Physiographic Divisions and 

Differential Uplift in the 

Piedmont and Blue Ridge

G£b :LOG-! A-L SURVEY P R O FC S SI ON A L PAPER

Physiographic Divisions and 

Differential Uplift in the 

Piedmont and Blue Ridge

ByJOHNT. HACK

GEOLOGICAL SURVEY PROFESSIONAL PAPER 1265

A study 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
UNITED STATES DEPARTMENT OF THE INTERIOR 

JAMES G. WATT, Secretary

GEOLOGICAL SURVEY 

Dallas L. Peck, Director

Library of Congress Cataloging in Publication Data

Hack, John Tilton, 1913-

Physiographic divisions and differential uplift in the Piedmont and Blue Ridge.

(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.

II. Series. 

GB124.P53H33 551.4'3'0975 82-600079

For sale by the Superintendent of Documents, U.S. Government Printing Office

Washington, DC 20402
CONTENTS

Abstract _______

Introduction ____

Plan of report _ 

Regional setting _

Page

1

1

2

6

10

11

Piedmont Lowlands ___________________________________

Stream patterns and evidence for sedimentary overlap ____

Quantitative analysis of drainage patterns _________ 12 

Regional physiographic features related to the Pine

Mountain belt ______________________ 14 

Differential movement along the Fall Zone ________ 14 

Northeastern Highlands _____________________ 15 

Southwestern Highlands ____________________ 20 

Foothill zone ___________________________ 23 

Northern Blue Ridge Mountains ________________ 25

Southern Blue Ridge province _. 

Chilhowee-Walden Creek belt.

.____________ 27

._____________ 28

Mountain highlands ______________ 28

Brevard zone and Hendersonville bulge __ 30

Intermontane basins ________ 32

Mount Rogers area _____________ 34

New River plateau _____ _ 34

Blue Ridge escarpment ______________ 35

Dan River capture ________________ 38

Linville River ________________ 40

Evolution of the escarpment ______ 44

Time frame of late tectonism _________ 45

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

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

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

26. Map of the headwaters area of the Dan River _____________________________________________ 39

27. Profile of the headwaters of the Dan River _______________________________________________ 40

28. Map of the Linville River and surrounding areas ______________ ____________________________ 42

29. Profile of the Linville River _______________________________________________________ 43

TABLES

	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

ill

PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT 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 Pennsylvania. 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 effect, 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 resistant rocks. Stream patterns as well as broad topographic forms indicate 

that although the southern part of the Piedmont Lowlands was probably once covered by younger sediments, this area has been exposed to 

erosion for a long time. In North Carolina, the inner part of the Piedmont 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, controlled 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 increased 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 

600m.

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 more 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 surface.

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 differences 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 differences 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 northwest. At the Grandfather Mountain window and south of it, the escarpment 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 suggested by Ahnerf s concept relating rates of relief reduction, rates of 

uplift, and rates of erosion. His analysis shows that if uplift is not continuous, any area such as the Blue Ridge would be reduced to 10 percent 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
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 interconnecting 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 significant. 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, I960). 

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 convinced 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 differences 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 acknowledge 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 effect 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 (Yatsu, 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. Furthermore, 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 tectonic effects and histories. Stream patterns may reflect 

stages during downwasting. Stream-profile analysis 

proves to be a powerful tool with which to identify possible 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 tectonism.

The writer is indebted to many colleagues in the U.S. 

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 distinctly 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 Appalachian 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

R i Richmond basin

Fa Farmville basin

Da Danville basin

DS Durham-Sanford basin

Wa Wadesboro basin

200 300 KILOMETERS

EXPLANATION

TALLAPOOSA-RAPPAHANNOCK LINE 

SUBDIVISIONS OF PIEDMONT LOWLANDS

CA Carolina slate belt 

UW Uwharrie Mountains

OP Outer Piedmont of North Carolina and Virginia 

KM Kings Mountain belt 

PM Pine Mountain belt

FIGURE l.-Map of Southeastern United States showing the subprovinces within the Piedmont and Blue Ridge and the TallapoosaRappahannock Line.
PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

76° X"

EXPLANATION

Crystalline basement

Clastic and volcanic rocks of low to medium

4 Mostly highly deformed metasedimentary

Metasedimentary rocks, chlorite-schist, 

slate, and volcanic rocks

300 KILOMETERS

32

Metasedimentary rocks, migmatites, and 

quartz-kyanite rocks

Relatively undeformed sedimentary and 

volcanic rocks (Carolina volcanic slate belt)

Plutonic rocks

Mainly quartz-feldspar gneiss and schist

Cataclastic rocks

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

JQ!________Jf

32'

100 

I

200 

I

300 

I

Contour interval is 150 m between 0 and 450 

Contour interval is 152 m between 0 and -1520

400 KILOMETERS 

I

EXPLANATION

450-850 m 850-1250 m Over 1250 m Areas in the Coastal Plain underlain 

by Miocene and younger deposits

NEGATIVE CONTOURS NE OF SAVANNAH RIVER ARE ON BASEMENT. NEGATIVE CONTOURS SW OF SAVANNAH

RIVER ARE ON BASE OF BEDS OF AUSTIN AGE

FIGURE 3. -Averaged topographic map of part of 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, and the Appalachian Plateaus. The contours on the land surface 

show mean altitudes within 3-minute rectangles and were prepared from digitized data (original contour map furnished by M. F. 

Kane). Generalized contours on the basement are from the "Tectonic map of the United States" (U.S. Geological Survey and American 

Association of Petroleum Geologists, 1961).
6 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

ly 6x5 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 l: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 between 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 example 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 Piedmont 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 1:1,000,000 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 vertical 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 Pennsylvania may have been formed in early Mesozoic time, 

as it is related to a low-gravity anomaly and also corresponds 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 individual 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) appears to be a major dividing line in the region. It is expressed 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 intersects 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 corresponds 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 embayment, however, the rise of the subaerial topographic surface 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 embayment collects the drainage of some of the largest 

rivers in the Appalachians, including the Delaware, Susquehanna, Potomac, James, and Roanoke. This embayment and its extension to the south in North Carolina 

received great thicknesses of sediment, particularly in 

Cretaceous and Miocene time (Owens, 1970). In late Tertiary 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 tectonic system that formed the Salisbury embayment. The 

embayment and a possible troughlike area to landward 

may have been responsible for gathering a fanlike network of large rivers.
REGIONAL SETTING

300 KILOMETERS 

I

EXPLANATION

Steep hills and mountains 

(relief more than 50 m in 

6-km2 area)

Low relief

(relief averages 60 m in 

100-km2 area)

Moderate relief 

( relief averages 150 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

80°

88°

300 

I

400 KILOMETERS

EXPLANATION

ATLANTIC-GULF DIVIDE

TALLAPOOSA-RAPPAHANNOCK LINE

FIGURES 5.-Drainage map of part of the Appalachian highlands (shaded area), including the Piedmont and Blue Ridge. Rivers are identified 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, Ocmulgee; 20, Flint; 21, Chattahoochee; 22, Tallapoosa; 23, Coosa; 24, New; 25, French Broad; 26, 

Tuckasegee; 27, Hiwassee; 28, Etowah; 29, Tennessee.
REGIONAL SETTING

.0 CO

o
10 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

B

METERS 

2000-1

Basement 

rocks

B 1

1800-

1600-

1400-

1200-

1000-

800-

600-

400-

200-

Ashe Formation

Appalachian 

Valley

Basement rocks

Grandfather Mountain window

Morganton basin

South Mountains

50 KILOMETERS 

I

Vertical exaggeration X38.5

FIGURE 7.-Topographic profile across the Blue Ridge from the Appalachian Valley near Johnson City, Tenn., to South Mountains south of

Morganton, N.C. 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 Atlantic coast (see rivers 13 to 16, fig. 5). If the basins of these 

rivers were at one time gathered together in a troughlike 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 sedimentary 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 lowrelief 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
PIEDMONT LOWLANDS 11

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 pronounced 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 finegrained 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 affected by the Warm Springs fault (White, 1965; 

Christopher and others, 1980). This occurrence of finegrained 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 deformation 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 spectacular; 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) contain 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 continental 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 contrast 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 patterns 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 major 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 secondorder 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 northwest 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 extended through South Carolina because the outer Piedmont 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 consequent 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 northeasttrending 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 narrower 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 described 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. Considerable 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 restrictions on basin shape, basins 100 km2 in area on the 

average have a width-length ratio of 0.33, whereas 

basins 10,000 km2 in area on the average have a widthlength 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 km2) in area; stream length 

was found on the average to approximate the function 

L = 1.4 A°-6 , 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 A0-6 . 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
PIEDMONT LOWLANDS 

TABLE 1. -Mensuration 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]

13

Basin Length 

(km)

Area 

(km*)

Average 

width 

(km)

Predicted 

length-law 

width

Difference 

(percent)

Width-length 

ratio

Piedmont south of Cape Fear River

Deep-Cape Fear, N.C.

Yadkin, N.C. _

Catawba, N.C.-S.C.

Broad, S.C. _

Saluda, S.C.

Savannah, Ga. _

Oconee, Ga.

Ocmulgee, Ga. _

Flint, Ga.

136

. _ 317

279

. _ _ 241

226

233

. _ 178

142

. _ 137

9,680 

18,663 

11,691 

13,263 

6,612 

18,277 

7,337 

6,128 

5,239

71 

58 

42 

55 

29 

78 

41 

43 

38

40 

53 

43 

46 

34 

53 

36 

33 

31

+ 77 

+ 9 

2

+ 19 

-14

+ 47 

+ 14 

+ 30 

+ 23

0.52 

.19 

.15 

.23 

.13 

.54 

.23 

.30 

.28

Outer Piedmont north of Cape Fear River

Po River, Va. _

North Anna, Va.

Little River, Va.

South Anna, Va.

Appomattox, Va.

Nottoway, Va.

Meherrin, Va.

Fishing Creek, N.C.

Swift Creek, N.C.

Tar, N.C. _

Neuse, N.C.

36

71

. _ 38

. _ _ 90

162

. _ _ 70

. _ 86

64

. _ 61

. _ _ 114

142

297 

1,052 

281 

1,113 

3,387 

1,193 

1,797 

1,161 

464 

1,793 

4,430

8.2 

15 

7.3 

12 

21 

17 

21 

7.6 

18 

16 

31

9.2 

16 

8.9 

16 

26 

16 

20 

11 

16 

20 

29

-11

6

-18 

-25 

19

+ 6 

+ 5 

-30 

+ 12 

20

+ 7

0.23 

.21 

.19 

.13 

.13 

.24 

.24 

.12 

.28 

.14 

.22

Coastal Plain

Edisto, S.C. _

Salkahatchie, S.C.

Ogeechee, Ga.

. _ _ 211

129

267

7,457 

3,129 

7,684

35 

24 

29

36 

25 

36

2

4

19

0.17 

.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; Mosely 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 A0-61 , expressed 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 A°-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 

andisW=0.82A°-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 explained, 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 Tertiary time, as studies in the Washington, B.C., area suggest (Barton, 1951). The area of overlap by upper Tertiary Coastal Plain sediments corresponds to a steep 

gradient of the submerged and buried crystalline basement and is associated with the deep Salisbury embayment (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 substantially 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 Piedmont. 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 gradients 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 resistant quartz rock within these steep reaches that would increase 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 km2 at 

the Fall Zone, compared with 5,240 km2 for the Flint. 

One would expect that the gradient of the Chattahoochee 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 neighboring 

streams to the east, one sees little similarity. The Ocmulgee and Oconee show a concave profile as they descend 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 Mountain 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 exceptions, 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'Connor, 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, compared 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, Appomattox, 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. (Bischinger, 1979). The Rappahannock River
NORTHEASTERN HIGHLANDS 15

RAPPAHANNOCK RIVER SALUDA-CONGAREE RIVER SYSTEM

SAVANNAH RIVER 

Fz

APPOMATTOX RIVER

OCONEE RIVER

ROANOKE RIVER

DEEP-CAPE FEAR RIVER SYSTEM 

Fz

OCMULGEE RIVER

YADKIN-PEEDEE RIVER SYSTEM Apalachicola River

CHATTAHOOCHEE RIVER

CATAWBA-WATEREE-SANTEE RIVER SYSTEM

Apalachicola River

200 400 600 800 

DISTANCE FROM SOURCE, IN KILOMETERS

0 200 400 600 800 

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 Newell, 1977).

NORTHEASTERN HIGHLANDS

The Northeastern Highlands is an area of substantially 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 Appalachian Valley. The highlands include a strip of interconnected 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 comparatively high. It was recognized by Campbell (1929, 

1933) as distinctly different in relief from the surrounding Piedmont; he referred to it as the Westminster anticline, 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 Stafford 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 km2 , whereas, in much of the Piedmont, 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 adjacent highlands. For example, in the Frenchtown 

quadrangle of the Newark basin (Drake and others, 

1961) a 100-km2 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 

m 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 longitudinal 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,

DELAWARE RIVER

:all Zone

SCHUYLKILL RIVER 

/Fall Zone

SUSQUEHANNA RIVER

Zone

PATAPSCO RIVER

Fall Zone

POTOMAC RIVER

0 200 400 600 800 

DISTANCE FROM SOURCE, IN KILOMETERS

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., 78 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 Piedmont. 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-
NORTHEASTERN HIGHLANDS 17

nel is -15 m. The average slope in this reach is 0.00095, 

somewhat steeper than the slope upstream. The opposite 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 Rappahannock River (fig. 8). These data seem to indicate

that the Susquehanna, Patapsco, Potomac, and Rappahannock Rivers were all affected by differential movement 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

Westminster

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 normal 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. Except 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 median 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). 

Median size 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

200

CO

cc tu

tu

tu 

Q

D

100

North Branch

0 1.0

Reservoir

PATAPSCO RIVER

Baltimore 

Complex

10 

DISTANCE FROM SOURCE, IN KILOMETERS

FIGURE 11.-Logarithmic profile of the Patapsco River, Maryland.

100
NORTHEASTERN HIGHLANDS 19

200-

DC 

UJ

H 

UJ

100

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 characteristic 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 resistant 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 northwest, 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.
20 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

-260

^

36m

| 122m . 

GILLIS FALLS ABOUT 3.4 KILOMETERS FROM SOURCE

58 m

91 m I 

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 underlain 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 clastic rocks of lower 

metamorphic grade, many of them fine grained and interbedded 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 

th6 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).

180

DISTANCE FROM TRUNK STREAM, IN METERS

FIGURE 14.-Profiles of some first-order tributaries in the Patapsco 

River basin. For locations, see figure 10.
SOUTHWESTERN HIGHLANDS 21

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 into 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 northeast (A, fig. 16). This escarpment corresponds to the 

drainage divide separating the headwaters of the 

Tallapoosa River from small tributaries of the Chattahoochee 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'

33°

34'

33'

PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

87° 86° 85° 84°

CONTOUR INTERVAL 122 m 

EXPLANATION

:*&;::l Areas higher than 488 m

EXPLANATION

ms Metasedimentary rocks 

v Volcanic rocks

9 Graphitic schist and phyllite

ga Garnet and graphitic schist

f Feldspathic schist and gneiss

Granitic rocks 

bg Biotite gneiss 

mg Metagraywacke and quartzite 

bv Brevard zone

50 100 KILOMETERS

FIGURE 16.-Simplified topographic and geologic maps of the Southwestern Highlands. Geologic data from Neathery and others

(1976).
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 U.S. Army Corps of 

Engineers show, in the lower 60 km of the Piedmont, 

the river was intrenched as deeply as 130 m before impoundment by a series of dams. The area of the drainage 

basin of this river is about 2,600 km2 . 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, N.C. 

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 (U.S. Army Corps of Engineers

400 

<£ 300

LLJ

uj 200 

Z 100

S °

ID 200

TALLAPOOSA RIVER 

Fall Zone 

Joins Coosa

- COOSA RIVER

\- 

<

100

Enter Piedmont

Fall Zone

. Joins 

[Tallapoosa

0 200 400 600 800 

DISTANCE FROM SOURCE, IN 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 Appalachian Valley (fig. 15). The topography in the Appalachian 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 Piedmont to near the South Carolina-North Carolina boundary. Its boundary on the northwest is defined by the 

southeast margin of the Blue Ridge physiographic province (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, consisting 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 contain resistant beds that underlie ranges of hills. These 

rocks are commonly quartzites of the Mount Athos Formation and of the Candler Formation, which also contains 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 exposed in a syncline (Alien, 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 anticlinorium 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

c/) DC

LLJ hUl

Q

13

TALLAPOOSA RIVER

200

100

Join Coosa River

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 Alien (1963) (James 

Clarke, oral commun., 1978). Thus, in this area, the effect 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 (U.S. 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

elude (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 

silts tone.

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, hornblende gneiss with amphibolite, and schist. Even the 

sillimanite-rich rocks are not restricted to the mountains. 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 Mountains rise 450 m above the Morganton Basin. The local 

relief within the mountains averages about 100 m, compared 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 gradient 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 characteristic 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 profile. 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 tectonic 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 farther 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 includes only the escarpment and mountains behind it that 

rise directly above the Piedmont. This common definition excludes, for example, the Great Smoky Mountains 

as well as some other ranges southeast of the Appalachian Valley. Physiographers have long used the 

term "Blue Ridge" as a name for a physiographic province. 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 probable 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 Piedmont 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 Mountain 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 referred 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 Precambrian and Cambrian rocks that form the northwest limb 

of the Blue Ridge anticlinorium of the geologists (Espenshade, 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

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 Precambrian 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; 

Alien, 1963). The "Pedlar Formation" is generally more 

resistant than other rock units in Brown's Virginia Blue

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BLUE RIDGE ANTICTLJMQRIUM

EXPLANATION

Chilhowee Group 

(clastic rocks)

Catoctin Formation 

(metabasalt)

Rwl

Metarhyolite of the 

Catoctin Formation

Swift Run Formation Virginia Blue Ridge

(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.
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 considerable 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 Mountains (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 Proterozoic 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 Formation in the south), a sequence of metamorphosed shale 

and siltstone and minor sandstone and quartzite; 

thickness ranges from 300 to 600 m. The topmost formation in the group is the Antietam Formation (Erwin Formation in the south), composed of siliceous sandstone 

and vitreous quartzite, 125 to 250 m thick. These formations generally dip steeply to the northwest and form 

the northwest flank of the Northern Blue Ridge Mountains. 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" Formations.

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 tectonic 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 controlled 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 interrupted 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. Contrasts 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 Formation. 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

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 m. 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 exceed 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 Mountain 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 example, 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 im35'

100 200 KILOMETERS

FIGURE 20.-Map of the Southern Blue Ridge province, showing the subdivisions discussed in the text. Ruled line pattern represents highrelief 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.
SOUTHERN BLUE RIDGE PROVINCE 29

pression given by the region does not suggest rock control 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 suggests an origin due in part to control by nonresistant 

rocks and in part to tectonic influences. The tectonic influences 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 correlations 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 Marble and Andrews Schist of Cambrian(?) age; these formations are the youngest in a great syncline folded into 

the rocks of the southwestern part of the Mountain highlands (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 spectacular 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 headwater 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 Appalachian 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, Chattanooga, and Rome sheets of the l:250,000-scale 

topographic-map series distributed by the U.S. 

Geological Survey.

Another spectacular trench is north of Asheville, 

N.C., 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 Marble. On the other hand, the presence of the fault suggests 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 contrast with basins having relief of less than about 200 m is 

great. Figure 22 is an outline map of part of the Mountain 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 central 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 fracture (or shear) zone (Pickering and Higgins, 1979; 

Georgia Geological Survey, 1976). Some other trench 

valleys are identified by names coined by the writer.
30

sw

PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE 

VALLEY RIVER NANTAHALA RIVER

1500

1250

co 1000

DC 

01

750

500

250 1-

5000

4000

3000

2000

1000

hLU 

LU

TRENCH VALLEY ON MURPHY MARBLE

CANE 

RIVER

LITTLE CRABTREE 

CREEK

A

1500

1250

COcciOOO

| 750

500

250

5000

4000

3000

2000

1000

LAUREL CREEK TRENCH VALLEY

30 KILOMETERS 

I

FIGURE 21.-Profiles along two trench valleys in the Mountain highlands. The wavy heavy lines represent the profile along ridge crests immediately 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 resemble 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, N.C., 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
SOUTHERN BLUE RIDGE PROVINCE 31

^^rSs^SSi^.- i'V

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).
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 into the Piedmont from the main Mountain highlands. For 

convenience of discussion, it is here called the "Hendersonville 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 Hendersonville bulge and the Mountain highlands on the northwest side of the Brevard zone must have another cause. 

The rocks of the bulge are considered part of the Piedmont province by geologists. The most extensive formation 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 maximum 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 Appalachian 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 lineament. 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 continue westward as the Bryson City trench.

The Tuckasegee trench trends northwestward from 

the Brevard zone. This crosscutting trench is noteworthy 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 vertical Mesozoic dike at least 175 m. The amount of vertical offset, if any, cannot be determined. This evidence 

indicates that fracture zones do exist in the area that involved 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 affected the gross morphology of the terrain.

INTERMONTANE 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 Thornbury, 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 

^^\

33

^\

20 30 40 

I

50 KILOMETERS

FIGURE 23. -Geologic 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; pGg, Max. Patch 

Granite and Cranberry Gneiss and migmatite; p6gn, layered gneiss and migmatite; pGgs, Great Smoky Group; hg, 

Henderson Gneiss and related rocks; pGgc, biotite schist and gneiss. (From Hadley and Nelson, 1971.)
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 Canton (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 Hendersonville, 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 symmetry 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 surface. The data seem to indicate only that the relief increases downstream and that the local topography is 

symmetrical and adjusted to the present river. The increase 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 effect.

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 intensive faulting and shearing along many zones. The absolute amount of movement in these zones may have 

been small, but the zones provided a network of fractures 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 determined 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 ARIA

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 surroundings. 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 differences in relief.

NEW RIVER PIATEAU

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 1:24,000 scale; 20-ft contours]

Quadrangle

Hendersonville

Asheville

Leicester

Marshall

Highest point 

(m)

664

670

670

701

Lowest point 

(river surface) 

(m)

633.5

600

524

505

Total relief1 

(m)

30.5

70

146

195

Mean relief 

(m)

12.8

30.5

84.0

100

Mean altitude 

(m)

646

631

608

605

Elevation-relief 

ratio2

0.42

.43

.57

.51

Difference in altitude between highest and lowest points.

Ratio of mean relief to total relief; essentially the same as the hypsometric integral (see Pike and Wilson, 1971).
SOUTHERN BLUE RIDGE PROVINCE 35

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 Formation 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 northeastern end. The Piedmont surface as well as the escarpment 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 section 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

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 impressive. 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 Mexico, 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 Appalachian Valley near Blacksburg, Va., forming an irregular 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 escarpment 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 Campbell (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, producing 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 introduced, as part of his hypothesis, the idea that the existence of barriers to erosion in the form of resistant 

bedrock on the northwest side of the divide would increase 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 Piedmont and are now upfaulted peneplain surfaces. Stose 

and Stose (1951) criticized the faulting hypothesis, pointing 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

36° __ 84

monoclinal uplift could produce an escarpment maintained 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 interesting 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 consistency 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 Piedmont 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

34'

EXPLANATION

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

36° 82°

37

81 C 37°

( ( U, /

?u 7 Mount ?Linville River

Ashev.Mef. Mitchell \/^

£^ \t> / dv4 ^ 

^^ S£J Brushy Mountains 

Mountains , ____^___

Asheville x' sp 

 Mount Mitchell

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, undifferentiated; 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; ip, 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 escarpment is within the Alligator Back Formation, a weak 

rock consisting of thinly laminated micaceous schist and 

gneiss (Espenshade and others, 1975). The stream profiles on the escarpment itself are highly concave, a 

characteristic of headwater profiles in weak rocks. 

Thus, the aspect of the scarp is impressively steep. On 

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 Appalachian 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 escarpment 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 Formation, 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 northwest 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 difference may be an accident of the local geology. The 

streams flowing toward the Atlantic have penetrated 

relatively resistant rocks such as massive quartz monzonite, granodiorite, and paragneiss along weak but narrow 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 present 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 escarpment 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 inherited 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 upper basin of Reed Island Creek. The geology of the headwater area is relatively simple, and as the entire headwater 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 channels and show in a general way the vertical configuration 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
SOUTHERN BLUE RIDGE PROVINCE

80°30'

39

3 10

I I

20 MILES 

I I

I I I

0 10 20 KILOMETERS

CONTOUR INTERVAL 152.5 m (500 ft) 

FIGURE 26. -Map of the headwaters of the Dan River, showing the area of stream capture. Streams shown by

dashed lines flow toward Atlantic Ocean. Subenvelope contours are drawn on altitudes of the stream channels. 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).
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 sharply 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 discharge, 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 component 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 introduction 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 gradient 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 gradient m. At the steepest reach, the value is 2,900 gradient m. It then drops gradually to 240 gradient m at the

N

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 headward 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 examination 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 erosion.

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 abqut 

55x25 km, on the southeastern margin of the Blue 

Ridge near Morganton, N.C. 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

1000

cc

UJ

t. 500

APPROXIMATE LEVEL OF UPLAND

abg

srs

10

aba

20 30 40 

KILOMETERS

50 60 70

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 approximately 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 northeast. 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 Grandfather Mountain window. The geology along the river is 

shown on the exaggerated profile of figure 29. In its upper course (north of lat 36°), the river is in the Grandfather Mountain Formation of Proterozoic Z age. The 

rocks of this formation, which range from phyllite and 

siltstone to felsitic volcanic rocks and arkosic conglomerate, 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 controlled 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 backwasting 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, N.Y.

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 profile 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 quartzite 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 between 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 satisfactory 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 escarpment. 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°

r \* 4 

v s 3* Jr \ X <»f ^

L S5c^

SJr^TABLEROCK^V ,

IlxFrLT .v^

20 MILES

10 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.
SOUTHERN BLUE RIDGE PROVINCE 43

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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 flowing 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 explained 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 Linville 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 hundreds 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 northwesterly drainage system. These belts of hard rock are responsible 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 escarpment 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 example, 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 escarpment 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 Linville River was lowering its valley through the Grandfather 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 

20 x 20 km, a crude estimate can be made of the time 

that it would take to lower the Linville River to its present position from a position above the window. If a mean 

relief of 500 m is assumed, it would take 7.5 m.y. to 

lower the landscape 500 m and 15 m.y. to lower it 1,000 

m. If the relief were higher when the river began to 

erode the rocks within the window, the time might be 

somewhat shortened. Presumably, then, the drainage 

divide in the Linville River area has been near its present position or at least within the Grandfather Mountain 

window since early Pliocene or late Miocene time.

It is generally well known that the major drainage 

divide within the southern Appalachian highlands corresponds closely to 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 major 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 outlying hills south of the escarpment, such as the Brushy 

Mountains. These facts suggest a long history of landscape evolution in which the drainage became established along major structural trends, including shear 

zones and joint sets. The process did not necessarily involve substantial migration of the continental divide, 

because it took place southeast of that divide. The headwaters 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 Formation. Migration is also going on in the headwaters of 

the Roanoke River, where the divide crosses the Appalachian 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 highrelief 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 determined at a much higher erosional level, possibly as long 

ago as the middle Mesozoic. The crest of the Appalachian 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 adjusted 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 Mountains, as described by Oberlander (1965). In the 

crystalline rocks, the controls of migration and headward 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 probably 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 example, 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. Unfortunately, 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, D.C. (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 PHYSIOGRAPHIC DIVISIONS AND DIFFERENTIAL UPLIFT, PIEDMONT AND BLUE RIDGE

Northeastern Highlands are, or have been recently, increasing in relief. The tectonic uplift or tilting responsible for the rejuvenation must have taken place in middle 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 tectonic 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. Furthermore, 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 escarpment 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 indicates that some kind of tectonic movement that broke

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 tectonic 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 300 m, 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 

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The area northwest of the Tallapoosa-Rappahannock 

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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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