, exceeding 900 m.
Along the northwest and southeast borders of the Moun-
tain highlands the ranges tend to be roughly parallel to
the northeast regional strike of the rocks. In the center,
however, the ranges have various trends that seem to be
determined by the complex pattern of deep basins and
valleys that cut through the mountains in different
directions.

Some rock control of topography is evident. For exam-
ple, the Great Smoky Mountains have many peaks
higher than 1,800 m. The highest part of the range is
underlain by massive conglomerate and sandstone of the
Thunderhead Sandstone. The trends of many valleys in
the Mountain highlands parallel the strike of the rocks,
and some deep basins are related to large outcrop areas
of readily weathered rock types, for example, the low
areas of the Spruce Pine district in which feldspar-rich
rocks are extensive. On the other hand, the general im-

82" 37°
\

 

 

ZPO KILOMETEHS

FIGURE 20. —Map of the Southern Blue Ridge province, showing the subdivisions discussed in the text. Ruled line pattern represents high-
relief areas in the Southern Blue Ridge province. Dotted areas are the principal basin and trench areas: AB, Asheville basin; SP, Spruce
Pine area; WB, Waynesville-Canton basin; MM, Murphy Marble belt; LC, Laurel Creek trench.

rt

 

SOUTHERN BLUE RIDGE PROVINCE 29

pression given by the region does not suggest rock con-
trol like that seen in the Ridge and Valley province or in
the Chilhowee—Walden Creek belt.

The intermontane basins are irregular, have angular
boundaries, and are often connected with other basins
by deep valleys that look like trenches. This appearance
is due to the fact that the valleys commonly are straight
and cut across drainage divides at high angles. They are
referred to herein as trench valleys. The intersecting
pattern of linear valleys and interconnected basins sug-
gests an origin due in part to control by nonresistant
rocks and in part to tectonic influences. The tectonic in-
fluences could have involved faulting in the late geologic
time or, more likely, differential erosion along brittle
fracture zones associated with older faults.

Several trenches and basins show interesting correla-
tions with other features that shed light on their origin.
The most impressive of these is the long deep trench
that corresponds to the outcrop area of the Murphy Mar-
ble and Andrews Schist of Cambrian(?) age; these for-
mations are the youngest in a great syncline folded into
the rocks of the southwestern part of the Mountain high—
lands (MM, fig. 20). The trench is approximately 150 km
long and extends north through Georgia into North
Carolina. The beds that floor the trench in the North
Carolina part are marble, phyllite, and mica schist
(Hadley and Nelson, 1971; Kish and others, 1975). In
Georgia, where the marble occurs only locally along the
belt, the relief is fairly high. In North Carolina, however,
the marble has a continuous outcrop and forms a spec-
tacular trench far below the mountains on either side.
The trench culminates in the Nantahala Gorge, whose
floor is 450 to 900 m lower than the mountaintops on
either side (fig. 21), and is in places less than half a
kilometer wide. The gorge ends at the termination of the
marble outcrops.

At no place does the trench contain more than a head-
water stream, although it crosses the large Hiwassee
River and its northern end is near the Little Tennessee
River. The trench cuts directly across the main
westward-flowing drainage system with little change in
altitude. This trench is significant because it
demonstrates conclusively that where contrasting rock
types occur, rock control of topography is important in
the Blue Ridge. The immediate local relief along the
trench matches in magnitude the relief in the Ap-
palachian Valley, where carbonate rocks are adjacent to
wide outcrop belts of rocks such as quartzite and
massive sandstone. The morphology of the trench on the

_Murphy Marble is well shown on the Knoxville, Chat-

tanooga, and Rome sheets of the 1:250,000-scale
topographic-map series distributed by the U.S.
Geological Survey.

 

I Another spectacular trench is north of Asheville,
NO, and trends approximately east (LC, fig. 20). It is
referred to in this report as the Laurel Creek trench
after a stream that occupies the trench near its western
end. The profile of part of this trench is shown in figure
21. The trench crosses several drainage divides. The
Cane River is a large stream that in crossing the trench,
occupies it for about 14 km. The South Toe River also
crosses the trench, occupying it for about 2.5 km. The
Laurel Creek trench is about 750 m lower than the
mountain ridges north of it (fig. 21). The trench crosses
the Spruce Pine basin at its eastern end and terminates
against the Grandfather Mountain window. The
geologic map of the Knoxville Quadrangle (Hadley and
Nelson, 1971) shows that the trench is occupied for
much of its length by a right-lateral fault having a
horizontal displacement of about 2.5 km near the
western end. The Laurel Creek trench crosses the strike
of the rocks at a 50° angle, so that it is not related to
weak-rock strata, as is the trench on the Murphy Mar-
ble. On the other hand, the presence of the fault sug-
gests that the trench may owe its origin to a shear zone
that has the same erosional effect as a narrow belt of
weak rock. A sheared zone, if it contains fractures that
allow the passage of ground water, would be a zone of
low resistance to weathering and erosion.

The outlines of intermontane trenches, valleys, and
basins can be seen clearly on map separates showing the
topography only, at 1:250,000 scale. On such maps, the
differences in topographic relief between mountains and
low-relief areas are shown by differences in spacing of
contours, and areas of low relief are much lighter in
tone. Because the mountainous relief is generally quite
high, as much as 1,000 m Within a 100-km2 area, the con-
trast with basins having relief of less than about 200 m is
great. Figure 22 is an outline map of part of the Moun-
tain highlands made by tracing the boundaries between
sparse and dense contours. The map also shows the
Brevard zone (Hadley and Nelson, 1971; Goldsmith,
Milton, and Wilson, 1978; Nelson and Clarke, 1978) and
the outcrop of one map unit referred to as gneiss and
migmatite on the map of Hadley and Nelson (1971). This
unit seems to be a preferred rock type for the formation
of intermontane basins, although it is not known to be
particularly nonresistant. It is outlined here simply to
show the general structural trend of the rocks in a cen-
tral part of the area. Several large trench valleys that
crosscut the structural pattern of the strata are clearly
shown. A long trench valley shown in part on figure 22
contains a fracture zone known as the War Woman frac-
ture (or shear) zone (Pickering and Higgins, 1979;
Georgia Geological Survey, 1976). Some other trench
valleys are identified by names coined by the writer.

30 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

VALLEY RIVER

 

 

NANTAHALA RIVER

 

 

 

sw NE
1500 5000
1250 4000
l—
g 1000 3000 w
E u.
w 750
2 500 2000
250 ////////////////// / .000
TRENCH VALLEY ON MURPHY MARBLE
LAUREL BALD CANE LITTLE CRABTREE
w CREEK CREEK RIVER CREEK E
A A A
I W 11 I
1500 5000
1250 SOUTH TOE 4000
RIVER
2 1000 +-
.”_J y 3000 53
E 750 ""

250

/W ////// ::::

LAUREL CREEK TRENCH VALLEY

0 10
L I

2'0 30 KILOMETERS
J

FIGURE 21.-Profiles along two trench valleys in the Mountain highlands. The wavy heavy lines represent the profile along ridge crests im-
mediately on the north or northwest flanks of the valleys. The lines on the shaded areas represent the profiles of the valley floors.

Silicified breccia zones mapped by Hadley and Nelson
(1971) are shown by lines of triangles.

As the map shows, many valleys have distinct linear
trends. Some valleys clearly parallel the strike of the
rocks or the trends of thrust faults that parallel the
strike. Examples are the trench valley that trends
southwest from Waynesville (W, fig. 22) to the
Tuckasegee trench and, of course, the Brevard zone.
Other valleys clearly crosscut the structural grain, as do
major basins such as the Little Tennessee and the
Asheville basins.

Acker and Hatcher (1970) have identified rectilinear

systems of valleys controlled by joint systems at right.

angles to the structural grain. Spectacular valleys as
deep as 100 m are found in the drainage areas of the
Chauga, Tugaloo, and Chattooga Rivers in northeast
Georgia and northwest South Carolina. The spacing of
the valleys is controlled by the drainage pattern, but the
flow directions and linearity are commonly controlled by
joint directions. Some of the examples illustrated resem-

 

ble the trench valleys and can clearly be seen on
topographic maps at 1:250,000 scale.

BREVARD ZONE AND HENDERSONVILLE BULGE

The Brevard zone is a broad zone of faults and fault
slices that extends from Virginia to Alabama. It has had
a long history involving several periods of activity
possibly extending as far back as the Precambrian
(Rankin, 1975). The youngest tectonic features exposed
at the surface are brittle-fracture phenomena (Hatcher,
1978). The Brevard zone is itself a major topographic
feature generally forming a narrow zone of low relief.
At its northern end near Winston-Salem, NO, the fault
zone is about 35 km south of the Blue Ridge escarpment.
In the Morganton basin, it is at the foot of the Blue
Ridge escarpment. For about 122 km southeast, it rises
to cross the Mountain highlands for a distance of about
120 km. In this region, it forms a prominent topographic
boundary. The mountain region southeast of the
Brevard zone is distinctly lower than the area northwest

.51
.0

L

41

V ‘ SOUTHERN BLUE RIDGE PROVINCE 31

 

 

 

 

 

 

       
  

\fl
,
\\|

600‘. l675

 

f‘
f

Y
‘r

 

10 20 30 40 50 KILOMETEBS
1 I I l J

 

 

 

 

 

FIGURE 22.-Map of a part of the Mountain highlands, showing the larger trench valleys and intermontane basins. Patterned areas are

" rugged terrain having generally high relief. Unpatterned areas are low-relief areas. Lines of triangles represent silicified breccia zones (from
-.. Hadley and Nelson, 1971). Fine dashed line represents boundary of gneiss and migmatite unit of Hadley and Nelson (1971). Letters in circles
represent Asheville (A), Canton (C), Hendersonville (H), Waynesville (W), Sylva (S), and Bryson City (B).
b
I i

32 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

of it and has a different structural grain (fig. 22). The
mountains and trench valleys within this lower region
are arranged in linear and arcuate patterns. The arcuate
features are related to the strike of the underlying
rocks. The area is like a huge berm, bulging outward in-
to the Piedmont from the main Mountain highlands. For
convenience of discussion, it is here called the “Hender-
sonville bulge.”

The topographic evolution of the Hendersonville bulge
is probably related in part to the drainage system and in
part to rock resistance. The drainage of the bulge is
mostly northwest of the eastern continental drainage
divide to the Tennessee River via the French Broad
River. The fact that its general level is higher than that
of the Piedmont or southeast drainage probably explains
the difference in altitude between the bulge and the
Piedmont (Hack, 1973b).

However, the difference in relief between the Hender-
sonville bulge and the Mountain highlands on the north-
west side of the Brevard zone must have another cause.
The rocks of the bulge are considered part of the Pied-
mont province by geologists. The most extensive forma-
tion is the Henderson Gneiss, described by Hadley and
Nelson (1971) as a medium- to coarse-grained foliated
biotite-microcline augen gneiss. The degree of foliation
and shear increases and becomes intense near the
Brevard zone. In my judgement, such a rock might be
expected to be much less resistant than many of the
rocks in the Great Smoky Group to the northwest. Thus,
the large outcrop area of Henderson gneiss may explain
wholly, or in part, the difference in relief.

The middle reaches of the French Broad River (fig. 23)
are in the Henderson Gneiss just southeast of the
Brevard zone. The river has a low gradient in this area
(0.00053 over a distance of 34 km). Where the river
leaves the Brevard zone to enter the Asheville basin, the
slope increases to 0.00097 for 19 km. The next reach
downstream passes through Asheville and has a slope of
0.0013. The gradient continues to increase to a max-
imum of 0.005 in the Longarm quartzite. This kind of
convex profile is possible in a system that has achieved a
state close to dynamic equilibrium, where the rocks
become increasingly more resistant downstream. On the
other hand, an increased rate of uplift northwest of the
Brevard zone, relative to the Hendersonville bulge,
could also result in a strongly convex-upward stream
profile if the stream were uplifted relative to the Ap-
palachian Valley. Thus, the explanation for the different
relief in the two areas may involve one or both of two
causes.

Several lineaments and conspicuous trench valleys
that may be structurally related to the Brevard zone
trend away from it at low angles. The Catawba River
lineament in the Piedmont is one of these features. This

 

lineament is followed by the river for a distance of 50
km. Only the western end is shown in figure 22. Surface
geologic mapping has disclosed no fault (Bryant and
Reed, 1970; Goldsmith, Milton, and Wilson, 1978). The
Laurel Creek trench west of the Brevard zone trends
east, in the same direction as the Catawba River linea-
ment. This feature, as mentioned above, does contain a
fault having an offset of 2.5 km. The Swannanoa trench
intersects the Brevard zone at an angle of about 20°. No
offset along this zone has been identified, but the trench
is a prominent lineament that connects the Morganton,
Asheville, and Canton-Waynesboro basins. It may con-
tinue westward as the Bryson City trench.

The Tuckasegee trench trends northwestward from
the Brevard zone. This crosscutting trench is note-
worthy because of its straightness and the fact that it
separates mountain masses having different general
altitudes. The mountain peaks south of the trench are
about 300 m lower than those north of it.

The silicified breccia zones in the southeastern part of
the area shown in figure 22 are of interest because they
show that zones of brittle fracture have formed that do
not displace the geologic contacts to a large degree. The
breccia zone at locality 1 (fig. 22) is crossed by a Triassic
dike that is not displaced within the limits of mapping
(Hadley and Nelson, 1971). However, Snipes (1979)
reported that the breccia zone at locality 2 offsets a ver-
tical Mesozoic dike at least 175 m. The amount of ver-
tical offset, if any, cannot be determined. This evidence
indicates that fracture zones do exist in the area that in-
volved offsets in post-Paleozoic time. Silicification of the
breccia zones in the Piedmont is evidence that solutions
have percolated through them. Brittle fracture zones
may have localized rapid erosion in many places in the
Mountain highlands and were perhaps the loci of small
differential movements in post-Paleozoic time that af-
fected the gross morphology of the terrain.

INTERMON TAN E BASINS

The origin of the basins within the Mountain highlands
is another difficult problem. Most geomorphologists who
have worked in the area have considered these basins to
be straths formed when the rivers that drained them
were flowing at a lower grade, generally correlative
with the Harrisburg surface of the north (see Thorn-
bury, 1965, p. 107). Wright (1931) thought that rock
resistance had a strong influence on the position and
altitude of these basins. The idea that a sill of resistant
rock slowed relative erosion rates along the streams so
that a low-relief area could migrate up the tributaries is
appealing, and some such mechanism may have been a
part of the process of basin formation. On the other
hand, the basins themselves have patterns and outlines

 

 

SOUTHERN BLUE RIDGE PROVINCE

 

 

 

 

 

{ pegs g/ff/f/ HENDERSONVILLE N
// Mn f/J’J/ .00053 ° /
K ‘ ’ W
w/Qfi
I ngc / //§&‘ I hg
/ // /fé-/ fi-/
/ /f} / r3
/ // éH/ / V~’f —-—
/ ,,
t~/\ / fij/ [xx/J

ck“%‘{,‘?s/§4 ’Z/
- 2%“
r / 0 10 20 30 40 50 KILOMETERS
’ I I I I I 4

 

FIGURE 23. —Geolog'ic map of the basin of the French Broad River (Asheville basin shaded). Numbers represent gradients of
reaches of the French Broad River between double bars. Letters in squares are the first letters of the quadrangle names
listed in table 2: H, Henderson; A, Asheville; L, Leicester; M, Marshall. Geologic units shown are: ms, metasedimentary
rocks of Proterozoic and Cambrian age that include quartzite, sandstone, graywacke, phyllite, and shale; peg, Max Patch
Granite and Cranberry Gneiss and migmatite; png, layered gneiss and migmatite; pegs, Great Smoky Group; hg,
Henderson Gneiss and related rocks; pch, biotite schist and gneiss. (From Hadley and Nelson, 1971.)

33

34 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

related to the rock structure and to the trench valleys
that suggest a more complex origin. Several examples
can be cited. The complex basin at Waynesboro and Can-
ton (W and C, respectively, fig. 22) is at the intersection
of at least two trench valleys and within the outcrop
area of the layered gneiss and migmatite, which is a rock
unit quite variable in composition. The broadest part of
the Asheville basin is within the outcrop area of the
same gneiss and migmatite unit. The complex basin at
Spruce Pine (SP, fig. 20) is on highly feldspathic rocks as
well as a fault zone. The Henderson basin (H, fig. 22) and
the low-relief area to the west along the Brevard zone
are in the Henderson Gneiss, a highly foliated rock.
Nevertheless, the evidence is not sufficient to indicate
that rock control by itself is responsible for all the
basins.

The Asheville basin is the largest (fig. 23), and the
problem of its origin is probably typical of other basins.
It forms a large low-relief area having an average
altitude of about 600 m. The floor of this area is almost
the same altitude. Table 2 shows hypsometric data for
four rectangular areas in this basin, within the Hender—
sonville, Asheville, Leicester, and Marshall Quadrangles
(fig. 23). As the table shows, the altitudes of the highest
points in these areas do not decline systematically
downstream. The highest point of all is in the Marshall
Quadrangle, the farthest downstream. However, the
mean altitude declines 41 m, whereas the river declines
128 m, reflecting the systematic increase in total and
mean relief.

The elevation-relief ratio is a measure of the sym-
metry of the topography. It is the ratio of the mean
relief to the total relief. It expresses the proportionate
amount of the land surface lower than the mean relief
and thus reflects the forms. As the ratio in this area is
close to 0.5 and does not change systematically
downstream, there is no indication that the landscape
has relict features caused by dissection of an older sur-
face. The data seem to indicate only that the relief in-
creases downstream and that the local topography is
symmetrical and adjusted to the present river. The in-
crease in relief may be due to the overall increase in rock

 

resistance downstream. The possiblity exists, however,

that tectonic uplift or tilting could produce the same ef-
feet.

The question of origin still remains. The relief in the
Asheville basin is significantly lower than that of the
mountains around it, and the broader outlines of the
area, at least, are not controlled entirely by rock types.
A possible explanation is that the area is affected by in-
tensive faulting and shearing along many zones. The ab-
solute amount of movement in these zones may have
been small, but the zones provided a network of frac-
tures along which erosion could proceed. The altitudes
of the floor of the basin might have been controlled by a
sill of massive and resistant rocks that crosses the
French Broad River at the lower end of the basin.
Possibly the course of the French Broad itself was deter—
mined by a tectonic line or zone of weakness transverse
to the structure. Also belts and areas of weak rocks may
occur in this and other basins that cannot be identified in
the available reconnaissance maps. Detailed fieldwork
may be required to correlate the bedrock patterns with
topographic patterns in order to resolve this problem.

MOUNT ROGERS AREA

Mount Rogers (1,750 m) and White Top Mountain
(1,630 m) just north of the North Carolina—Virginia
boundary rise conspicuously higher than their surround-
ings. They are underlain by porphyritic welded rhyolite
tuff unique to this area. The surrounding lowlands are
underlain by rocks of the same formation (Mount Rogers
Formation) but consist of metasedimentary rocks that
include tillite, graywacke, siltstone, shale, and tuff
(Rankin and others, 1972). North of this local area, the
Iron Mountains, underlain by the Chilhowee Group, rise
to a lesser height (1,410 m). Thus, in the Mount Rogers
area, rock resistance is closely correlated with dif-
ferences in relief.

NEW RIVER PIA'IFAU

The northeastern part of the Southern Blue Ridge
province is a plateau region having comparatively low
relief, relieved in a few places by low mountains. Most of
the area is graded to the New River, which flows out of

TABLE 2. —Comparison of topography at four locations in the Asheville basin

[Based on altitudes of 64 points in a rectangular grid 20 km2 in area. At all three localities, the grid is centered on or close to the channel of the main stream.
Quadrangles at 124,000 scale; 20-ft contours]

 

Lowest point

 

 

 

 

Quadrangle Highest point (river surface) Total relief‘ Mean relief Mean altitude Elevation-relief
(m) (m) (m) (m) (m) ratio?
Hendersonville 664 633.5 30.5 12.8 646 0.42
Asheville 670 600 70 30.5 631 .43
Leicester 670 524 146 84.0 608 .57
Marshall 701 505 195 100 605 .51

 

 

1 Difference in altitude between highest and lowest points.
2 Ratio of mean relief to total relie ;

essentially the same as the hypsometric integral (see Pike and Wilson, 1971).

 

 

 

 

 

SOUTHERN BLUE

the area and into the Appalachian Valley at an altitude
of 610 m. The New River plateau is mostly less than 900
m in altitude. Where there are no mountains, the
average relief in a 100-km2 area is about 200 m. In such
places, the relief of the area is higher than that of the
Piedmont but much less than that of other parts of the
Blue Ridge. The streams cross the New River plateau in
sinuous courses, and form exaggerated meanders that
have long straight reaches trending northwest at right
angles to the strike of the bedrock, much like the
meanders that are common in Martinsburg shale in the
Appalachian Valley (Hack and Young, 1959).

The New River plateau as a whole appears to owe its
distinct character largely to rock control. Most of the
area is underlain by thinly layered schist and gneiss of
the Ashe Formation. The southeastern margin is
underlain by a thinly laminated schist and gneiss of the
Alligator Back Formation (Rankin and others, 1972;
Espenshade and others, 1975). These laminated and
thinly bedded fine-grained rocks contrast sharply with
the quartzite in the Chilhowee—Walden Creek belt, the
massive gneiss of the “Elk Park plutonic group,” and the
large bodies of massive amphibolite in the Ashe Forma-
tion that border the plateau in the Mountain Highlands
to the west. The southeastern boundary of the plateau,
however, except for one small area, is coincident with
the drainage divide separating the New River drainage
from streams that flow to the Atlantic. This divide is 610
m higher than the Piedmont surface at the southwestern
end of the boundary but only 450 m higher at the north-
eastern end. The Piedmont surface as well as the escarp-
ment rising to the plateau is underlain by the laminated
schists of the Alligator Back Formation, the same rocks
that underlie the adjacent plateau. The origin of the
escarpment is treated in some detail in the following sec-
tion of this report, but the higher altitude of the New
River plateau is believed to be related to the extensive
outcrops of resistant sandstones and quartzite cut by the
New River in its lower reaches that form a high base
level for the entire river basin (Hack, 1973b).

BLUE RIDGE ESCARPMENT

The Blue Ridge escarpment, as defined here, is the
narrow strip of steep land that drains southeastward to
the Piedmont and Atlantic Ocean. It extends southward
from the Roanoke River and ends near Rossman, N.C.,
on the headwaters of the Keowee River, a major
tributary of the Savannah River. Most geologists who
have written about the escarpment have been concerned
only with the steep and spectacular part northeast of the
Catawba River headwaters. In that region, the scarp
rises abruptly above the Piedmont surface. The crest
corresponds generally to the drainage divide and can be

 

RIDGE PROVINCE 35

seen from many places at the foot. Southwest of the
Catawba River, the crest is separated from the low
country by one or more narrow ranges of mountains.
Nevertheless, the steep rise from the Piedmont is im-
pressive. Southwest of the Keowee River, the southeast
slope of the Blue Ridge physiographic province is a little
broader, the drainage is mostly toward the Gulf of Mex—
ico, and the topography is too complex to be called an
escarpment. At the northeastern end of the Southern
Blue Ridge province, the drainage divide crosses the Ap-
palachian Valley near Blacksburg, Va., forming an ir-
regular but impressive escarpment in the valley between
the Blue Ridge and the Appalachian Plateau. This part
of the escarpment has been retreating southwest by
headwater extension of valleys in weak rocks, primarily
limestone. It is not considered in this report.

Four hypotheses explaining the Blue Ridge escarp-
ment have been discussed in the past (see Thornbury,
1965, p. 104—106). The first hypothesis was that of
Hayes and Campbell (1894), who believed that the
escarpment is essentially a monoclinal flexure of a
peneplain that was later dissected. According to Camp-
bell (1896), as uplift takes place on the up or northwest
side of the monoclinal flexure, stream erosion toward
the southeast is accelerated and works headward, pro-
ducing a scarp.

A second hypothesis was that of Davis (1903), who
suggested that the scarp is a topographic discontinuity
separating a set of two dissected peneplains of the same
age but on opposite sides of the major drainage divide
between streams flowing to the Atlantic Ocean and
those flowing to the Gulf of Mexico. He thought that
streams flowing to the Atlantic had an advantage over
streams flowing to the Gulf of Mexico by way of the
Mississippi River system because they had a shorter
distance to travel to reach the sea. Davis also intro-
duced, as part of his hypothesis, the idea that the ex-
istence of barriers to erosion in the form of resistant
bedrock on the northwest side of the divide would in-
crease the advantage of the eastern streams. There is a
serious problem with the idea that distance to the sea,
the ultimate base level, determined the escarpment,
however, because the altitude of the base of the Blue
Ridge escarpment on the southeast in some places is
higher than that of the base of the mountain front on the
opposite side of the Blue Ridge province. For example,
the Catawba River enters the Piedmont at an altitude of
about 1,500 feet. The French Broad River, however,
which shares a drainage divide with the Catawba River,
enters the Appalachian Valley on the Mississippi River
side of the divide at an altitude of only 1,200 feet.

Wright (1931) emphasized Davis’ idea that the greater
height of the northwestern drainage of the Blue Ridge is
related to the greater resistance of the rocks on the
northwest side of the divide.

36 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

White (1950) introduced the hypothesis that the scarp
was caused by faulting on a series of en echelon faults at
the foot of the scarp, and he extended this concept
northward along the entire Blue Ridge province to
Maryland. In his hypothesis, the Asheville basin and
New River plateau were once at the level of the Pied-
mont and are now upfaulted peneplain surfaces. Stose
and Stose (1951) criticized the faulting hypothesis, poin-
ting out that the Brevard zone is the youngest large
fault zone identified in the region and is a zone of
thrusting, the Piedmont being on the upthrown side.
Furthermore, the zone coincides with the escarpment
for only 50 to 60 km. It departs from the escarpment
both to the northeast, where it is farther east in the
Piedmont, and to the southwest, where it is within the
Blue Ridge Mountains.

Another idea is that, although faulting may have been
a factor, the faults were some distance from the present
scarp. The discontinuity in the topography produced by
faulting was maintained as the divide slowly shifted
northwest (Thornbury, 1965, p. 105). This idea or the
older idea of Hayes and Campbell (1894) seems close to a
satisfactory explanation. The peneplain is not a
necessary part of the explanation. All that is needed are
a discontinuity between the average altitudes of the
drainage systems on opposite sides of the divide and a
way to maintain that discontinuity. Either a fault or a

84.

 

 

monoclinal uplift could produce an escarpment main-
tained by resistant rocks in the lower reaches of the
main streams northwest of the divide (fig. 24). The
divide may be near its original position at the time of
uplift, or it may have retreated some distance as
downwasting took place (Hack, 1969, 1973b).

When the escarpment is viewed as a whole, it is in-
teresting that the average altitude of the escarpment
base does not change over its entire length of 350 km.
The base is approximately 450 m above sea level and
generally varies less than 60 m above or below. This con-
sistency of altitude is related to the general low slope
and low relief of the inner Piedmont. It is also related to
the sharply defined rise of the escarpment from the foot.
The outlying ranges of mountains southeast of the
escarpment have basal margins 60 to 100 m lower,
depending on their distance from the escarpment.
Figure 3 shows the regular configuration of the Pied-
mont surface along its inner margin.

The width of the escarpment and the height of the
escarpment crest above its base, on the other hand, do
change regionally. These dimensions are related to the
regional geology and the configuration of the drainage
system (see fig. 25). Near the Roanoke River, at the
northeastern end, the height of the escarpment is 275 m.
In this area, the escarpment is steep, and its width, that

 

  
  
 

 
 

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EXPLANATION

 

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Gneiss and plutonic rocks of
older Proterozoic basement

Low-grade metasedimentary
rocks

FIGURE 24. — Map of the drainage system of the Southern Blue Ridge physiographic province, showing its asymmetry. The patterned areas
represent two major geologic units containing many resistant rocks. Dotted areas are clastic sequences of Proterozoic Z and Cambrian
age, including graywacke-shale sequences and many quartzites. Hatched areas represent the crystalline basement rocks of Proterozoic
Z age, mostly gneiss and plutonic rock. Geology from Williams (1978).

SOUTHERN BLUE RIDGE PROVINCE 37

36° 82°

 

 

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36° 82°

 

 

o 50 100 KILOMETERS
EXPLANATION

 

Blue Ridge escarpment

Brevard zone

Henderson Gneiss

FIGURE 25. — Strip maps showing the southeastern margin of the Southern Blue Ridge and the foothills in the adjacent Piedmont. A, Shaded
area represents the area of the Blue Ridge southeast of the drainage divide as well as the slopes of the outlying foothills in the Piedmont.
B, Geologic units are adapted from maps by Hadley and Nelson (1971), Rankin and others (1972), Espenshade and others (1975),
Goldsmith, Milton, and Wilson (1978), and Milici and others (1963). brt, gneiss, schist, and plutonic rocks of Blue Ridge thrust sheets, un-
differentiated; sp, rocks of Spruce Pine area (mostly coarse-grained feldspathic gneiss); gfw, rocks of Grandfather Mountain window; a,
Ashe Formation, mostly thinly layered biotite schist and gneiss; ab, Alligator Back Formation, finely laminated gneiss; sra, gneiss and
schist of Smith River allochthon; smw, mostly gneiss and plutonic rock in Sauratown Mountain window; hg, Henderson gneiss; i p, rocks
of Inner Piedmont, includes biotite and amphibolite gneiss, sillimanite-mica schist, granite, migmatite, and quartz monzonite; ch,
Chilhowee group; brc, Blue Ridge complex; and s, carbonate rocks and shale of early Paleozoic age.

is, the horizontal distance from foot to crest, is 2 to 3
km. From this area southwest to Stone Mountain, the
height and width change very little. The entire escarp-
ment is within the Alligator Back Formation, a weak
rock consisting of thinly laminated micaceous schist and
gneiss (Espenshade and others, 1975). The stream pro-
files on the escarpment itself are highly concave, a
characteristic of headwater profiles in weak rocks.
Thus, the aspect of the scarp is impressively steep. 0n
the Blue Ridge, or northwest side of the crest, the
topographic relief is low, and the area drains northward
to the New River, a meandering stream that flows
northeast almost parallel to the scarp toward the Ap-
palachian Valley.

In the headwater area of the Ararat River (fig. 25A),

 

the crest of the escarpment is farther from the base.

From this place to the northeastern margin of the
Grandfather Mountain window, the escarpment has an
average width of about 15 km. The crest rises to a height
of 460 m above the base, much higher than the escarp-
ment farther north. The streams drain the face of the
broad escarpment through deep valleys separated by
ridges. The different aspect of the escarpment when it is
compared with the area to the northeast is related to
two factors: First, although the country rock on the
escarpment face is still the weak Alligator Back Forma-
tion, many dikes and sills of granitic rocks are intruded
into the formation in this region (Rankin and others,
1972). Second, on the upland to the northwest, the New
River flows through a phase of the Ashe Formation that
contains amphibolite and hornblende gneiss. The local
relief is higher than the relief downstream, and the gra-

38 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

dient is steeper; thus, the small tributaries that drain
toward the New River have their heads at altitudes
higher than those of streams farther northeast along the
escarpment crest.

In the Grandfather Mountain window and southwest
of it, the local geology is comparatively complex and
changes along the escarpment. The escarpment is
generally broad but varies in height above its base. In
the window, the width reaches 23 km, and the maximum
height above the base in the Piedmont is 1,000 m. This
high point is at Grandfather Mountain, where the
escarpment crest is underlain by an arkosic quartzite
conglomerate. The lowest point within the window is in
the Blowing Rock Gneiss, a feldspathic augen gneiss.

Between the Grandfather Mountain window and the
head of the Morganton basin on the Catawba River, the
escarpment narrows to 11 km. The height of the crest
above the base ranges from 400 m near Spruce Pine to
1,000 m at a spur of Mount Mitchell immediately north-
west of the escarpment. Because of the great height and
narrow width of the escarpment, the Morganton basin is
a spectacular place from which to view it. The northwest
margin of the Morganton basin is underlain by sheared
migmatite and sheared rocks of the Brevard zone. To
the northwest, the rocks underlying Mount Mitchell are
coarse metagraywacke containing quartzite mapped by
Hadley and Nelson (1971) as part of the Great Smoky
Group. It is the juxtaposition of contrasting rock belts
that gives the escarpment its grandeur at this point.

Southwest of the Catawba River, the character of the
escarpment changes. It becomes a less regular mountain
front that lies southeast of the drainage divide. This is
the area of the Hendersonville bulge already described,
a half-moon—shaped basin drained by headwaters of the
French Broad River. The basin is bordered on the
southeast by mountains, many of which rise to altitudes
greater than 900 m. The drainage divide is close to the
inner margin of these mountains. Near Hendersonville,
the divide is on the basin floor, and three headwater
streams drain through the mountains to the Piedmont
by way of the Green and Broad Rivers. As already
described, many of the streams have probably worked
headward in the mountains along fracture zones and
other places of weakness. This area is different from the
areas of the escarpment previously described, in that the
crest is not on the drainage divide. However, this dif-
ference may be an accident of the local geology. The
streams flowing toward the Atlantic have penetrated
relatively resistant rocks such as massive quartz mon-
zonite, granodiorite, and paragneiss along weak but nar-
row zones.

Important questions related to the history and origin
of the escarpment remain. Geologic evidence reinforced
by recent geologic mapping makes it clear that faulting

 

is not the direct cause of the escarpment. At present,
the escarpment is maintained by a difference in stream
profiles on rocks of different resistance on either side of
the divide. Can evidence be found that the escarpment
has migrated from some position southeast of the pres-
ent one? If so, has it migrated far? One way to appraise
this problem is to look for recent stream captures.
Presumably, if a southeast-flowing stream captured a
stream on the Blue Ridge upland as it worked headward,
it would take some time to destroy the low relief in the
captured part of the upland. If the Blue Ridge escarp-
ment is migrating rapidly toward the northwest, one
might expect to find areas on the upland surface that
now drain toward the Piedmont, areas in which gorges
that are working headward have formed. Only two areas
of significant size were found in which streams draining
to the Piedmont are northwest of the escarpment crest.
One is the head of the Dan River, which, upstream from
the crest, has a basin 92 km2 in area. The evidence that
this upland basin has been captured by the Dan River is
unequivocal. The Linville River has a headwater basin
108 km2 in area northwest of the escarpment crest, and
this occurrence has often been cited as an example of
stream capture. However, I believe that this river in-
herited its course down the escarpment by superposition
on resistant rocks. Both occurrences can provide clues
to the origin of the escarpment.

DAN RIVER CAPTURE

The Dan River has its source on the Blue Ridge
uplands about 15 km upstream from the escarpment
crest. As Dietrich (1970) discussed, the river has worked
headward from the escarpment and has captured the up-
per basin of Reed Island Creek. The geology of the head-
water area is relatively simple, and as the entire head-
water area of the Dan River is in the same geologic unit
(fig. 25), the anomalous profile is probably not related to
differences in the rocks. The unit in which the capture
took place is the Alligator Back Formation (Espenshade
and others, 1975), typically a finely laminated gneiss
composed of quartzofeldspathic layers a few millimeters
thick separated by thin micaceous partings. It underlies
areas of low relief on the upland and can be regarded as
a weak rock. The drainage pattern is shown in figure 26,
a subenvelope map of the upper Dan River basin. The
contours are drawn on the bottom of the stream chan-
nels and show in a general way the vertical configura-
tion of the drainage.

The drainage pattern and its relation to the crest of
the escarpment suggest that a capture took place some
distance downstream from locality B (fig. 26). The Dan
River, cutting northward into the upland, captured the
former headwaters of Reed Island Creek, which, at the

LA.

 

SOUTHERN BLUE RIDGE PROVINCE
80°30'

' \

 

 

 

 

 

0 m 20 MILES
i 1 1 I L I
0 10 20 KILOMETERS

CONTOUR INIEHVAL 1525 m (500 ft)
FIGURE 26. —Map of the headwaters of the Dan River, showmg the area of stream capture. Streams shown by

dashed lines flow toward Atlantic Ocean. Subenvelope contours are drawn on altitudes of the stream chan-
nels. Major geologic boundaries are shown by hachured lines. Geologic units modified from Espenshade and
others (1975). sra, gneiss, schist, and gneissic granite of Smith River allochthon; a, Ashe Formation; ab,
Alligator Back Formation; sp, granitic rock of “Spruce Pine plutonic group” (Mount Airy Granite of
Stuckey and Conrad, 1958).

39

40 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

time, was close to the crest of the escarpment and
flowed parallel to it. This idea is strongly supported by
the modern profile of the Dan River (fig. 27). A striking
feature is the sharp break shown near locality B on the
profile. The profile downstream from the break is sharp-
ly concave. The concavity is inherited in part from the
original profile of the headwater stream before the time
of capture. At that time, the stream had reduced dis-
charge, less power, and steep slopes. As a result of the
capture and the increased discharge and power, the
steeply concave section or knickpoint worked headward.

Figure 27 shows that the profile above the knickpoint
is slightly convex. This is related to the downward com-
ponent of erosion at the knickpoint. The figure also
shows that the stream is intrenched for about 22 km
above the knickpoint, as measured along the channel. At
about 16 km from the streamhead, the valley is only 60
m deep, normal for valleys on the upland.

The observed conditions show that the river has had a
complex response upstream from the capture, as wOuld
be expected from the experiments of Schumm (1973, p.
306—307). The knickpoint is decreasing in height as it
works headward because of the capture and the in-
troduction of greatly increased discharge. In the Dan
River, the interpretation of the facts does not seem to be
complicated by differences in bedrock lithologies, as is
common.

Differences in stream power along the channel can be
expressed in a semiquantitative way by use of the gra-
dient index, which is a function of the slope or rate of fall
of a reach along its logarithmic profile (see fig. 10 for
definition). This index is expressed here in meters. On
the upland in the headwater area, the gradient index is
only 14 gradient m. The value graually increases in the
convex section, reaching 510 gradient m above the
knickpoint, where the index rises sharply to 1,700 gra-
dient m. At the steepest reach, the value is 2,900 gra-
dient m. It then drops gradually to 240 gradient m at the

 

base of the escarpment. Because of its high discharge,
the power of this river is much higher than that of
streams that head on the escarpment. Hookers Creek,
for example, immediately east of the Dan River, has an
index of 100 gradient m, 1.22 km from its head on the
escarpment crest. This index increases to 460 gradient
m at 1.9 km, then declines to 140 gradient m at 5.6 km
near the base of the escarpment.

The upper Dan River basin is bordered on both sides
by small upland areas that drain to the Piedmont. Two
small tributaries of the Smith River have worked head-
ward past the escarpment crest to capture 5 km2 of the
upland terrain. The Ararat River just west of the Dan
River has captured about 2.5 km2 of the upland. There is
no evidence, however, that any flowing streams were
captured by the headwaters of these streams. Close ex-
amination of the details of the Blue Ridge escarpment
will show many places where steep headwater streams
have penetrated small areas on the escarpment crest.
The simple geometry of the escarpment in this region,
where the rocks are essentially the same on both slopes,
dictates that, as downwasting takes place, the steeper
slope must migrate toward the gentler slope. The scarp
retreat is the horizontal component of downward ero-
s10n.

LINVILLE RIVER

The Linville River has often been cited as an example
of capture of the headwaters of a stream northwest of
the drainage divide by a southeast-flowing stream. The
river makes its descent to the Piedmont entirely within
the Grandfather Mountain window, an area of about
55x25 km, on the southeastern margin of the Blue
Ridge near Morganton, NC. In the window, rocks of
Proterozoic Y and Z age and rocks of Cambrian age are
exposed beneath great thrust sheets (see Bryant and
Reed, 1970). The Blue Ridge escarpment crest runs
diagonally across the window. The Linville River drains

S

A APPROXIMATE LEVEL OF UPLAND

METERS

o I i l
o 10 20 30

 

40 50 60 70

KILOMETERS

FIGURE 27. —Profile of headwaters of the Dan River. (Geologic units from Espenshade and others, 1975.) abg, typical finely laminated gneiss
of the Alligator Back Formation; aba, amphibolite, greenstone, and muscovite gneiss; srs, biotite quartz schist of Smith River allochthon.

SOUTHERN BLUE RIDGE PROVINCE 41

a low-relief area on the upland above the escarpment but
south of the divide.

Figure 28 is a drainage map showing the Linville
River and its relation to adjoining streams in the Blue
Ridge. Subenvelope contours having an interval of ap-
proximately 152.5 m connect points on the streambeds.
The position of the Blue Ridge escarpment is shown by
the closely spaced subenvelope contours between 457
and 914 m that cross the area from southwest to north-
east. As the subenvelope shows, the upper Linville River
basin is higher than the basins of rivers on either side.

The geology of the area (Byrant, 1965; Reed, 1964)
shows that the difference in altitude of the rivers is due
to rock control. The North Toe River flows in softer
rocks consisting of feldspathic schist and gneiss of the
Spruce Pine area (Hadley and Nelson, 1971; Rankin and
others, 1972). The Linville River, on the other hand,
owes its higher altitude to an extensive outcrop area of
quartzite in the Chilhowee Group exposed in the Grand—
father Mountain window. The geology along the river is
shown on the exaggerated profile of figure 29. In its up-
per course (north of lat 36°), the river is in the Grand-
father Mountain Formation of Proterozoic Z age. The
rocks of this formation, which range from phyllite and
siltstone to felsitic volcanic rocks and arkosic con-
glomerate, vary widely in their resistance. The river
tends to flow in the phyllite and siltstone or in alluvium
deposited on them. North of locality A in figure 28, the
river crosses the north tip of the lower quartzite unit of
the Chilhowee Group and moves over to the Cranberry
Gneiss, which it follows as far as Linville Falls, which is
at the 915-m contour and at the fault contact between
the Cranberry Gneiss and the Chilhowee Group.

At locality C in figure 28, the river winds steeply
through the quartzite unit in a series of angular
meanders. Although Wright (1928) thought that the
meanders were inherited from a peneplain at a higher
level, it is much more likely that the river channel is con-
trolled by structures in the quartzite, especially by minor
folds that trend northeast (see the geologic map, in
Reed, 1964). Where the river enters Wilson Creek
Gneiss, a massive rock, its profile steepens sharply and
is highly concave as far downstream as the faults that
bound the Grandfather Mountain window on the south.
The lower relief in the drainage basin of its type locality
in the Wilson Creek basin indicates that this gneiss unit
is somewhat less resistant than is the quartzite of the
Chilhowee Group. The profile of the Linville River in the
gneiss is extraordinarily steep, however, for a river of
its large discharge. The gorge bottom is cluttered with
huge boulders, some more than 6 m across, giving the
impression that the gorge is downcutting and back-
wasting rapidly. The quartzite forms a resistant capping

that insures steep slopes in the reach below, analogous ‘

 

to the Lockport Dolomite that forms a caprock above
shale beds at Niagara Falls, NY.

The North Fork of the Catawba River is also in rocks
of Cambrian age, the Shady Dolomite, the youngest
beds within the window. The dolomite itself does not
crop out because it is covered by alluvium and residuum,
but its presence in the valley bottom is revealed in an
underground cavern as well as by outcrops of residuum.
This dolomite explains the comparatively concave pro-
file of the river. Probably the gradient would be gentler
were it not for the many quartzite boulders that have
been brought down by short tributaries from the quartz-
ite slopes on the east. Details of the relation of these
rivers to the geology and topography are clearly shown
in the geologic maps and cross sections of Bryant (1965)
and Reed (1964).

The hypothesis that the Linville River has captured a
former tributary of the North Toe River has been
favored by those who believe that the drainage divide
and scarp are migrating northwestward by a series of
captures. However, no really satisfactory evidence of
the place of capture has been found. Low divides be-
tween the two drainages occur only above Linville Falls.
The lowest of these is Montgomery Gap (M, 'fig. 28),
where the head of a tributary of the North Toe is 75 m
higher than the Linville River less than 1 km to the east.
Wright (1931) rejected this divide as a point of capture
because it would divert only the upper part of this North
Toe tributary and would not explain the origin of the
reach of the Linville River from Montgomery Gap to
Linville Falls.

Other low divides (A and B, fig. 28) are not satisfac-
tory for other reasons. Suppose that a capture of the
Linville River took place at the place shown at locality A
on figure 28, where the drainage divide between the
North Toe and the Linville River is at an altitude of only
1,006 m. This low part is only 175 m above the Linville
River but about 230 m above the North Toe River. The
divide itself is on the Cranberry Gneiss, and the Linville
River flows on the Linville Falls fault along the contact
between the Cranberry Gneiss and the quartzite of the
Chilhowee Group. On the other hand, the North Toe is in
the relatively nonresistant rocks of the Spruce Pine
area. It seems impossible that the Linville River could
have captured a major tributary of the North Toe,
because the Linville River would have to work headward
through resistant quartzite to do so.

Wright (1931) believed that the capture took place at
the Linville gorge below the meander belt (D, fig. 28) at
a time before the cutting of the gorge, when the ridge
crest on the west side of the present river was, in his
view, part of a peneplain surface north of the escarp-
ment. The stream that is now the upper Linville River
was then a tributary of the North Toe but had a curving

42 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

82°

 

915

36°

 

 

 

 

 

 

20 MILES
l

 

20 KILOMETERS
CONTOUR INTERVAL 152.5 m (500 ft)

EXPLANATION

 

Chilhowee Group Grandfather Mountain Formation

FIGURE 28. —Map of the Linville River and surrounding areas. Drainage toward the Atlantic Ocean is shown by dashed lines. Barbed lines
show Table Rock fault and the Linville Falls fault, which bounds the Grandfather Mountain window. Contours show the altitudes of the
stream channels. Geology from Reed (1964) and Bryant (1965). Letters M, A, B, C, D, and E refer to localities discussed in the text.

   
 

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SOUTHERN BLUE RIDGE PROVINCE

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44 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

course across what is now the North Fork of the
Catawaba River and joined the North Toe near locality
E (fig. 28). Wright understood the main difficulty with
this idea: the valley of the North Fork of the Catawba
River has gentler slopes, or, as he put it, is more mature
than the valley of the Linville River. He also recognized
the speculative nature of the idea. His preferred
hypothesis was an attempt to fit the facts into a
framework of ideas in which the Blue Ridge escarpment
is retreating by a series of piracies and in which the main
streams, at altitudes between 760 and 1,060 m, are flow-
ing in valleys that, as he believed, still contain remnants
of a valley floor or Harrisburg peneplain. One can truly
admire Wright’s careful observation. His conclusion was
almost inescapable within the limitations of the
theoretical concepts of the time. It seems more likely,
however, that the Linville River might at some future
time be captured at locality A (fig. 28) by a tributary of
the North Toe or at locality B by a tributary of either the
North Toe or the North Fork of the Catawba River.

The course of the Linville River is peculiar because it
is on the most resistant rocks of the area. It is best ex-
plained by superposition that took place when the
topographic surface was hundreds of meters higher than
it is now. The original course may have been established
in the Blue Ridge thrust sheet above the Linville Falls
fault plane. The fault plain dips northwest, and the river
probably migrated downdip in a southwesterly direction
within the least resistant rocks. Penetration of the
quartzite probably first took place at some point
downstream near the Piedmont. As the base of the
quartzite became exposed by continued downwasting,
the lower river established a course in the Wilson Creek
Gneiss. The Linville Falls and steep reach below are
presumably working headward and will continue to do
so unless the river is captured by another stream such as
the North Fork of the Catawba or the North Toe.
However, because of the complexity of the geology, it is
impossible to reconstruct with any certainty the factors
that influenced the changes that could have taken place
at higher erosion levels.

The important conclusion, however, is that the Lin-
ville River did not enlarge its drainage basin by capture
of a stream on the northwest side of the divide unless it
did so at a much higher erosion level. Thus, the drainage
divide in this area has probably been near its present
position for some time during a period involving hun-
dreds of meters of downwasting. Migration of the
escarpment is restrained by the extremely resistant
rocks on the face of the scarp. -

EVOLUTION OF THE ESCARPMENT

The present escarpment is clearly the steeper side of a
drainage divide between stream systems draining to the

 

Piedmont and those draining to the Applachian Valley
and Plateaus. Although a simple explanation may not be
possible, the asymmetry is caused in part by massive
and resistant rocks far downstream in the northwester-
ly drainage system. These belts of hard rock are respon-
sible for the long and steep reaches in the Blue Ridge
and Appalachian Valley, as well as the Appalachian
Plateaus. They help to maintain the highland northwest
of the divide. Differences in the rocks along the escarp-
ment itself cannot be the cause of the steep slopes on the
southeast, because in one segment, the escarpment foot
and crest are within nearly homogeneous rocks for more
than 150 km. Local variations in the rock types, for ex-
ample, in the Grandfather Mountain window, may cause
the escarpment to be steeper and higher in some areas
than in others.

The data show that parts of the escarpment crest are
retreating toward the Mississippi River drainage, as
proved by the Dan River capture. This capture took
place in the Alligator Back Formation, which the escarp-
ment follows for 150 km. The evidence in the Linville
River area, on the other hand, shows that a capture
probably could not have taken place there while the Lin-
ville River was lowering its valley through the Grand-
father Mountain window, a vertical distance of 500 to
1,000 m.

By using Ahnert’s (1970) formula for the relation of
the mean relief to the mean denudation rate in areas
20x20 km, a crude estimate can be made of the time
that it would take to lower the Linville River to its pres-
ent position from a position above the window. If a mean
relief of 500 m is assumed, it would take 7.5 my. to
lower the landscape 500 m and 15 my. to lower it 1,000
m. If the relief were higher when the river began to
erode the rocks within the window, t time might be
somewhat shortened. Presumably, then, the drainage
divide in the Linville River area has been near its pres-
ent position or at least within the Grandfather Mountain
window since early Pliocene or late Miocene time.

It is generally well known that thesma'or drainage
divide within the 'uthern Appalachian hiJghlands cor-
responds closely t? a topographic high and a negative
gravity anomaly (p. 6). This coincidence suggests a long
period of stability for the divide and suggests that a
hypothesis of escarpment evolution should be considered
that does not involve substantial migration of the divide.

The configuration of the southeastern margin of the
Blue Ridge from the Roanoke south into Georgia (fig.
25A) suggests such a hypothesis. In figure 25A, the ma-
jor valleys on the escarpment are shown. Many of them
are structurally controlled, especially in the area
southwest of Mount Mitchell, where the escarpment
crest becomes an interior divide. Many valleys have
easterly trends. The Catawba River follows an easterly
trending lineament parallel to the Laurel Creek trench

 

 

SOUTHERN BLUE RIDGE PROVINCE 45

valley, but offset from that valley by the Brevard zone
(fig. 22). The Yadkin River follows the Brevard zone.
Many small valleys follow structural trends in the outly-
ing hills south of the escarpment, such as the Brushy
Mountains. These facts suggest a long history of land-
scape evolution in which the drainage became estab-
lished along major, structural trends, including shear
zones and joint sets. The process did not necessarily in-
volve substantial migration of the continental divide,
because it took place southeast of that divide. The head-
waters of the Green River follow structural trends and
show the process at an early stage. As the stream slopes
are lowered to match the gradient of the Piedmont, the
highlands on either side recede, leaving residual hills or
mountains. Migration of the divide may also take place,
but it depends on the gradients and relative resistance
of the rocks on either side of the divide.

North of the Yadkin River, the divide is migrating
northward in the weak rocks of the Alligator Back For-
mation. Migration is also going on in the headwaters of
the Roanoke River, where the divide crosses the Ap-
palachian Valley. In that area, the migration is rapid in
the limestones.

The South Mountains and Brushy Mountains and
other hills south of the Yadkin and Catawba Rivers are
true residuals or inselbergs formed by the backwasting
of steep slopes, as discussed by Kesel (1974). Recent
mapping by Goldsmith, Milton, and Wilson (1978) at
1:250,000 scale shows that the boundaries of the high-
relief areas cut across the geology, although there is a
tendency for the high areas to be in rocks such as
granite, sillimanite schist, and Henderson Gneiss. The
mountain areas south of the Hendersonville bulge are in
similar rocks, and the margin of this high area cuts
across the geology, an indication that it too has been
backwasting.

Divide migration on a large scale need not have taken
place. The general form of the topography was deter—
mined at a much higher erosional level, possibly as long
ago as the middle Mesozoic. The crest of the Ap-
palachian highlands was probably not too far from its
present geographic position, that is, trending northward
from northeast Georgia to western Pennsylvania.
Whatever the initial drainage pattern, it became ad—
justed to the weak rocks, especially in the limestone and
shale of the Appalachian Valley, probably in a manner
analogous to the drainage evolution of the Zagros Moun-
tains, as described by Oberlander (1965). In the
crystalline rocks, the controls of migration and head-
ward erosion leading to stream captures may have been
largely determined by structural features inherited from
late Paleozoic or early Mesozoic diastrophism. Locally,
however, migration of the continental divide as well as
of many lesser divides is still going on.

 

Another problem is the sharp change in form between
the escarpment foot and the Piedmont, where within a
short distance, the relief may vary within a factor of 5 to
10. The residual outlying hills have similar abrupt
changes in slope at the foot. In some places, the changes
are within the same rock unit. This phenomenon pro-
bably is explained by the development of saprolite,
which can be retained on a low-relief landscape more
readily than it can on one having steep slopes. A
threshold value for local relief and slope probably exists
that determines the retention of a thick blanket of
saprolite. When this threshold is reached, as for exam-
ple, along a retreating escarpment, a sharp change in
relief would take place.

TIME FRAME OF LATE TECTONISM

An important problem that has not been resolved is
the time frame of the late tectonic activity. Unfortunate-
ly, stratigraphic information is available in only a few
areas adjacent to the Coastal Plain, and the history of
the Blue Ridge can be inferred only by an indirect and
tenuous argument. Two areas are considered here. One
is the Northeastern Highlands, where we have the most
convincing geomorphic evidence of renewed uplift, as
well as some stratigraphic evidence. The other area is
the Southern Blue Ridge province, where we have only
geomorphic evidence to go on.

The southern part of the Northeastern Highlands is
marginal to the Coastal Plain at a place where a fault
system has been carefully studied, between
Fredericksburg, Va., and Washington, DC. (Mixon and
Newell, 1977). A monocline and a zone of en echelon
faults lie directly west of the Potomac River estuary
along the reach where the river trends parallel to the
Coastal Plain boundary. The age of the faulting is
Cretaceous and middle Tertiary, but later faulting is
possible. The sense of movement is up on the northwest
side of the fault system. The Brandywine fault system is
southeast of the river and parallel to the Stafford
system. The sense of motion on this fault is up to the
southeast. The Potomac estuary is bracketed by the two
fault systems and is within the downthrown block; thus,
the southwesterly course of the river was probably
determined by these fault systems. We also know that
the original course of the river was toward the southeast
when it flowed in sediments of Miocene age (Schlee,
1957). Thus, the diversion must have taken place in
Miocene time or later.

Other faults along the strike of the Stafford fault zone
cut low gravel terraces in the Washington, D.C., area
and are probably as young as Pleistocene. Therefore,
this region was subject to tectonic stress until very
recently. The landforms show clear evidence that the

46

Northeastern Highlands are, or have been recently, in- '

creasing in relief. The tectonic uplift or tilting respon-
sible for the rejuvenation must have taken place in mid-
dle to late Tertiary and possibly Quaternary time. If
rates of erosion of 20 to 30 m per million years are
assumed (Hack, 1980), the geomorphic evidence is
unlikely to have survived unless some uplift continued
into the Pleistocene. One can speculate that, after an
earlier period of uplift in the Cretaceous, the
topographic relief became subdued. A second uplift of
the area started in late Tertiary time, producing the
present forms, including the convex lower reaches of the
Potomac and Susquehanna Rivers, but leaving gentle
slopes in the upland areas, such as those in the Patapsco
River basin.

Because of the spectacular difference in relief between
the southern Piedmont Lowlands and the Southern Blue
Ridge province, it is of interest to speculate on the tec-
tonic history of that general area. The average relief in
the Blue Ridge highlands is at least 5 to 10 times greater
than the average relief in the adjacent Piedmont. Such
an abrupt difference in relief could not be the result of
lithologic differences between the two regions, because
the physiographic and topographic boundary between
the two does not correspond to the boundary between
the Piedmont rocks and the Blue Ridge rocks. Further-
more, an assessment of the common rock types in the
region does not suggest that great differences in their
resistance can be the cause of the different relief. The
example of the different topographic expression of
granite and other rocks along the Blue Ridge escarp-
ment and the outlying hills shows that the same rocks do
have greatly different topographic expression Within
short distances.

Little empirical evidence of a stratigraphic nature is
available related to late tectonism in the Blue Ridge.
The Brevard zone is probably the youngest major fault
zone in the region, and it is believed to be Paleozoic
(Hatcher, 1978), or it may have existed for a long span of
time, the latest motion being in the Triassic Period
(Bryant and Reed, 1970). A diabase dike crosses the
Brevard zone in the Grandfather Mountain window
without offset. It was intruded into cool rocks, as chilled
contacts show, and is not metamorphosed. It is believed
to be Late Triassic but could be as young as Early
Cretaceous (Bryant and Reed, 1970, p. 161). Several of
the large trench valleys in the Blue Ridge highlands end
at the Brevard zone. Presumably, they are deeply
eroded brittle-fracture or shear zones related to the
Brevard zone. The age of the shear zones is probably
also Paleozoic or early to middle Mesozoic. At least one
silicified breccia zone in the Piedmont near the foot of
the Blue Ridge escarpment offsets a diabase dike. It in-
dicates that some kind of tectonic movement that broke

PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

the rocks probably took place as late as middle Mesozoic.
Although no direct evidence of younger tectonism has
been identified in this region, differential movement
associated with regional warping of the crust could have
taken place along existing fracture zones. Small offsets
would be difficult to detect.

Outlying hills in the North Carolina Foothill zone that
are residual from a more extensive upland suggest that
the Southern Blue Ridge province owes its present relief
to vertical uplift younger than the features discussed
above. It probably belongs to the same erosional and tec—
tonic block as the Valley and Ridge province to the
northwest. This is suggested by the deep valleys on
weak rocks along its northwest margin that are similar
to those in the Valley and Ridge province. The Murphy
Marble belt in the Mountain highlands suggests the
same, because it is eroded hundreds of meters below the
peaks on either side.

Ahnert’s (1970) work on the relation of denudation to
relief provides some basis for speculation. It suggests
that if the only uplift is isostatic, then the mean relief of
any terrain will probably be reduced to 10 percent of its
original value at the end of a 30-m.y. period. Thus, if the
Blue Ridge has been absolutely stable and the present
mean relief is approximately 300m, the area must have
had a relief of 3,000 m in Oligocene time and 30,000 m in
Paleocene time. That such relief is virtually impossible,
supports the idea that later tectonic uplift was involved.

In conclusion, the evidence indicates that the major
features of the topography of the crystalline rock area
are related to differential uplift of very late geologic age
that has involved both the Piedmont and Blue Ridge.
The area northwest of the Tallapoosa-Rappahannock
line has been uplifted more or at a later time than the
Piedmont Lowlands southeast of the line. The tectonic
histories of both areas, however, have many complex-
ities. In general, the area in which relief is highest and
erosion most rapid shows the effects of rock control to
the greatest degree. The present drainage divide may
not be very far from its position at the end of
postorogenic time, that is, early to middle Mesozoic.
Nevertheless, some migration has taken place locally in
the southern Blue Ridge area and probably somewhat
more in the Appalachian Valley.

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w U.S. GOVERNMENT PRINTING OFFICE: 1982— 361-6l4/H7