Hydrology and Geochemistry of 

Thermal Springs of the Appalachians

GEOLOGICAL SURVEY PROFESSIONAL PAPER 1044-E

Hydrology and Geochemistry of 

Thermal Springs of the Appalachians

By W.A. HOBBA, JR., D. W. FISHER, F. J. PEARSON, JR., and J. C. CHEMERYS 

GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

GEOLOGICAL SURVEY PROFESSIONAL PAPER 104 4-E

A study of the geologic and hydrologic setting 

and the physical and chemical character 

of selected major warm springs 

of the Appalachians

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON : 1979
UNITED STATES DEPARTMENT OF THE INTERIOR 

CECIL D. ANDRUS, Secretary

GEOLOGICAL SURVEY 

H. William Menard, Director

Main entry under title:

Hydrology and geochemistry of thermal springs of the Appalachians.

(Geohydrology of geo thermal systems) (Geological Survey professional paper; 1044-E)

Bibliography: p.

Supt. of Docs, no.: I 19.16:1044-E

1. Hot springs-Appalachian Mountains. 2. Geochemistry-Appalachian Mountains. 

I. Hobba, W.A. II. Series. III. Series: United States. Geological Survey. Professional 

paper; 1044-E. 

GB1198.3.A57H92 551.2'3'0974 79-607987

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

Washington, D.C. 20402

Stock Number 024-001-03218-0
CONTENTS

Page

Abstract ________________________________ E 1

Introduction ____________________________ 1

Purpose and scope _____________________ 1

Previous studies _______________________ 1

Acknowledgments _____________________ 3

Historical background-____________________ 3

Geologic and hydrologic setting, by W. A. Hobba, Jr ______ 4

General description of the springs _______________ 4

Physical setting _______________________ 4

Geology and topography ________________ 4

Rock structure ____ _____ ______ ____ _ 4

Hydrology ____________________________ 11

Probable flow systems___________________ 11

Selected flow systems __________________ 13

Temperature of thermal waters _ ______ _ _ ____ 17

Fluctuations ________________________ 17

Wells and thermal gradients______________ 18

Heat discharge __________________________ 20

Geochemistry, by D. W. Fisher, F. J. Pearson, Jr., and

J. C. Chemerys ________________________ 20

General principles _____________________ 20

Geochemistry—Continued

Geochemical indicators ___________ ——————— _ E21

Dissolved gases ____———— ———————— ————— 21

Trace elements _____________________ 23

Isotopes ____________________-————— 23

Carbonate chemistry and isotopes ______————— 26

Chemical geothermometers ____———————————— 27

Discussion _______________ ——————— ————— 29

Silica-rock spring systems ___ ______ _____ 29

Perry County, Pa., area _____________ 30

Berkeley Springs, W. Va., area _________ 30

Minnehaha Springs, W. Va., area ________ 30

Warm Springs, Ga., area ____________ 30

Carbonate-rock spring systems______________ 31

Lebanon Springs, N. Y., area ________———— 31

Bolar Spring, Warm Springs, Hot Springs,

and Falling Spring, Va., areas __________ 32

Hot Springs, N. C., area________________ 33

Conclusions ___________________________ 34

References cited _________________________ 35

ILLUSTRATIONS

Page

FIGURE 1. Photograph of Homestead Hotel and Hot Sulphur Spring at Hot Springs, Va_________________________ E 3

2. Satellite images of thermal spring areas in Virginia, West Virginia, Georgia, and New York ________—————— 5

3. Side-looking radar images showing prominent lineaments and thermal springs and wells in Virginia, West Virginia,

and Georgia __________________________________________________——————————— 6

4. Map showing geology, spring, and well locations in Minnehaha Springs area, West Virginia ——————————————— 8

5. Cross section parallel to fault passing near Webster Springs and Minnehaha Springs, W. Va., and

Warm Springs, Va _______________________________________________——————— 9

6. Map showing lineaments and their relationship to thermal springs in Virginia and West Virginia ________ ——— _ 10

7. Photograph of sandstone quarry looking southwest toward Berkeley Springs, W. Va ________——————————— 11

8. Cross section showing theoretical movement of ground water through a homogeneous isotropic aquifer with

an impermeable layer at 1,500 m _____________________________-___—————————————— 11

9. Cross section showing possible movement of ground water through a multilayered folded, faulted, and fractured

aquifer such as those in Warm Springs Valley, Va __________________________———————————— 11

10. Block diagram showing possible circulation of water to a warm spring located near the crest of an anticline at the

intersection of two faults _________________________________________—————__———— 12

11. Diagram showing possible movement of ground water through a faulted and fractured anticlinal ridge____ ————— _ 12

12. Diagram showing possible movement of ground water through a faulted and fractured anticlinal ridge bordered

by a synclinal ridge ___ __ _ ____ ___ ____ __________ ______ _______ ______ _ _ ___ 13

13. Plan view showing section lines for figures 12 and 14 and possible flow lines in the plane of the sandstone

formation ________________________________________________________ 13

14. Diagram showing possible movement of water along a vertical plane parallel to the strike of an anticlinal ridge ——— 13

15. Block diagram showing possible circulation of water to the warm springs at Berkeley Springs, W. Va __________ 14

16. Map showing geology and spring and well locations near Berkeley Springs, W. Va _________________ ——— _ 15

17. Graphs showing water temperature of hot spring and stage of French Broad River at Hot Springs, N.C ______——_ 17

18. Graphs showing temperature and discharge of Bolar Spring at Bolar, Va ________________———— —— _ 18

19. Logs showing temperature for a cold-water well and temperature and fluid conductivity for a warm-water well — —— _ 19

20. Temperature log for air-filled "dry" gas well (Pocahontas 21) near Minnehaha Springs, W. Va _________————— 20

21. Graph showing isotopes in Appalachian warm springs area waters ________________ —————— —— _____ 25

22. Graph showing dissolved silica and temperatures, Appalachian warm spring areas ________________———— 28

III
IV CONTENTS

TABLES

Page

TABLE 1. Thermal gradients at wells producing warm or cold water ______________———————————————————— E19

2. Estimated depths of circulation based on thermal gradients and conduction losses ————— ———— ——————————— — 20

3. Flow, temperature, and power output of thermal springs __ _ _______—— ——— ———— —— ————— ————— —— 21

4. Temperatures and selected analytical data for waters in Appalachian warm-spring areas—— ——— ——— __———————— 22

5. Saturation indexes and carbon-isotope ratios for waters in Appalachian warm-spring areas— —— ————————————— 24
GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS

OF THE APPALACHIANS

By W. A. HOBBA, JR., D. W. FISHER, F. J. PEARSON, JR., 

and J. C. CHEMERYS

ABSTRACT

Thermal springs in nine areas in the Appalachians from Georgia 

to New York were studied in 1975 and 1976 using satellite imagery, 

local well and spring data, and results of current and early studies by 

other investigators. All the springs investigated discharge from 

folded and faulted sandstone or carbonate rocks in valley areas. 

Where geologic structure is relatively uncomplicated, ground water 

discharging from thermal springs probably has circulated to great 

depths roughly parallel to the strike of the bedding and has moved 

upward rapidly where a fault or faults cross the bedding. Hydrologic 

and chemical data suggest that most of the water discharging from 

warm springs in the Devonian Oriskany Sandstone is derived from 

recharge entering and circulating through that formation. The temperature of the thermal springs can be accounted for by "normal" 

geothermal heat flow if water circulates to minimum depths of 250- 

1,600 meters along paths of sufficient length and contact area to 

absorb enough geothermal heat to raise the temperature of the water 

to that required to produce the given thermal springs. The discharge 

at springs where temperature fluctuates very little is primarily 

water from deep circulation. The discharge at springs where temperature fluctuates widely is warm water mixed with variable proportions of shallow-circulating cool water.

Observed temperatures of the warm springs range from 18° to 

41°C; the highest chemical thermometer temperature is 84°C. 

Agreement among observed, chalcedony, and cation temperatures of 

the warmest springs suggests reservoir temperatures of 30°-50°C. 

Dissolved helium, arsenic, potassium, and 818O are considered as 

geothermal indicators. Tritium analyses are used to calculate fractions of old and modern components of mixed waters. Computer calculations of carbonate saturation indices show (1) considerable undersaturation in silica-rock warm spring waters and (2) carbonate 

equilibrium in the limestone and dolomite thermal waters. Better 

values of saturation indices are obtained when analyzed carbon 

dioxide rather than field pH is used in the computer input data. A 

method is described for adjusting 813C to correct for carbon dioxide 

outgassing from water samples.

INTRODUCTION

PURPOSE AND SCOPE

The primary purpose of this study was to investigate 

the geologic and hydrologic setting and the physical 

and chemical character of several of the major warm 

springs in the Appalachians.

Investigative and analytical techniques in 

geochemistry, hydrology, and remote sensing have advanced significantly in recent decades, but few detailed 

studies of the thermal springs of the Appalachians 

were made during this time. By using new techniques 

and investigating the larger thermal springs, an attempt was made to identify common characteristics 

that would show the character of the geothermal systems of the region. The first part of this report, prepared by Hobba, describes the geologic and hydrologic 

setting of the springs; the second part, prepared by 

Fisher, Pearson, and Chemerys, discusses the 

geochemistry of the springs and inferences of thermalreservoir temperatures indicated by the chemical data.

The present study was a 1-year reconnaissance. 

Springs and wells were inventoried and water samples 

collected during the autumn of 1975 and the spring of 

1976. Several cold springs and wells were also sampled 

to provide comparative data from which properties 

peculiar to the warm springs could be discerned. Complete hydrogeological and chemical data and descriptions of the methods of sample collection and 

analyses are given in Hobba, Chemerys, Fisher, and 

Pearson (1977).

PREVIOUS STUDIES

Various aspects of the warm springs in the Appalachian Mountains have been described in numerous scientific papers, some of which date from well back into 

the nineteenth century. In addition to descriptions of 

flow, temperature, and chemical character of the 

springs, many of the papers speculate as to the geologic 

structural setting and the origin and the depths of circulation of the waters. Most of the early inferences to 

depths of circulation of water were based on simplified 

assumptions of prevailing gradients and their effect on 

spring-discharge temperatures. Recent conceptual 

models, such as those of Lowell (1975), provide more

El
E2 GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

realistic estimates of depths of circulation. Likewise, 

many of the earlier interpretations of geologic structure associated with the springs were based on field 

observations. Now, gross structural features may be 

revealed by remote sensing from satellites or highaltitude aircraft. Similarly, geochemical techniques 

were formerly very simple; recently techniques have 

been vastly improved so that water-rock reactions and 

chemical equilibria can be modeled and even age dating of water is possible. From this prespective the following paragraphs summarize the earlier papers most 

useful to the present study.

McCallie (1908) reported the chemical quality and 

discharge of Warm Springs, Ga. In a later paper, 

McCallie (1913, p. 13) inferred a depth of circulation of 

about 450 m on the basis of the discharge temperature, 

the mean annual air temperature, and an assumed 

"normal" geothermal gradient. Because of errors inherent in the simplified assumption, 450 m probably 

represents, at best, only a minimum depth of circulation. In an extensive study of Warm Springs and the 

surrounding area, Hewett and Crickmay (1937, p. 34, 

35) concluded that the water emerging from the 

springs enters a bed in the lower part of the Hollis 

Quartzite on Pine Mountain, moves downdip to a depth 

of about 1,200 m where the bedding is offset by a fault, 

moves upward along the fault, and then moves updip 

along a second permeable bed near the top of the Hollis 

Quartzite to discharge where the bed crops out.

Keith (1904), Stose and Stose (1947), and Oriel 

(1950) generally agreed that the area surrounding Hot 

Springs, N.C., is faulted and may be described as a 

structural window. However, Trapp (1970) suggested 

that the structure is a graben rather than a window 

and stated (p. 28) that "A fault probably forms part of 

the artesian system supplying the thermal springs at 

Hot Springs." Stose and Stose (1947, p. 644) concluded 

that the spring water circulates to a depth of about 

1,500 m and discharges along a nearly vertical fault.

Rogers (1884) and Reeves (1932) describe the warm 

springs of Virginia. Rogers (1884, p. 577-597) observed 

that many of the warm springs are on the axes of anticlines and suggested that water enters the rocks in the 

anticlinal ridges and sinks along joints and fissures 

until it reaches a permeable bed along which it then 

rises to discharge at the outcrop in an adjacent valley. 

Reeves (1932, p. 28) suggested that water enters a 

permeable bed at a relatively high altitude on the crest 

or limb of one anticline and travels along this bed 

through a syncline to an outcrop at a lower altitude in 

another anticline. According to Reeves (1932, p. 26), 

some anticlines are broken by thrust faults, but most of 

the springs in the region are not on faults, and hence 

few are fed by water rising along fault planes or associated fissures. Rogers and Reeves both postulated, because of similarity in chemical composition, that the 

water from the warm springs is of the same meteoric 

origin as that from the cold springs.

More recently Kulander and Dean (1972) studied 

gravity and structure across Browns Mountain, Wills 

Mountain, and Warm Springs anticlines. They reported that the rocks beneath Browns Mountain "anticlinorium" are cut by at least seven faults, most of 

which are interpreted as thrust faults dipping to the 

east. Displacement along a thrust fault cropping out 

just west of Minnehaha Springs is about 450 m. A similar structure is proposed for the rocks beneath Warm 

Springs anticline.

Heltz and Sinex (1974) studied the chemical equilibria in the thermal spring waters of Virginia and reported that; (1) neither concentration of dissolved constituents nor water temperature has changed in the 

past 140 years; (2) the waters with temperatures below 

25°C are undersaturated with respect to calcite and 

dolomite, whereas those above 25°C are in equilibrium 

with calcite and dolomite; and (3) the waters with temperatures less than 25°C are probably mixtures of 

warm and cool ground water.

Costain (1976) recently evaluated the geothermalresource potential in the vicinity of the Warm Springs 

anticline in Bath and Allegheny Counties, Va. He 

(1976, p. 40) concluded that meteoric water percolates 

downward through the ridges, southeast of the anticline, to a depth of about 3 km; and that it moves from 

there along thrust faults that dip gently southeastward. The water is assumed to rise relatively rapidly 

along high-angle cross-faults which intersect one or 

more of the thrust faults at depth.

Hobba, Friel, and Chisholm (1972), in a description 

of the water resources of the Potomac River basin in 

West Virginia, showed with a triangular diagram how 

the chemical composition of the water discharging 

from Berkeley Springs varies with time; the spring 

water probably is a mixture of (1) cold water from 

near-surface sandstone and carbonate rocks and (2) 

warm sodium chloride water from deeper aquifers. Assuming steady-state conditions, a geothermal gradient 

of 21.6°C/km, and sufficient length and area of flow 

path to intercept all the geothermal heat flow, the 

water is moving upward from a depth of about 530 m.

Lowell (1975, p. 355, 356), using a mathematical 

model, computed depths of circulation of 1,100 m for 

Warm Springs, Ga., 1,300 m for Warm Sulphur 

Springs, Va., and 1,700 m for Hot Springs, N. C. Lowell's model consists of two vertical fractures and an 

intervening horizontal conduit. In the model, he considers flow rates and minimum contact areas required 

to transfer geothermal heat flow (in a region of so-
HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS OF THE APPALACHIANS E3

called "normal" geothermal gradient) to the water and 

also conduction losses as the water ascends to land surface.

The spring at Lebanon Warm Spring, N. Y., has not 

been investigated in detail. However, as early as 1843, 

W. W. Mather observed that the spring was at or near 

the contact of limestone and slate beds. Mather (1843, 

p. 106) also observed " * * * traces of a fault, and of 

much derangement of the strata in the immediate 

neighborhood."

Similarly, little has been done on the Perry County 

Warm Springs in Pennsylvania. Early mention of the 

springs was made by Claypole, in 1855 (p. 348), who 

suggested that they may derive their water from surface streams. Little scientific study was done until the 

recent study by Weinman (1976), who reports that the 

springs are at the intersection of two lineaments he 

observed on ERTS (Earth Resources Technology Satellite) color composite imagery.

ACKNOWLEDGMENTS

Many people provided valuable assistance during 

various phases of the investigation. J. B. Foster and R. 

L. Wait made geophysical logs of wells. T. K. Ellis and 

G. Goold provided records of wells in Virginia and New 

York, respectively. R. Rogers, P. Gentry, R. Taylor, and 

J. D. Henry monitored spring temperatures in 

Virginia, North Carolina, and West Virginia, respectively. S. M. Pickering, Jr., R. P. Lowell, J. K. Costain, and B. Weinman provided helpful discussions and 

information regarding thermal springs in Georgia, 

New York, Virginia, and Pennsylvania, respectively. 

Personnel of the Pennsylvania Geological Survey 

supplied well information from their files. A. H. Truesdell helped greatly with advice on the interpretation of 

the geochemistry of hot-water systems and of various 

geothermometers. L. N. Plummer provided valuable 

assistance in devising computer programming techniques for the calculated carbonate geochemistry data 

in this report. F. W. Trainer and F. H. Olmsted furnished guidance and support throughout the study; 

Olmsted assisted in combining two separate manuscripts on hydrology and geochemistry into the present 

paper. Finally, the cooperation of all the individuals 

who permitted access to their property, wells, and 

springs is particularly appreciated.

HISTORICAL BACKGROUND

Many of the warm springs included in our investigation have long and interesting histories. While George 

Washington was surveying for Lord Fairfax in 1748 he

bathed in "ye fam'd warm springs" at the present site 

of Berkeley Springs, W. Va. Later Lord Fairfax 

granted the springs and 50 acres (20 ha) to what was 

then the colony of Virginia, with the stipulation that 

the waters were "to be forever free to the public for the 

welfare of suffering humanity." Even before Washington visited these springs the Indians used the springs 

for bathing, as they did those at Warm Springs, Ga. 

Later the Georgia springs were to be used by President 

Franklin D. Roosevelt in treatment of his polio.

After discovery of the warm springs by white men, 

development soon took place. At some springs elaborate pools and hotels were built to accommodate 

.visitors seeking the healing and medicinal powers 

claimed for the warm waters. Many of these spas 

flourished in the late 1700's and early 1800's when 

there was no treatment or cure available for ailments 

such as rheumatism, arthritis, polio, diabetes, gout, 

stomach disorders, or kidney stones. The hotels and 

spas built at many of the springs continue to flourish as 

resorts today, as does the Homestead Hotel at Hot 

Springs, Va. (fig. 1). Some like the Warm Springs Hospital at Warm Springs, Ga., continue to flourish as 

hospitals. Many of the buildings at some spas fell into 

disrepair, were destroyed, or were converted to other 

uses, but some warm springs, not included in this 

study, still sustain prosperous resorts such as those at 

White Sulphur Springs and Capon Springs, W. Va. u

FIGURE 1.—Homestead Hotel and Hot Sulphur Spring (foreground) 

at Hot Springs, Va.
E4 GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

GEOLOGIC AND HYDROLOGIC SETTING

By W. A. Hobba, Jr. 

GENERAL DESCRIPTION OF THE SPRINGS

The part of the eastern United States covered by this 

report extends from western Georgia to eastern New 

York. This area lies primarily in the Valley and Ridge 

province, which is characterized by a series of 

northeast-trending valleys and ridges. The maximum 

altitudes in the central part of this area (West 

Virginia, Virginia, and North Carolina) range from 

1,200 to 1,400 m above sea level. The valley floors, 

where the warm springs commonly emerge, are at elevations of 670-700 m in the Virginias and at about 400 

m in North Carolina. Toward the north and south the 

maximum altitudes decrease; in New York the 

maximum altitude is about 700 m and the valleys are 

at about 210 m; in Georgia the maximum altitude is 

about 430 m and the valleys are at about 240 m.

The thermal springs are generally at the valley 

margins or a short distance up the valley walls. One 

spring in North Carolina discharges upward through 

about 4.5 m of alluvium on the valley flat. Most of the 

others discharge directly from bedrock or through weathered residuum. Most of the thermal springs are in or 

near small rural communities surrounded by forested 

hills and farmed valleys. Growths of green algae and 

water cress are common in the warm springs and their 

runoff channels.

Temperatures of the spring waters range from 17.7° 

to 41.7°C. On cool days vapor clouds commonly form 

over the warmer springs and their runoff channels. Recent discharges measured at the springs range from 4 

to 293 L/s; these are probably minimum discharges because at some sites leakage occurs around the measuring site through the alluvium or other unconsolidated 

material.

PHYSICAL SETTING

GEOLOGY AND TOPOGRAPHY

Most of the warm springs in the eastern United 

States are in the folded and faulted Valley and Ridge 

province of the Appalachian Mountains. Ten of these 

spring groups having relatively large yields and temperatures above 15°C were selected for this study. 

These spring groups comprise: Lebanon Springs, N.Y.; 

Perry County Warm Springs, Pa.; Berkeley Springs 

and Minnehaha Springs, W. Va.; Bolar Spring, Hot 

Springs, Warm Springs, and Falling Spring, Va.; Hot 

Springs, N.C.; and Warm Springs, Ga. Although many 

of the springs are widely separated geographically, 

similar geology and topography characterize each of 

the warm-spring areas. All the springs discharge from

sandstone or limestone; and, with two exceptions, 

every spring is in a valley on the crest or flank of an 

anticline. (The exceptions are Lebanon Spring in New 

York and Hot Springs in North Carolina. They discharge from faulted limestones of complex structure, 

Lebanon Spring on a hillside and Hot Springs in a 

synclinal valley that has been described by some writers as a graben and by others as a horst.) As early as 

1840, W. B. Rogers observed that 46 of 56 thermal 

springs he studied in the Appalachians of Virginia and 

West Virignia are on or near anticlinal axes. Regarding the thermal springs of Virginia, he (Rogers, 1843, 

p. 331) concluded that, in general, the springs "issue 

from the lines of anticlinal axes, or from points very 

near such lines."

Topography, lithology, and structure are important 

factors in the occurrence and character of thermal 

springs in the Appalachian Mountains. Topography 

controls potentiometric head relations that make possible deep circulation of the water and its return to the 

surface. The large yields of some of the thermal 

springs, and the fact that the water retains appreciable 

heat until it reaches the surface, indicate the presence 

of effective flow conduits. These conduits probably are 

related to both lithology and structure. Both the 

sandstone and the limestone of this region can provide 

the permeability required to permit the rapid upward 

migration of warm water—sandstone through fractures and intergranular openings, limestone through 

solution-widened fractures. In the more "plastic" or 

less competent shale, the weight of the overlying rock 

may lead to decrease in permeability with depth; under 

this condition the upward movement of warm water is 

slow, and the water has an opportunity to cool as it 

moves upward just as it warmed when it moved 

downward. The depth of circulation in shale is not 

known; the scarcity of thermal springs in shale may 

reflect either lack of deep conduits in this type of rock 

or a tendency for the water to cool after rising from 

depth, as already suggested. The effect of structure on 

occurrence of the thermal springs is considered in the 

next section of this report.

ROCK STRUCTURE

The occurrence of warm springs in the eastern 

United States suggests a relationship between the 

structure of the rocks and the deep flow systems that 

develop and produce the warm springs. As noted earlier, the warm springs are typically on the crests or 

limbs of anticlines. Several structural features in these 

settings favor vertical permeability: openings along 

bedding planes, open tension fractures, and open faults 

and other fractures common to folded rocks. Interest-
HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS OF THE APPALACHIANS E5

ingly, thermal springs are unknown along faults in the 

Appalachian Plateaus province where the rocks are 

relatively flat lying and faulted and the relief is high. 

Similarly, at many localities in the Valley and Ridge 

province the rocks are folded, faulted, and fractured, 

topographic relief is great, and the same geologic formations from which warm springs issue at other localities are present, but all the springs are cold. Apparently deep circulation paths have not been developed at 

these sites. The warm springs are found where deeper 

circulation paths have been developed preferentially to 

the shallower circulation paths.

It is relatively easy to identify the probable dip of 

bedding planes and the general orientation of tension 

fractures, but it is more difficult to identify faults. 

However, satellite imagery is helpful in identifying 

linear features, some of which represent faults or fracture zones. Some faults, especially those that cross the 

strike of the bedding of the rocks, stand out well on 

ERTS imagery. Figure 2 shows a NW-SE trending 

fault passing through Webster Springs, Huntersville, 

and Minnehaha Springs, W. Va., and through Warm 

Springs, Va. This fault gives rise to a warm spring at 

Huntersville and perhaps to two other warm springs 

just northwest of Lexington, Va. In fact, a warm spring 

occurs each place the fault cuts the sandstonelimestone aquifer formed by the Oriskany Sandstone 

and Helderberg Group. The apparent northwest end of

the fault lies near Webster Springs, where it may yield 

the salty water observed in springs and wells. (One 

salty cold spring discharging in the bed of the Elk 

River was reported by Price and others (1936, p. 101) to 

contain about 4,000 mg/L chloride. During this investigation an attempt was made to sample the same 

spring but it could not be found. However, a well in 

Webster Springs, sampled during the present study,

FIGURE 2.—Satellite images of thermal spring areas in A, Virginia 

and West Virginia, B, Georgia, and C, New York. (Lettered arrows 

indicate lineaments.)
E6 GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

contained 3,240 mg/L chloride.) What appears to be 

this same fault is also apparent on the SLAR (sidelooking airborne radar) image (fig. 3). Although the 

SLAR image is much sharper and at a larger scale than 

the ERTS image, faults are more readily apparent on 

the ERTS image. However, the offset of ridges by faults 

is more apparent on the SLAR image.

In the Minnehaha Springs area, hydrologic data can

be used to aid in the interpretation of fault data obtained from the imagery. Figure 4 shows the geology of 

the area around Minnehaha Springs. Well 404, 14 m 

deep, penetrated only sandstone. The water had a 

specific conductance of 140 /^mhos/cm and a pH of 5.8; 

both these parameters are typical of water from the 

sandstone. According to the geologic map, well 405 

should be in limestone, but the well is 26 m deep and

FIGURE 3.—Side-looking radar images showing prominent lineaments and thermal springs and wells in A, Virginia and West

Virginia, and B, Georgia. (Lettered arrows indicate lineaments.)
HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS OF THE APPALACHIANS E7

yields water with a specific conductance of 80 /umhos/ 

cm and a pH of 4.9; the water quality indicates that 

this well also taps the sandstone formation. If both 

wells tap sandstone, the northwest-trending fault 

passes south of well 405 and the northern block is 

downdropped relative to the southern block, thus displacing the outcrop of the sandstone on the northern 

block westward.

Figure 5 represents an interpretation by Kulander 

and Dean (1972) of the geology in a vertical section 

parallel to this fault and about one mile north of it (see 

fig. 6). Their section is based on seismic, well, gravity, 

and magnetic data. The warm springs at Minnehaha 

Springs and at Huntersville both occur at points in the 

outcrop of the Oriskany Sandstone where the fault 

parallel to the line of the section intersects northerly 

trending faults. The northerly trending fault at Huntersville was mapped by Kulander and Dean (1972). 

Another fault passing directly through Minnehaha 

Springs is visible on the ERTS image and was mapped 

by the authors (fig. 2). It trends about N. 3° W. and 

extends about 74 km. A second fault just east of Minnehaha Springs strikes about N. 3° E. It appears to be 

about 105 km long, reaching 68 km south of Natural 

Well, Va. The latter fault was found in the field at 

Ruckman Spring (a large cold spring, 408 fig. 4) prior 

to observing it on the ERTS imagery. These observations of the occurrence of springs along faults or lineaments are in agreement with those of Weinman (1976), 

who pointed out that the warm spring in Perry County, 

Pa., lies at the intersection of two lineaments observed 

on color-composite ERTS imagery. He also found that 

the warm water at the spring is detectable on lowaltitude thermal imagery. Werner and Medville (1975) 

pointed out that in West Virginia both warm and cold 

springs occur on lineaments or at the intersections of 

lineaments as observed on ERTS imagery.

The lithology and structure at Minnehaha Springs, 

Berkeley Springs, and Perry County Warm Springs are 

very similar (fig. 7): each group of springs is in a valley 

and discharges from the Oriskany Sandstone on the 

east flank of an anticline. Weinman (1976) mapped two 

lineaments in the vicinity of Perry County Warm 

Springs which intersect at the springs just as similar 

lineaments do at Minnehaha Springs. At Berkeley 

Springs no intersecting pair of lineaments was observed on the satellite imagery; however, a single subtle but extensive lineament was observed that extends 

from Berkeley Springs northwestward across Pennsylvania along much of the course of the Allegheny River 

to the southeastern shore of Lake Erie. The southeastern extent of this lineament is not well defined, but it 

may follow part of the lower course of the Potomac 

River.

The lithology and structure at Warm Springs, Ga., 

are somewhat similar to that at warm springs in the 

Oriskany Sandstone in West Virginia and Pennsylvania. The warm springs in Georgia, in a valley on the 

northwest flank of an anticline formed by the Hollis 

Quartzite, discharge at the contact of the quartzite 

(which forms Pine Mountain) and the overlying Manchester Schist. Both schist and quartzite are cut by 

faults and lineaments (see arrows in fig. 2). One lineament passes through the large cold spring southeast of 

Warm Springs; another one passes northeast through a 

73-m well which produces warm water. What may be a

FIGURE 3.—Continued.
E8 GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

y&S.?£}0/ '•^••/mWW rJf&'--iW$f{ ys-jtm^jg/^iA^T--- "M:*y~M<£i»&

y3v(^:fi-^/m/ <S' &£^//mw^.

-m/-/i«fc»SKm^

^'Si \ / y-.'*''\',- >/x / s* v^s>v%/ •/ ,*/ ,

'-^ .*>' * ssjs >r , >,A"

EXPLANATION

— 38°15'

Base from U.S. Geological Survey 

Warm Springs 1:62,500,194B

Helderberg Group

I i i i i i i i i i i

1 MILE

1 KILOMETER

CONTOUR INTERVAL 40 FEET

FIGURE 4.—Geology, spring, and well locations in Minnehaha Springs area, West Virginia.

third lineament strikes eastward and passes close to 

the warm spring.

Many lineaments have been mapped in the warm 

springs area of Virginia by the authors. Most of the 

known warm springs in this area are on lineaments, 

observed on the satellite imagery (fig. 6). Dennison and 

Johnson (1971, p. 506) suggested that the thermal 

springs here may be related to Eocene volcanism or 

that they represent a still younger deep source of heat 

(magmatic). Regardless of whether intrusion has occurred, the combination of deep fracturing, the local 

geology, high relief and structure favors deep circulation of ground water here.

The lineaments mapped near Lebanon Springs, N. Y., 

stand out well on imagery with snowcover on the 

ground (fig. 2C). One prominent northwest-trending 

lineament passes through the spring, as does a lobate

northeast-trending feature (not indicated) that may be 

structural.

The satellite imagery offers good evidence that the 

thermal springs occur in valleys where lineaments or 

faults cut steeply dipping beds of sandstone or limestone. Although many thermal springs occur on the 

crest or flanks of anticlines, the anticlinal structure 

itself is not the only factor contributing to their occurrence. The tension fractures that parallel the anticlinal 

axis and the dipping rocks are conducive to deep circulation parallel to the structure. However, vertical

FIGURE 5.—Cross section parallel to fault passing near Webster 

Springs and Minnehaha Springs, W.Va. and Warm Springs, Va. 

(Modified after Kulander and Dean, 1972. Approximate line of 

cross section is along lineament trending southeasterly from Huntersville to Warm Springs in fig. 6.)
HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS OF THE APPALACHIANS E9

1000

24 32 

KILOMETERS

EXPLANATION

30,000

z o -<

w

Q

Chemung 

Formation

[ Db | 

Brallier Shale

I Dh I

Harrell Shaft

Marcellus 

Shale

Oriskany 

Sandstone

Helderberg 

Group

Tonoloway 

Formation

Wills Creek 

Formation

| Sm I 

McKenzie 

Formation

Clinton 

Formation

Tuscarora 

Sandstone

Z

u 

o 1

Q

| Qj |

Juniata

Formation

ly'.Om.";

Martinsburg 

Shale

Trenton Group 

(Limestone)

| Obi~]

Black River 

Group

Rocks of 

Chazy age

I Ob i

Beekmantown 

Group
E10 GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

38°15'

38° 00

80°00'

FIGURE 6.—Map showing lineaments and their relationship to thermal springs in Virginia and West Virginia.
HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS OF THE APPALACHIANS Ell

Marcel lus Shale Oriskany 

• Sandstone

FIGURE 7.—Sandstone quarry looking southwest toward Berkeley 

Springs, W.Va. The spring is about 5 km from this site and discharges from the sandstone near the shale contact.

faults or lineaments, rock type, dip of the rock, and 

subsurface thrust faults may be the most important 

factors contributing to the occurrence of the warm 

springs. For example, a warm spring may be just as 

likely to develop on a monoclinal fold as it would on an 

anticlinal fold having similar lithology faults, topography, and hydraulic head. Table 3, in a later section, 

summarizes the structural features at each of the 

thermal springs.

HYDROLOGY

PROBABLE FLOW SYSTEMS

All the springs investigated discharge in valleys 

whose adjacent ridges are underlain by fairly steeply 

dipping rocks. The warm springs in Georgia, West 

Virginia, and Pennsylvania are on the flanks of anticlines; in Virginia they are commonly near the crests of 

anticlines. It is of interest to consider likely hydrologic 

conditions where a warm spring is (1) in a valley near 

the crest of an anticline, (2) on the flank of a ridge 

formed by an anticline, or (3) on the flank of a ridge 

formed by a limb of an anticline.

The Virginia springs are in valleys at or near the 

crests of anticlines, where any bedding planes open to 

depth would be expected to permit flowlines of upward 

moving water to converge at the crest of the anticline. 

Figure 8 shows the theoretical flow lines of ground 

water in a homogenous, isotropic sand aquifer. Figure 

9 shows inferred paths of ground-water movement

2000

FIGURE 8.—Theoretical movement of ground water through a 

homogeneous isotropic aquifer with an impermeable layer at 1,500 

m. (After Toth, 1963.)

FIGURE 9.—Possible movement of ground water through a multilayered folded, faulted, and fractured aquifer such as those in 

Warm Springs Valley, Va. Much of the flow (not indicated by arrows) is into the figure and upward beneath the hot spring.

through the consolidated folded rocks in Virginia, in a 

plane normal to the strike of the rocks and at a place
E12 GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

where the anticline crest forms a topographic low. To 

complete the picture, add low angle thrust faults at 

depth, a vertical fault or fracture at the spring, and 

possibly a fault or fracture in the plane of the cross 

section. Any such factors would contribute to the development of a system of conduits that permits 

ground-water circulation to great depths where head is 

sufficient to drive the system. The overall flow system 

may look something like that in figure 10, but with 

much of the flow being into the plane of the figure 

along open tension fractures, bedding, or fault planes 

and upward beneath the spring.

The occurrence of springs along the crests of anticlines in valleys or in water gaps in the Valley and 

Ridge Province of Virginia and West Virginia has been 

noted by Rogers (1884), Price, Haskins, and McCue 

(1936), Hobba, Chemerys, Fisher, and Pearson (1977), 

and Clark, Chisholm, and Frye (1976). Clark, 

Chisholm, and Frye (1976) made statistical analyses of 

the yields of wells in various geologic and topographic 

situations. Working mainly with wells in limestone, 

they determined that yields along the axes of anticlines are significantly better than those along axes of 

synclines. Statistical tests by Siddiqui and Parizek 

(1972) showed that variation in structural control is 

not a significant factor in the productivity of wells not 

on fracture traces. The relatively high well yields reported by Clark, Chisholm, and Frye (1976) may reflect 

the presence of fractures that parallel the anticlinal 

axes. These fractures, along with deep vertical faults 

and perhaps deep horizontal thrust faults, could permit 

a deep flow system such as that in figure 9 to develop.

Fault 

trace;

FIGURE 10.—Possible circulation of water to a warm spring located 

near the crest of an anticline at the intersection of two faults. 

Much of the flow is into the plane of the figure along open tension 

fractures, bedding, or fault planes and upward beneath the spring.

Predominantly deep circulation favors warm springs; 

predominantly shallow circulation favors cold springs.

The second type of flow system, which may develop 

on an anticline forming a topographic high, is represented by figure 11, which shows the flow pattern in a 

vertical plane normal to the strike of the rocks. Water 

moves down through the open fractures (joints and 

bedding-plane openings) to as great a depth as permeability and head allow, and then it returns to the 

land surface quickly along a fault or through other 

fractures. It seems unlikely, because of the generally 

short distances between the crest of the ridge and the 

warm spring, that a system like this would be capable 

of developing deep enough circulation to produce a 

warm spring (fig. 11). However, theoretical computations by Toth (1963) for a homogenous isotropic aquifer 

with an impermeable boundary at 1,500 m depth suggest that circulation to a depth of about 700 m is possible (fig. 8). His model roughly approximates the situation at some of the warm springs. Although circulation 

to this depth is sufficient to produce some of the 

moderately warm springs, it is insufficient to produce 

the warmest springs.

Figure 12 depicts the third type of geologic setting 

common at some of the warm springs. Here a ridge is 

formed by the flank of an anticline. The warm springs 

discharge from a sandstone unit underlain by carbonFIGURE 11.—Possible movement of ground water through a faulted 

and fractured anticlinal ridge bordered by a synclinal ridge. 

Aquifer has impermeable boundary at 1,500 m depth. Flow pattern 

is in a vertical plane normal to the strike of the rocks.
HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS OF THE APPALACHIANS E13

FIGURE 12.—Possible movement of ground water through a faulted 

and fractured anticlinal ridge bordered by a synclinal ridge. A 

ridge is formed by a flank of the anticline so that the water may 

flow by way of two routes—along a vertical fracture normal to 

A-A' and along the aquifer shown by the stippled pattern, in the 

plane of the cross section.

ate rocks and overlain by shale and siltstone. Water 

enters the ground and may flow to the spring by way of 

two routes or flow systems: The one indicated by the 

arrows in figures 12 and 13 shows the flow to be almost 

entirely within the plane of the sandstone formation. 

Figure 14 shows recharge and movement to the spring 

along a vertical fracture approximately normal to the 

line of A-A' (The lines of the sections in figures 12 and 

14 are shown in figure 13.) This type of setting exists 

at the Warm Spring in Georgia and in West Virginia 

and Pennsylvania. A significant amount of hydrologic 

data is available in the vicinity of the warm spring at 

Berkeley Springs, W.Va. The situation is depicted in 

figure 15, and the data suggest that the flow system 

shown in figures 12 and 13 could be operating there 

with a small amount of water being added from the 

formations above or below the sandstone. Similar flow 

systems may be operating at the Perry County Warm 

Springs and at Minnehaha Springs.

SELECTED FLOW SYSTEMS

Figures 15 and 16 show the warm spring at Berkeley 

Springs and the wells along the contact of the shale 

and sandstone along the ridge that trends southwest 

from the spring. The water table in figure 15 is based 

on reported and observed water-level data from the

FIGURE 13.—Plan view showing section lines for figures 12 and 14 and 

possible flow lines in the plane of the sandstone formation.

FIGURE 14.—Diagram showing possible movement of water along 

a vertical plane parallel to the strike of an anticlinal ridge.

wells in the diagram. Wells 302, 304, and 308 are 

drilled, at least in part, into the Oriskany Sandstone. 

Wells 304 and 308 were reportedly 46 m deep and are 

either dry or produced only wet clayey sand near the
E14 GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

Fai

FIGURE 15.—Block diagram showing possible circulation of water to 

the warm springs at Berkeley Springs, W.Va. Water quality data 

suggests that most of the water discharged at the spring has circulated only through the sandstone.

bottom. Well 302 is the only one in the sandstone that 

has not been destroyed. It is 121 m deep; in the upper 

116 m it penetrates the Marcellus Shale, and in the 

bottom 4.5-6.0 m it penetrates the Oriskany 

Sandstone. During drilling some water entered the 

well at less than 18 m, but it was cased off with 19 m of 

casing. When this well was first visited on August 19, 

1975, the water level was 31m below land surface. At 

this time the well had filled with loose sand to a depth 

of less than 107 m. In order to get a water "sample from 

the Oriskany Sandstone, the well was cleaned out on 

October 16, 1975. At this time water was observed 

squirting into the well above 49 m, and the water level 

recovered to about 49 m. On May 19, when the well was 

sampled again, the water level was 50.1 m, and some of 

the water, still squirting into the well, was entering 

between 29 and 32 m. The higher head in the shale 

favors movement of water from the shale into the 

sandstone. Prior to removal of the sand, the water level 

was higher because fine sand had partially plugged the 

well bore, impeding flow down the hole and into the 

Oriskany Sandstone. The water on January 14, 1976, 

was at about 189 m altitude—the same as the warm 

spring at Berkeley Springs. This level was about 37 m 

lower than the altitude of the main stream channels 

draining to the northeast on either side of the Warm 

Springs Ridge. Wells 304 and 308 reportedly were 

drilled to depths of about 46 m and were either dry or 

had wet clayey sand near the bottom. If these reports 

are accurate, then the water table at these wells may

be lower than the altitude of the warm spring. Despite 

a difference in water-level, this area could still be the 

recharge area for the warm spring, because the column 

of warm water rising at the spring could be driven by a 

column of lesser head of cooler, more dense water in the 

recharge area. For example, a 600.0-m column of water 

at 22° C at the springs could be balanced by a 599.0-m 

column of water at 13° C in the recharge area.

If ground water from the area of wells 302, 304, and 

308 does not discharge to the warm spring, it must 

discharge to Sir Johns Run, or to the Potomac River, 

north or west of the warm spring. Subsurface flow toward the north or west is possible, but permeability 

across the strike of the bedding probably is not great 

enough to permit a high flow rate. Also, if such flow 

were happening, the water that now discharges at the 

spring probably would be diverted underground to Sir 

Johns Run or the Potomac River.

Assuming recharge of about 880,000 L/day per km2 

into the sandstone-limestone aquifer, the recharge 

area necessary to maintain the flow of the warm spring 

is about 10.6 km2 . If the actual area of recharge is 

assumed to be all the ridge to a distance of 6.1 km 

south of the spring and parts of the valleys on either 

side of the ridge (which is 1.3 km wide), the recharge 

area is about 9.8 km2 , or approximately the inferred 

area of recharge. However, if a large amount of recharge were derived from the streams on either side of 

Warm Spring Ridge, the actual recharge area would be 

indeterminate.

The chemical content of the water discharging from 

Berkeley Springs suggests that nearly all the water is 

derived from the Oriskany Sandstone. However, the 

water shifts from a calcium magnesium bicarbonate 

type at high flow toward a sodium chloride type at low 

flow. This shift indicates that a part of the water is 

derived from either a deeper formation or the overlying 

shale.

Rough computations can be made to estimate the 

likelihood that all the spring water is derived from 

recharge entering the outcrop of the Oriskany 

Sandstone. First, recharge is assumed to be 880,000 

L/day/km2 or 320 mm/yr in the 9.8-km2 outcrop area. 

This high rate is believed reasonable because of the 

permeable nature of the Oriskany Sandstone in this 

area. Grimsley (1916 p. 224) described an outcrop at 

Rock Gap Run eight miles south of Berkeley Springs as 

forming "an apparently solid wall sloping at a high 

angle to the east valley, but when struck with a hammer it is found to be soft and crumbly, it being almost 

impossible to secure a solid hand specimen of the rock." 

Grimsley (1916, p. 224) described another outcrop, 

about four miles north of Berkeley Springs, as a conglomerate "composed of small or large white pebbles 

imbedded in a loose white sand matrix which crumbles
HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS OF THE APPALACHIANS E15

EXPLANATION

Helderberg Group

(Geology after Grimsley, 1916)

1/2 1 MILE 

J

1/2 1 KILOMETER

CONTOUR INTERVAL 20 FEET 

DATUM IS MEAN SEA LEVEL

Base from U.S. Geological Survey 

Hancock 1:24,000, 1951 (Pr 1971), 

Great Cacapon 1:24,000, 1958 (Pr 1974) 

and Stotlers Cross roads 1:24,000, 1958

FIGURE 16.—Map showing geology and spring and well locations near Berkeley Springs, W. Va.

to a gravel." Thus, it is reasonable to infer a large 

amount of recharge from precipitation on poorly 

cemented rocks such as these. If all the recharge for the

spring enters a 300-m wide outcrop of the Oriskany 

Sandstone from 18 km south of the spring and 4.8 km 

north of the spring, only two-thirds of the measured
E16 GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

flow of the spring can be accounted for. However, additional recharge is known to come from the overlying 

Marcellus Shale, and perhaps some also comes from 

the underlying limestones. If the spring derives water 

from as far as 18 km to the south, additional recharge 

could also be derived from Rock Gap Run, which flows 

across the sandstone outcrop (where Grimsley described the rocks). Thus, at least two-thirds of the discharge of the spring is probably derived from direct 

recharge on the Oriskany.

R. T. Hurr (Lohman, 1972, p. 53) reports an average 

hydraulic conductivity value of 9 m/day for medium to 

fine alluvial sand. This value converts to a 

transmissivity value (T) of about 900 m2/day for the 

Oriskany Sandstone, which ranges in thickness from 

90 to 120 m and averages about 100 m. Using a slope (I) 

of 2.33 x 10~3 for the water table from the spring to 

Rock Gap Run, the length of aquifer (L) required to 

supply the actual discharge of the spring (Q), 110 liters 

per second, was computed by the relationship Q = TIL 

and found to be 4,500 m. Thus, the water discharging 

at Berkeley Springs could be derived from recharge 

entering 4,500 m of the Oriskany Sandstone outcrop if 

sufficient recharge were available. However, all or 

most of the recharge comes from precipitation, and if 

our earlier assumption of 880,000 L/day/km2 for recharge is reasonable, then recharge must be entering 

the rocks over as much as 18-23 km of the outcrop.

It should be mentioned here that other nearby areas 

apparently duplicate the geology at Berkeley Springs 

(Hobba and others, 1972). The hydrology is similar too 

in that the water levels in the sandstone hills (formed 

by the Oriskany Sandstone), are lower than water 

levels in the adjacent shale valleys, much the same as 

the area south of Berkeley Springs. However, in these 

other areas, large cold springs are usually located some 

distance downgradient and along the strike of the 

rocks. At Berkeley Springs the only known large 

spring downgradient and along strike is the Warm 

Spring. Apparently the only difference between this 

warm spring and the cold ones is the depth of water 

circulation. The hydrologic situation near Baker, W. 

Va. (96 km southwest of Berkeley Springs) suggests 

that deep circulation may have occurred there in the 

past. The geologic conditions at this site are essentially 

the same as those at Berkeley Springs, except that the 

rocks dip to the west and a large cold spring discharges 

in the valley from a cavern in the Tonoloway Formation, which lies stratigraphically about 107 m below 

the Oriskany Sandstone. A dye-tracer study showed 

that the spring is hydraulically connected to a large 

sinkhole in the same limestone in an upland valley 

about 3 km northeast of the spring. The large cavity at 

the spring is below the intersection of a thrust fault

and a vertical fracture. At some time in the geologic 

past, when this large conduit was smaller and buried 

beneath hundreds of meters of rock, it is conceivable 

that it supplied water to a thermal spring. The age of 

the cavern is unknown, but it may be very old, since it 

is supplied water from such a large sinkhole which 

drains an area of about 2.5 km2 . Regardless of whether 

this system ever supplied water to a warm spring or 

not, cavernous circulation systems do develop parallel 

to the anticlinal ridges along certain formations and 

perhaps along faults. If developed to a great enough 

depth, such a system may give rise to a warm spring.

The hydraulic system at Hot Springs, Va. may be 

similar to the one near Baker, W. Va. One well (607) at 

Hot Springs obtains warm water from a 0.6-m-high 

cavern at a depth of 230 m. No significant amount of 

water was encountered in the well above that depth, 

and well yield exceeded 30 L/s for a 53 hour pumping 

period. This high sustained yield suggests a waterfilled cavern system at depth. Thermal springs having 

various temperatures and yields are found at both Hot 

Springs and Healing Springs, Va. which are 5.6 km 

apart along the crest of the anticline in Warm Springs 

Valley (fig. 10). A topographic high separates the two 

thermal-spring areas in the Valley. A well drilled in 

this topographic high (near well 611) was 91 m deep; 

this well penetrated caverns and was dry. If the water 

table is at or below 91m, its altitude is near that of the 

thermal springs at Hot Springs but well above that of 

the springs at Healing Springs. On the basis of the 

yield of Healing Springs and Hot Springs and an assumed rate of recharge of 730,000 L/day/km2 (270 

mm/yr) in the area between the springs, this recharge 

would supply half the total water discharged at the two 

spring areas. The other half of the water required is 

available in the part of the valley north of Hot Springs 

and that part south of Healing Springs plus any contribution from the ridges on either side of the valley.

Warm Springs in Georgia discharges from a 

sandstone formation which dips 25°-30° beneath the 

Manchester Schist. Hewett and Crickmay (1937), inferred that the recharge for Warm Springs is derived 

from precipitation on Pine Mountain (fig. 3B). They 

concluded (1937, p. 33) that there are lower and upper 

permeable zones in the Hollis Quartzite, separated by a 

nearly impermeable middle unit. Water enters the 

lower permeable bed on Pine Mountain and flows 

downdip to where the rocks are broken by Towaliga 

Fault. At this point the water moves upward into the 

upper permeable beds and then back updip to discharge at the warm spring. Not enough new data are 

available to draw definite conclusions as to the pattern 

of circulation at this warm spring. If Warm Springs 

derives its water from the area of Pine Mountain south
HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS OF THE APPALACHIANS E17

and southwest of the springs, and if the water flows to 

depth and if it is heated by a normal geothermal heat 

flow, then the springs may be part of a flow system 

similar to the type Hewett and Crickmay (1937) have 

described. However, a flow system similar to the one 

described for Berkeley Springs is also possible. That is, 

recharge entering the Hollis Quartzite on Pine 

Mountain may circulate to a depth parallel to schistosity along open fractures, joints, and faults and rise 

along a highly permeable zone to discharge at the 

springs. Figure 2 indicates that a lineament may pass 

through the warm springs. In either case, the water at 

the springs would have been derived principally from 

recharge on exposures of the Hollis Quartzite.

TEMPERATURE OF THERMAL WATERS

FLUCTUATIONS

The temperatures of the thermal springs investigated range from 17.7°C at Perry County Warm 

Springs in Pennsylvania to 41.7°C at Hot Springs, N.C. 

Continuous recording thermographs were installed at 

Warm Springs, Va., and at Minnehaha Springs, W. 

Va., in the fall of 1975. The thermographs showed temperature fluctuations of about 0.3°C at Warm Springs 

over an 8-month period and less than 1°C at Minnehaha Springs over a 3-month period.

Recording instruments were also used by Hewett 

and Crickmay (1937) at Warm Springs, Ga., and by 

Costain (1976) at Hot Springs, Va. The temperature at 

a central discharge point at Warm Springs varied from 

30.9°-31.2°C in a year and a half. However, the temperature at opposite ends of the 150-m-long discharge 

belt were 26.1°C and 28.3°C. Thus, there is a difference 

of about 5 degrees between the warmer and cooler waters discharged by the springs. Our measurements of 

the temperature of the combined flow from all the 

sources varied only 0.4°C when sampled in December 

1975 and May 1976. Costain (oral commun., 1976) 

found that the temperature of Boiler Spring at Hot 

Springs, Va., was essentially constant during a 

1-month period during which a thermograph was installed. Our measurements at the times of sampling 

showed a temperature variation of only 0.1°C at Boiler 

Spring. However, the temperatures of other nearby 

warm springs fluctuated as much as 12.2°C during the 

same period.

Weekly measurements were made by observers at 

Berkeley Springs, W. Va., and at Hot Springs, N. C., 

using maximum-temperature thermometers. The temperature at the discharge point observed at Berkeley 

Springs varied from 22.0°C to 22.2°C; that at Hot 

Springs varied from 36.8° to 41.7°C (fig. 17). (To date, 

41.7°C is the highest known temperature measured at 

any of the thermal springs in the eastern United 

States.) The major temperature fluctuations here are 

caused by cool river water moving into the spring 

house during times of high flow in the river. Figure 17 

shows that when the river stage rises to a level higher 

than about 1.5 m, water may move into the overflow 

pipe or through the alluvium and reduce the temperature of the spring.

From 1949 to 1955 flow and temperature measurements were made at Bolar Spring by the Virginia 

District of the U. S. Geological Survey. Figure 18 

shows that there is generally an inverse relationship 

between discharge and temperature. This indicates 

that the spring water probably is a mixture of deep 

circulating warm water and shallow circulating cool 

water.

December 1975 January 1976 February March April May June July

5 10 15 20 25 30 5 10 15 20 25 30 5 10 15 20 25 5 10 15 20 25 30 5 10 15 20 25 30 5 10 15 20 25 30 5 10 15 20 25 30 5 10 15 20 25

° 39 

o. co

< LU

w uj 2

cc S

,Spring overflow, submerged 

in river

FIGURE 17.—Water temperature of hot spring and stage of the French Broad River at Hot Springs, N.C.
E18 GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

500

- ai

UJ (/)

0 (T IE ui 

< 0. 300

££

Q t

200

100

24

LJJ CO

5! UJ 20

UJa! DC

1949 1950 1951 1952 1953 1954 1955

— oo

uj a:

10 a LJJ u tr Q.

<\-

I LJJ

O LJJ

CO U.

Q

FIGURE 18.—Temperature and discharge of Bolar Spring at Bolar, Va.

Weinman (1976) measured the temperature of Perry 

County Warm Springs periodically; he found that it 

fluctuated from 17.5° to 19.5°C. He also found that even 

though this spring is not very warm, the temperature 

difference is great enough to appear on low altitude 

thermal imagery. Our measurements at the times of 

sampling varied only from 17.8°C in the fall to 17.7°C 

in the spring.

The temperatures at Lebanon Springs, N. Y., and 

Falling Spring, Va., were measured only at times of 

sampling. Lebanon Springs varied from 21.9°C in the 

fall to 18.7°C in the spring. Falling Spring varied from 

25.0°C in the fall to 20.8°C in the spring. These lower 

temperatures during the spring suggest that the discharge is a mixture of deep-circulating warm water 

and cool shallow water.

WELLS AND THERMAL GRADIENTS

During this investigation three wells were inventoried that produced warm water. Two are in Virginia 

and one is near Warm Springs, Ga. Temperature and 

other geophysical logs were obtained on one of the open 

wells (607) at Hot Springs, Va. The temperature and 

fluid conductivity logs are shown in figure 19. This hole 

is reported by the driller to be cased to 47.2 m with 

150-mm casing. The only water reported is warm water 

from a cavern between 229.8 and 230.4 m. However,

the fluid conductivity log shows two types of water in 

the well: one of lower conductivity above 109.7 m and 

one of higher conductivity below 109.7 m. Although the 

log shows that conductivity increases from 750 to about 

1,100 /amhos/cm below 122 m, much of this increase is 

probably caused by the increase of temperature with 

depth because the conductivity of point samples collected at 122 and 227 m were 700 and 750 /amhos/cm, 

respectively, when cooled to the same temperature at 

land surface. The yield of the well is substantial; (it 

was test pumped at 30 L/s for 53 hours). When sampled 

with a pump that would pump only 0.41L/S, the temperature and specific conductance of the discharge increased from 15.0°C and 380 micromhos to 20.2°C and 

670 micromhos, after 2 l/2 hours of pumping. During the 

period, the water level rose 70 mm in the well. The rise 

in temperature, conductance, and water level with 

pumping is attributed to the warm water of higher 

conductivity entering the bottom of the well and moving upward. The expansion of this warm water was 

enough that the water level in the well increased although pressure head actually decreased.

Figure 19 shows the temperature log for well 302 

near Berkeley Springs, W. Va., and that of the warm 

well at Hot Springs, Va. At well 302 the cool water is 

flowing into the well from the Marcellus Shale, down 

the well bore, and out into the Oriskany Sandstone, the
HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS OF THE APPALACHIANS E19

40 -,

60 -

80-

100 -

120 —

• 140 —

160-

180-

200-

220

230

Cold-water well.

Berkeley Springs,

West Virginia

Warm-water well;

Hot Springs,

Virginia

-1100 

"//mhos

11.7° 12-2° 15-6° 21-1° 26-7° 32-2° 

TEMPERATURE, IN DEGREES CELSIUS

FIGURE 19.—Logs showing temperature for a cold-water well and 

temperature and fluid conductivity for a warm-water well.

top of which lies at about 116 m. This well is thought to 

be located in the recharge area for the warm spring of 

Berkeley Springs. The warm water in the well at Hot 

Springs is flowing in at the bottom of the well, up the 

well bore, and out into the overlying formations. This 

well is thought to be located in the same discharge area 

as the thermal springs at Hot Springs, Va.

Increases in temperature and conductance with time 

of pumping were common to all the wells that produced 

warm water. The temperature and conductance of well 

702 near Falling Spring, Va. steadily increased to 

27.2°C and 1,050 micromhos after more than one hour's 

pumping. With continuous pumping the temperature 

of water from this well is reported by the tenant to 

reach 35°C. After one and one-half hours of pumping, 

the temperature and conductance of well 907 near 

Warm Springs, Ga., had increased to 33.8°C and 100 

micromhos. This is 2.6°C higher than the highest temperatures observed at the Warm Springs (Hewett and 

Crickmay 1937, p. 5). Possibly this well is tapping the 

conduit that supplies warm water to the spring, or 

perhaps the water from the well is mixed with a 

smaller amount of cool water than the water from the 

spring.

Thermal gradients were computed using bottom hole 

temperatures or maximum temperatures of water 

pumped from wells (table 1). It is apparent from table 1 

that some of the shallow hot wells have thermal gradients about 100 times that computed for well 302 near 

Berkeley Springs. The small gradient at well 302 can 

be explained by the downward flow of cold water in the 

well. The large gradient in the warm wells can be explained by the upward movement of warm water from 

depth. The thermal gradients for test well 703 in 

Virginia and gas well Pocahontas #21 in West 

Virginia are considered about normal for the eastern 

United States.

Figure 20 is the temperature log for gas well 

Pocahontas #21. It shows a nearly linear increase in 

temperature with depth. This well is 3,111 m deep and

TABLE 1.—Thermal gradients at wells producing warm or cold water

Wells

Virginia: 

Homestead hot well (607)

Webb hot well (702) _____ _

West Virginia: 

Wachter domestic well (302)

Gas well Pocahontas #21

Georgia: 

Phillips hot well (907) ___ __

Thermal 

gradient 1

(°C/Km)

— — 94.2

____470

— _ 6.6

_ 21.4

_ _234

Depth

(m)

230.4

50.9

121.3

3,111.4

73.1

Bottom-hole 

temperature

(°C)

32.8

35

11.9

78.2

33.8

'Thermal gradient computed from bottom hole temperature and local shallow ground-water 

temperature. For water wells, highest measured or reported water temperature was assumed to be bottom hole temperature.
GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

3000 —

3111 =

50 55 60 65 70 75 

TEMPERATURE, IN DEGREES CELSIUS

FIGURE 20.—Temperature log for air-filled "dry" gas well (Pocahontas 21) near Minnehaha Springs, W. Va.

is 3 km southwest of Minnehaha Springs, W. Va. It 

penetrates the Cambrian and Ordovician rocks that 

underlie the warm spring areas of West Virginia, 

Virginia, and Pennsylvania. The thermal gradient 

here is 21.4°C/km. On the basis of this thermal gradient, the depth of circulation at Minnehaha Springs is 

estimated to be more than 450 m. Table 2 shows 

minimum depths of circulation at these warm spring 

areas inferred on the basis of known geothermal graTABLE 2.—Estimated depths of circulation based on thermal 

gradients and conduction losses

Springs

Perry Co. Warm Springs, Pa

Berkeley Springs ^VV

Hot Springs, Va

Hot Springs, NC _ — —

Thermal c 

gradient

(°C/Km)

1 26.6

1 26.6

. _ _ ' 23.0

. _ - 21.4

_____ ' 3 20.0

. _ ' 3 20.0

.____ ' :! 20.0

._ _ ' 3 20.0

._ _ fi 24.8

._ __ 7 21.2

Estimatec 

:onductioi

loss

(%}

5

3

3

2

2

6

8

2

3

5

1 Minimum 

i depth of 

circulation

(mi

500

250

500

450

600

1,300

4 1,600

5 1,450

1,300

750

'Am. Assoc. Petroleum Geologists and U. S. Geol. Survey, 1976.

•Measured in nearby gas well Pocahontas #21.

"Lowell, 1975, p. 356.

4For the hottest spring.

5Costain (1976), p. 31) reports a the

dients, assumed temperature equilibrium at depth, 

and estimated conduction losses. Conduction losses 

were estimated from measured spring flows, temperature, and the following assumed values (Lowell, 1975, 

p. 353-356): thermal conductivity of the rock (2.5 W 

m" 1 ^-M, specific heat of water (4.2 x 10:i j kg- 1 °C- 1 ), 

thickness of the resistance (heat-transfer) layer equals 

depth of circulation, and fracture lengths equal 1 km 

for small springs and 10 km for large springs.

HEAT DISCHARGE

An impressive amount of heat is carried to the 

earth's surface and released by the thermal springs. 

Computations based on flow and temperature data 

show that the heat released at larger but cooler springs 

such as Falling Spring and Bolar Spring, Va, is much 

greater than that released at the smaller but warmer 

springs (table 3). Since the discharge at these cooler 

springs is a mixture of warm and cold water, the components of hot water must be either larger or hotter 

than that emerging at the smaller warmer springs. 

Thus, it seems that areas of the springs with the 

greatest heat output (as power equivalent) would be 

the best areas to investigate for geothermal potential.

GEOCHEMISTRY

By D. W. Fisher, F. J. Pearson, Jr., 

and J. C. Chemerys 

GENERAL PRINCIPLES

Several combinations of processes can lead to the appearance of thermal water 1 at the earth's surface. The 

chemical and isotopic measurements discussed here

'Thermal water is considered herein to be water having a temperature above the mean 

annual temperature near the earth's surface.
HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS OF THE APPALACHIANS

TABLE 3.—Flow, temperature, and power output of thermal springs

E21

Spring 

No.

101

201

301

401

501

601

602

603

604

605

606

631

632

633

701

801

901

Spring owner or name

Lebanon Springs, N.Y

Perry County Warm Springs, Pa

Minnehaha Springs, W. Va

Hot Springs, Va

Octagon Spring

Boiler Spring

Hot Sulphur Spring

Magnesia Spring

Cold Magnesia Spring

Total

Warm Springs, Va

Children's Pool

Men's Pool

Total _

Falling Spring, Va

Hot Springs, NC

Warm Springs, Ga

Date

(1975)

___ 10- 4

___ 10- 6

___ 10- 2

___ 11-18

___ 12-16

___ 11-21

___ 11-21

_ _ 11-21

___ 11-21

_ _ 11-21

___ 11-21

^__ 11-26

___ 11-25

___ 11-25

_ 11-25

___ 11-25

___ 12-15

_ _ 12-6

___ 12-10

Flow

(1/s)

6.6

8.83

103

13.88

130

3.78

.03

8.5

63

11.6

28.4

11.7

51.7

204

.5

55.8

Temperature

ro

21.9

17.8

21.8

20.4

99 9

36.5

39.9

36.5

16.2

16.3

13.4

20.8

35.4

35.0

Q A Q

34.8

24 9

39.3

30.9

Power output 

at land surface

( Watts xlO6 )

.369

.287

4.55

.580

6.35

9 7Q

5.29

12.3

OQ

3.37

Average annual 

air temperature

8.3

10.0

11.1

10.6

10.6

10.6

10.6

10.6

10.0

16.7

Structural 

features 

at spring

F

A T

A, F

A,L

A, F

A, F

A, F

A,F

F

A, F

A Crest or flank of anticline.

F Fault or lineament.

L Intersection of lineaments or faults.

were intended to identify which of several possible 

combinations of processes gave rise to the observed surface temperatures of the thermal springs and wells 

sampled.

The processes can be envisioned as follows: Ground 

water is in some way heated—for example, by deep 

circulation in a region of so-called "normal geothermal 

heat flow." The water presumably attains temperature 

equilibrium with the aquifer rock at depth. This water 

may then flow to the surface rapidly enough so that it 

neither loses appreciable heat by conduction to nearsurface cooler rocks nor mixes with cooler, shallowcirculating ground water. The surface temperature of 

such a water is that of the region in which it was originally heated. On the other hand, the heated water may 

ascend slowly, losing heat by conduction, in which case 

its surface temperature will be below its maximum 

temperature at depth, but it will show no evidence of 

having been mixed with shallow, cooler water. Finally, 

if the heated water mixes with cool, shallow water before issuing forth at the surface, its surface temperature will also be below its maximum temperature at 

depth, but it should have characteristics which show 

that it is a mixture.

The measurements made on the samples from the 

systems described here were intended to permit the 

estimation of the maximum temperatures attained by 

the heated component of each of the systems. The measurements and their use are described in detail below. 

Briefly, they include tritium measurements to estimate the amount of cool shallow water present in each 

sample, measurements of dissolved constituents which 

are useful as geothermometers or as qualitative indicators of heat-generating reactions, and consideration 

of the carbonate geochemistry of the samples as indicators of the aquifer rock type through which the water 

has passed.

GEOCHEMICAL INDICATORS

Cation and anion distributions and total mineralization are useful for delineating differences between 

waters from various sources. Values of dissolved-solids 

concentrations, a measure of total mineralization, calculated as the simple sums of analyzed cations, anions, 

and nonvolatile neutral constituents, are listed in table 

4. Variations in cation and anion distributions are discussed under the individual warm-spring areas.

DISSOLVED GASES

Dissolved gas concentrations can also be helpful in 

understanding the geochemistry of waters. High dissolved oxygen is typical of a water that has had little 

contact with organic matter or sulfide minerals during 

its subsurface travel. On the other hand, the presence 

of methane and (or) hydrogen sulfide in a water 

signifies highly reducing conditions in the aquifer 

through which the water flows. Dissolved oxygen concentrations of the sampled Appalachian warm springs 

exceed >3 mg/L except for those in the Hot SpringsWarm Springs area of Virginia, where levels range 

from 0.1 to 2 mg/L. A hydrogen sulfide concentration of 

0.75 mg/L in the Warm Springs sample (631) suggests 

that sulfide minerals may be consuming the dissolved 

oxygen of the recharge to these warm springs.

Roberts, Friedman, Donovan, and Denton (1975) 

have proposed the use of helium concentrations in soil
E22 GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

TABLE 4.—Temperatures and selected analytical data for waters in Appalachian warm-spring areas

[Locations and hydrologic data for these wells and springs are given in Hobba, Friel, and Chisholm (1977)]

Sample

No.

Owner or name

Temperature (°C) 

DisDis- solved CE 

solved silica ciu 

Date Ob- Chal- solids (Si02 ) (C 

served cedony Quartz Cation (mg/L) (mg/L) (mg

Poll- tasim Sodium sium I

a) (Na) (K) 

/L) (mg/L) (mg/L)

Tritium

lelium 

(He) 

(.ig/L) TU ±1(7

S180

°/00

New York — Lebanon Springs area

101 

101 

109 

109 

114 

116

B. Schell __ ___ ______ ___ _____ __ __

F. Amlaw _

5-76 

- __ 10 75

5-76 

______ 10-75

______ 10-75

21.9 10.7 45.6 6.6 271 12 

18.7 1.1 36.2 -.4 266 9.0 

14.0 7.8 42.7 10.0 271 11 

12.2 7.8 42.7 10.3 253 11 

10.9 -11.3 23.9 -9.5 698 6.0 

10.4 16.1 50.8 -23.3 352 14

35 

40 

31 

31 

133 

49

6.9

5.5 

7.0 

7.3 

28 

2.8

1.2 

1.0 

1.3 

1.3

.8 

.4

Tr<2 

1.0 

<2 

1.3 

<2 

<2

11.6 

19.9 

1.0 

-.4 

.3 

.6

+ 1.0

±1.3 

+ 1.4 

±.5^ 

±.9 

±1.0

-9.9 

-10.7 

-10.0 

^-9.8

-9.4

Pennsylvania — Perry County Warm Springs area

201 

201 

202 

216 

218

Warm Springs

M. Loy Spring

______ 10-75

5-76 

_ _ 10-75

10-75

17.8 1.1 36.2 -19.9 197 9.0 

17.7 -5.0 30.1 -19.8 193 7.4 

12 -21.4 13.9 -17.6 117 4.2 

12.8 -5.4 29.7 -8.0 96 7.3 

12.0 -11.8 23.4 -12.1 308 5.9

38 

40 

25 

17 

65

1.6 

1.8 

.6 

1.4 

3.1

.5 

.5 

.6 

.6 

.9

<2 

<.5 

<2 

<2 

<2

6.0 

8.3 

55.6 

15.5 

78.1

±1.0 

±.6 

±3.0 

+ 1.4

±4.1

-7.7 

-8.2

-8.8 

-7.5

West Virginia — Berkeley Springs area

301 

301 

302 

302 

303 

307

F. Wachter ___ ____________ _ ____

F. Wachter _ __ _____ _

H. Hovermale

_ — 10-75

5-76 

______ 10-75

5-76 

_ _ &-75

- _ _ 10-75

22.2 2.9 37.9 -3.8 258 9.5 

21.9 1.1 36.2 -6.1 249 9.0 

13.0 29.5 63.7 -17.6 430 20 

12.8 23.3 57.7 -17.5 418 17 

13.9 25.4 59.8 -14.4 405 18 

16 -3.4 31.7 -20.9 596 7.8

45 

47 

80 

83 

55 

106

4.1 

4.3 

6.2 

6.8 

17 

3.8

1.0 

.9 

.6 

.6

.4 

.7

<2

<.5

<.5 

<2 

<2

2.7 

2.6 

36.2 

40.2 

45.3 

71.8

±.8 

±.6 

±2.4 

±2.1 

±2.6 

±3.7

-8.7 

-8.6 

-7.9

-8.2 

-7.8

West Virginia — Minnehaha Springs area

401 

401 

402 

402 

403 

406 

407 

410

501 

501

A. T. White _ __ __ __ _

T. Rose _ _ __ __ _ _ .___

_ 11-75

6-76 

______ 11-75

6-76

U 7E

__ _ _ 11 75

_______ 6-76

6-76

20.1 -1.1 34.0 -10.1 212 8.4 

20.3 -2.2 32.9 -9.8 213 8.1 

11.6 16.1 50.8 -23.5 364 14 

13.5 13.5 48.3 -24.5 363 13 

13.3 -7.6 27.6 -12.8 160 6.8 

10.3 16.6 18.7 19.2 164 5.0

11.8 -7.6 27.6 17.5 80 6.8 

13.8 -5.9 29.3 96.9 5650 7.2

Virginia — Bolar Spring area

22.2 7.8 42.7 3.0 323 11 

20.2 -2.2 32.9 1.2 288 8.1

39 

38 

61 

64 

36 

36 

14 

129

58 

54

2.2 

2.2 

4.2 

3.9 

1.2 

.4 

2.6 

1920

1.6 

1.6

.8 

.8 

.4 

.4 

.8 

.8 

1.5 

12

2.3 

2.0

Tr<.l 

<.l 

•c.3 

<.3

<.l 

<.l 

.6

Tr<,2 

•c.3

2.5 

1.3 

41.0

62.9

87.6 

-.3

24.4 

27.2

±.9 

±.6 

±2.4

±3.6 

±4.5 

±.6

+2.0

±1.6

-8.5 

-8.6 

-8.1

-7.2 

-8.0 

-7.6

-8.0 

-8.1

Virginia — Hot Springs — Warm Springs area

602 

602 

605 

605 

607A 

607B 

607C 

607D 

619 

619 

625 

627 

631 

631 

633

Cold Magnesia Spring

Homestead Hot Well _ __ _ _

T. H. Bonds Well _ _ ___ _ __

Ingall's Airport

_ ___!! 75

6-76 

__ __ 11-75

6-76 

_ _ 11-75

2-76 

2-76 

2-76

6-76 

_ — 11-75

__ __ 11 75

11-75

6-76 

__ __ 12-75

39.9 31.4 65.5 41.4 802 21 

40.0 29.5 63.7 44.1 788 20 

15.2 -1.8 33.3 7.7 302 8.2 

14.0 -6.7 28.5 5.5 271 7.0 

19.6 18.6 53.2 17.4 653 15 

17.9 -5.9 29.3 13.2 317 7.2 

19.5 16.1 50.8 35.9 697 14 

32.2 27.5 61.8 40.7 735 19 

15.8 .8 35.9 3.0 394 8.9 

14.4 -5.9 29.3 -1.0 342 7.2 

12.0 -11.3 23.9 -10.3 349 6.0 

10.5 -14.4 20.9 -20.1 446 5.4 

35.4 31.4 65.5 25.0 607 21 

35.4 29.5 63.7 22.4 601 20 

9.1 -24.1 11.2 8.5 9 3.8

132 

126 

60 

53 

113 

53 

119 

123 

80 

73 

60 

96 

112 

119 

.5

7.0 

6.9 

1.7 

1.5 

2.9 

2.6 

4.9 

5.6 

2.1 

1.8 

1.5 

5.7 

3.7 

3.7 

.4

13 

14 

2.9 

2.5 

5.8 

3.0 

11 

13 

2.6 

2.1 

1.2 

.6 

7.4 

6.9 

.2

0.7 

.3 

<.l

1.0

Tr<.2 

•c.3

•C.I 

•C.I

.1 

<.3 

<.l

-3.1 

.7 

41.0 

43.5 

2.4

36.0 

34.2 

4.9 

64.6 

.9 

1.6 

68.3

±1.8 

±.5 

±2.4 

±2.2 

±2.3

±2.8 

±2.0 

±1.8 

±3.5 

±.9 

±.6 

+3.4

-7.4 

-8.3 

-7.9 

-9.1 

-7.9

-7.7 

-8.3 

-7.9

-7.9 

-8.5

Virginia — Falling Spring area

701 

701 

702

Falling Spring

H. D. Webb Well _ _ _

_ _ 12-75

6-76 

_ _ 11-75

25.0 25.4 59.8 38.4 848 18 

20.8 10.7 45.6 30.0 610 12 

27.2 36.8 70.7 -4.3 1260 24

158 

117 

234

3.8 

3.1 

6.6

16 

10 

2.5

<.3 

Tr<.3 

2.0

24.4 

32.6 

10.9

±2.2 

±1.9 

+ 1.0

-7.7 

-7.8 

-7.9

North Carolina — Hot Springs area

801 

801 

803 

803 

804 

805 

810

City Well _____ ____ _ ______ _ ____

P. Lovin

___ 12-75

5-76 

______ 12 75

5-76

__ __ 12-75

_ __ 5-76

39.3 47.7 80.9 37.3 719 31 

41.4 50.5 83.6 35.4 719 33 

15.2 38.5 72.3 54.0 202 25 

15.2 41.7 75.3 56.3 194 27 

12.5 -5.0 30.1 25.4 54 7.4 

13.2 7.8 42.7 5.9 155 11 

15.0 -15.4 19.8 4.8 210 5.2

135 

144 

13 

11 

4.9 

19 

23

10 

9.6 

6.6 

6.9 

.9 

.3 

1.8

10 

9.8 

4.5 

4.3 

1.5 

2.2 

1.3

.7 

.4 

<.l 

<.l 

<.l 

<.l

-.2 

-.5 

40.6 

33.3 

52.2 

6.0 

48.9

+2.0 

±.1 

±2.4 

±1.9 

+3.0

+ 1.0 

+2.4

-6.3 

-6.4 

-6.2 

-7.0 

-5.9 

-6.2 

-6.8

Georgia — Warm Springs area

901 

901 

905 

905 

907 

909

Warm Springs _

J. Phillips ________________________

______ 12-75

5-76

1 9 7^

5-76 

______ 12 75

_ _ 12 75

30.9 29.5 63.7 26.0 192 20 

31.6 25.4 59.8 26.8 181 18 

17.7 -4.6 30.5 49.4 14 7.5 

17.8 -9.8 26.3 50.8 12 6.5 

33.8 31.4 65.5 35.6 119 21 

17.0 47.7 80.9 44.0 170 31

22 

20 

.3 

.2 

12 

19

1.2 

1.2 

.9 

.8 

1.0 

10

3.8 

3.7 

.6 

.5 

4.0 

3.5

<.l

<.2 

<.l 

<.l 

<.2

.8 

.8 

34.6 

30.3 

-1.0 

15.1

±.9 

±.5 

±2.4 

±1.7 

±2.2 

±1.4

-4.5 

-5.0 

-4.4

-5.2 

-4.3 

-4.2
HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS OF THE APPALACHIANS E23

gas as a geothermal prospecting tool. They observed 

close correlation between helium levels and proximity 

to hot-water seeps at several locations in the western 

states; the effect was ascribed to escape of helium from 

the hot water into the local soil gas. Heat and helium 

are generated during the radioactive decay of uranium 

and thorium minerals. Decay of 238U to stable 206Pb and 

4He evolves ^3.5 x 10 10 calories of heat per gram of 

helium created (from data in Birch, 1954). Calculations 

based on these data show that the heat produced by the 

number of decay events needed to add 1/ugHe/L is about 

35 x 103 calories. If all this heat were absorbed and 

retained by the water, its temperature would rise 

about 35°C. Of course, in a ground-water system, much 

of this heat is dissipated by conduction in the aquifer 

material itself, so that the observed temperature increase will be smaller than the theoretical 35°C/ 

jiigmHe. However, an elevated dissolved-He concentration is still an indication of a subsurface heat source 

consisting of a concentration of radioactive minerals in 

or near the aquifer, and independent of the depth of 

circulation of the water.

Dissolved helium cannot be used as a quantitative 

geothermal indicator largely because of the nature of 

helium gas. It is the least soluble and among the most 

mobile of all gases and may escape from solution before 

the water reaches the sampled spring or well, producing the soil gas effect noted by Roberts and at the same 

time reducing the dissolved helium level. Helium concentrations of the Appalachian water samples are 

listed in table 4. Water at equilibrium with 5 x 10~6 

atmospheres of He (sea-level value for the normal atmosphere) will contain less than 0.01 /ug/L of the dissolved gas. Thus, all positive findings of He noted in 

the table (i.e., Trace, >0.1^tg/L and greater) indicate 

substantial supersaturation with respect to the atmosphere.

Dissolved carbon dioxide analyses provide an independent check on the field pH and aklalinity determinations. In this study, the gas analysis CO2 values help 

to provide better carbonate system data. Accurate values of carbonate ion activities are obviously essential 

for geochemical studies of water-carbonate mineral 

reactions. Ordinarily, field measurements of pH and 

bicarbonate alkalinity, together with calculations of 

activity coefficients and ion pairing of the major solutes, are used to determine the carbonate ion activities of a water. An alternative method, involving 

calculation of a pH from the analyzed CO2 pressures 

and field alkalinities, was used in determing carbonate 

species activities for several of these Appalachian waters. The calculations are described by Pearson, Fisher, 

and Plummer (in preparation). Carbonate mineral saturation data derived by both methods are listed in 

table 5.

Carbonate ion activities are very sensitive to pH. 

The analyzed CO2-calculated pH method may be advantageous for calculating carbonate ion activities of 

high PCo 2 waters, which may evolve CO2 , thereby increasing pH during pH determinations. We consider 

the errors in analyzed PCOa values to be <5 percent for 

Pco2 >5 x 1°~3 atm- Calculation of pH from the Pco2 

depends on the relationship

co2 + H2o ^ H+ + HCO.;

with an equilibrium constant

aH + &HCO.,

CO-2

CO., "H-,O

Noting that pH = -log aH + , or 10pH = 

and rearranging, we obtain

'CO-, H.,O CO.,

From (A), an error in PCo 2 of 5 percent would correspond to a change of about 5 percent in 10pH, or a pH 

difference of 0.02 units. In contrast, direct pH measurements in ground waters are normally no more precise than ±0.05 units (Langmuir, 1971; Pearson and 

Rettman, 1976).

Recent studies of the CO2 system in the oceans have 

made increasing use of measured PCo 2 rather than pH 

as one of the parameters employed for determinations 

of carbonate ion activity (for example, Li and others, 

1969). A discussion of the relative merits of various 

approaches to these determinations was presented by 

Park (1969).

TRACE ELEMENTS

Dissolved trace elements are sometimes useful qualitative indicators of geothermal waters. Truesdell 

(1976) has reviewed recent geothermal studies in 

which trace element concentrations help to define 

geothermal reservoirs. Arsenic was found in 8 of the 10 

warm springs sampled, indicating a possible correlation between this element and the thermal history of 

the waters. However, the levels of arsenic determined 

(0.5-0.2 /itg/L, the lower detection limit) do not permit 

an assessment of the role of arsenic as a thermal indicator in these waters.

ISOTOPES

Ratios of stable-oxygen isotopes and stable-carbon 

isotopes, and the concentration of the radioactive 

isotope of hydrogen, tritium, were measured in many of 

the samples. Oxygen isotopes are qualitative indicators of the geography of the area of recharge of a
E24 GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

TABLE 5.—Saturation indexes and carbon-isotope ratios for waters in Appalachian warm-spring areas

Sample 

No.

Owner or name Date

Carbonate chemistry based on measured pH

Calcu- Total Saturation index Measured Pr02 bonate Cal- Dolo- 8 I3CPDB 

pH (atmT) (mM/L) cite mite °/oo

Carbonate chemistry base"

Calculated 

pH

Analyzed

Pro, 1

(atm.) (

d on analyzed CO2

Total Saturation index

senate 

mM/L)

CalciteDolomite

Adjusted 813CPDB

°/00

New York — Lebanon Springs area

101 

101 

109 

109 

114 

116

John Koepp

B. Schell __________ _ —

F. Amlaw

-— 10-75

5-76 

.— 10-75

5-76 

.____ 10-75

.— _ 10-75

7.60 0.0038 2.86 -0.11 -0.33 -11.9

7.72 .0029 2.89 +.03 -.19 -11.9 

7.60 .0034 2.74 -.30 -.71 -13.1 

8.00 .0013 2.63 +.07 -.05 -11.9 

7.20 .0093 3.53 -.20 -1.32 —

7.55 .0055 4.23 -.04 -.33 -11.5

7.60 

7.55 

7.87 

7.86

0.0039 

.0043 

.0018 

.0018

2.86 

2.95 

2.66 

2.66

-0.11 

-.14 

-.03 

-.06

-0.32 

-.52 

-.18 

-.32

-11.9 

-12.1 

-12.8 

-12.0

Pennsylvania — Perry County Warm Springs area

201 

201 

202 

216

218

Lloyd Hetrick

M. Loy Spring _ _

.— _ 10-75

5-76 

.— _ 10-75

—— _ 10-75

.____ 10-75

7.20 .0077 2.53 -.59 -2.05 -13.4 

7.62 .0029 2.29 -.16 -1.28 -12.6 

6.85 .0088 1.62 -1.45 -4.01 — 

6.50 .0160 1.73 -2.04 -4.84 -12.9

7.20 .0092 3.34 -.37 -.95 -12.3

7.43

7.48

.0045 

.0040

2.39 

2.34

- 36

-.30

-1.58

-1.56

-13.0

-12.8

West Virginia — Berkeley Springs area

301 

301 

302 

302 

303 

307

Berkeley Springs

F. Wachter

F. Wachter _

H. Hover-male

.— _ 10-75

5-76 

. ___ 10-75

5-76 

______ 9- 75

._ _ 10-75

6.99 .0163 3.35 -.59 -1.85 -15.2

7.46 .0054 2.89 -.12 -.98 -13.9 

6.80 .033 5.78 .34 1.80 15.5

7.07 .0174 4.88 -.27 -1.33 -15.5 

7.30 .0095 4.14 -.24 -.75 — 

6.97 .037 8.35 -.02 -.50 -15.2

6.98 

6.94

.0166 

.0180

3.36 

3.36

-.60-1.86 

-.64-2.02

-15.2 

-14.9

West Virginia — Minnehaha Springs area

401 

401 

402 

402 

403 

406 

407 

410

Minnehaha Springs

Camp Minnehaha

A. T. White ______________

F. Davis _ _

T. Rose _ _ .

.— _ 11-75

6-76

.___- 11-75

6-76

. _ 11-75

.____ 11-75

-—11-75

. _ 6-76

7.88 .0014 1.94 +.04 -.51 -11.4 

7.90 .0013 1.90 +.04 -.54 -11.4 

7.30 .0095 4.24 -.21 -.89 -14.4 

7.30 .0096 4.23 -.16 -.86 — 

7.20 .0045 1.59 -.87 -3.30 — 

7.63 .0022 1.96 -.36 -2.06 -13.1

5.50 .055 3.11 -3.59 -8.10 -22.9

7.68 .0039 4.61 +.28 +.17 +2.1

7.82 

7.72

7.95

.0016 

.0020

.0021

1.95 

1.93

4.50

-.02 

-.13

+ .58

-.62 

-.90

+.71

-11.4 

-11.5

+2.3

Virginia — Bolar Spring area

501 

501

Bolar Spring .— _ 12-75

6-76

7.48 .0063 3.58 +.06 -.25 -9.2 

7.42 .0065 3.28 -.09 -.68 -10.1

7.30 

7.40

.0097 

.0067

3.71 

3.29

-.12 

-.10

-.61

-.72

-9.5

-10.1

Virginia — Hot Springs — Warm Springs area

602 

602 

605 

605 

607A 

619 

619 

625 

627 

631 

631 

633

701 

701 

702

801 

801 

803 

803 

804 

805 

810

Homestead Hot Well

Valley View Cottages

T. K Bond's well

Warm Springs 

(Children's Pool) _ .

Ingall's Airport

Falling Spring

H. D. Webb well ___ ___

Hot Springs

City well _ __ -

Fairview \Vater

Bubbling Springs

P. Lovin

.— _ 11-75

6-76

.— _ 11-75

6-76

—— 11-75

— _ 11-75

6-76

— — 11-75

11-75 

______ 6-76

.— _ 12-75

___ _ 12-75

6-76

.— _ 11-75

.____ 12-75

5-76 

.____ 12-75

5-76 

.____ 12-75

—— 12-75

______ 5- 76

6.65 0.117 10.23 +0.07 -0.05 -5.6

6.55 .147 10.92 -.05 -.32 -5.8 

7.20 .0095 3.28 -.37 -1.35 -8.8

7.15 .0099 3.13 -.51 -1.70 -9.8 

7.15 .022 6.85 +.14 -.01 -7.5 

7.32 .0080 3.56 -.11 -.77 -10.4

7.42 .0058 3.19 -.09 -.90 -7.4 

7:30 .0103 4.57 -.17 -.77 -7.7 

7.48 .0080 5.32 +.26 -.50 — 

7.32 .0101 3.46 +.24 -.21 -8.2 

7.20 .0132 3.50 +.14 -.06 -8.8 

4.50 .055 3.07 -7.0 -14.3 —

Virginia — Falling Spring area

7.32 .0139 5.63 +.41 +.38 -6.8 

7.32 .0109 4.49 +.17 -.16 -8.4 

7.80 .0275 7.80 +.49 +.66 -8.0

North Carolina — Hot Springs area

7.50 .0045 2.15 +.31 +.30 -9.5 

7.50 .0043 2.03 +.35 +.27 -9.2 

6.75 .0179 2.68 -1.61 -3.16 -19.3 

6.88 .0127 2.36 -1.57 -3.06 -19.2 

7.10 .0020 5.75 -2.26 -4.53 -12.0

7.75 .0016 1.84 -.49 -1.07 -13.4 

8.10 .0009 2.25 +.04 +.05 —

6.61 

6.69

6.68

7.26

7.08 

7.10

7.13 

7.10 

6.60

7.28

0.127 

.107

.066

.0134 

.0175 

.0165

.0217

.0177 

.098

.0074

10.48 +0.03 

9.96 +.09

8.58 -.32

5.61 +.03 

3.66 .00 

3.59 .04

5.91 

4.77 

10.09

2.23

+ .22 

-.05 

-.05

+ .10

-0.12 

-.04

-.94

-.95 

-.26 

-.26

.00 

-.60 

-.43

-.12

-5.7 

-5.4

-8.9

-8.5 

-9.0

-7.1 

-8.8 

-9.5

-9.7

Georgia — Warm Springs area

901 

901 

905 

905 

907 

909

Warm Springs

South Spring

J. Phillips

A. Baxley

.__ - 12-75

5-76 

. __ 12-75

5-76 

.— _ 12-75

.____ 12-75

7.40 .0053 2.15 -.48 -.97 -13.7 

6.95 .0146 2.35 -.97 -1.89 -14.5 

4.00 .288 1.22 -7.4 -14.8 -30.1 

4.90 .024 1.05 -6.9 -13.8 -24.3 

6.42 .028 1.82 -1.93 -3.69 -16.4 

7.02 .0081 1.89 -1.22 -2.98 -20.0

7.20 

7.19

.0083 

(.0083)

2.24 

2.17

-.67 

-.73

1 98

-1.40

-14.0 

-14.0
HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS OF THE APPALACHIANS E25

ground water and of whether a ground water has 

reacted with rock at temperatures greater than about 

150°C. Carbon isotopes indicate the types and extent of 

chemical reactions between ground water and carbonate minerals. Tritium has a relatively short half life 

(12.3 years), and nearly all tritium now in the hydrosphere was injected within the last 20 years. The 

tritium content reflects whether a ground water was 

recharged more than about 20 years ago or more recently, and frequently it is possible to use intermediate 

tritium values to estimate the proportions of young and 

old water in a mixed sample.

Stable isotope ratios are reported in 8 notation. 8 

represents the deviation in permil (parts per thousand) 

of the isotope ratio in a sample from that of a standard. 

The 18O/ 16O ratio of the oxygen standard (SMOW) is 

similar to that of ocean water; the 13C/ 12C ratio of the 

carbon standard (PDB) is similar to that of marine 

limestone. Tritium concentrations are reported in 

tritium units (TU). One TU equals 3.2 picocuries per 

liter (pCi/L).

Waters on the land surface are generally depleted in 

18O relative to SMOW—that is, their 8 18O values are 

negative. This depletion occurs because the vapor pressure of water molecules containing the heavier isotopes 

is slightly less than that of common water, H2 16O. During evaporation and condensation in the hydrologic cycle, molecules containing heavier isotopes are concentrated in the liquid phase. As water evaporates from 

the ocean, the vapor is depleted in 18O and the amount

of depletion becomes greater as the temperature of 

evaporation decreases. Further isotopic fractionation 

takes place as water is condensed and reevaporated 

during atmospheric transport. The amount of fractionation is inversely proportional to temperature, and also 

varies with altitude, latitude, and distance from the 

ocean. 

Figure 21A shows the variation of 818O values of

samples from this study with mean annual temperature. The lines on the figure have a slope of 0.5 permil/ 

°C, within the range of 0.4 to 0.7 permil/°C found in 

more extensive studies elsewhere (Dansgaard, 1964). 

The two lines on the figure represent the seasonal 

range in 8 18O values to be expected. Samples collected 

in the spring are characterized by more negative 818O 

values which reflect the lower temperature winter precipitation; the less negative 818O values of the autumn 

samples reflect higher temperature summer precipitation.

Oxygen isotope ratios of ground water can be 

changed by chemical and isotope exchange reactions 

between the water and the aquifer rock. Because the 

8 18O value of rock-forming minerals is strongly positive, reactions between water and rock will make the 

18O values of the water more positive. Such reactions 

occur only very slowly at temperatures below 150- 

200°C. Thus, anomalously positive 8 18O values may indicate that a ground water has been exposed to high 

temperatures (White and others, 1973).

As figure 2L4 shows, the 8 18O values of the thermal

CO

D

CO 20

_l

w 

o

1R16

EC

UJ 8

EL

EXPLANATION

O Warm springs

x Wells and cold springs

Spring Autumn

Per

Autumn, 1975

Warm Springs area, Ga

Hot Springs area, N.C. 

ry County area, Pa.'

Other Va. and W. Va. 

spring areas

Berkeley Springs area, W. Va. 

Lebanon Springs area, N.Y:

S 18 0SMOW IN PERMIL

I I

30

34

38

42

46

40 60 80

TRITIUM UNITS

FIGURE 21.—Isotopes in Appalachian warm springs area waters. A, Mean annual temperature and variation of 818 O values of samples.

B, Latitude and variation of tritium concentration of samples.
E26 GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

spring samples are within the ranges of values of the 

cold wells and springs in the same areas, allowing for 

the seasonal variations. This is evidence that the areas 

of recharge to the thermal spring systems are within 

kilometers or at most tens of kilometers from their discharge points. Further, it provides no evidence that 

would suggest that the thermal spring waters have 

been as hot as 150-200°C.

Figure 2 IB shows the variation of the tritium concentration of samples from this study with latitude. 

The line on the figure is drawn through the highest 

tritium values of samples collected in the autumn of 

1975. The increase in tritium with latitude, by a factor 

of roughly 3 from 32°-42°N, and the tritium values 

themselves are what one would expect to find in ground 

waters which are a mixture of waters recharged since 

the period of major tritium injection began in the early 

1960's (Stewart and Farnsworth, 1968; Stewart and 

Wyerman, 1970). The line may therefore be taken to 

represent the approximate tritium content of shallow, 

recently recharged ground waters at these latitudes.

Samples from the thermal springs at Warm Springs, 

Ga., Hot Springs, N.C., and Hot Springs and Warm 

Springs, Va., contain no tritium. Water from these 

springs is thus at least 20 years old. Tritium in the 

thermal springs at Bolar and Falling Spring, Va., was 

about 25 TU in 1975. From the line in figure 215, the 

maximum tritium content of shallow ground water at 

the latitude of these springs can be estimated to be 64 

TU. If the tritium content of the thermal component of 

these springs is zero, the proportion of thermal (T) to 

shallow, cool (C = 1 —T) water is:

25 = T(0) + 64(1 -T) 

T = (64 - 25)/64-60%.

The remaining thermal springs contained from 2 to 12 

TU in 1975, a level corresponding to a thermal-water 

fraction of from 97 to 81 percent.

CARBONATE CHEMISTRY AND ISOTOPES

The carbonate chemistry of ground waters—whether 

the waters are in equilibrium with carbonate aquifer 

minerals—and the stable carbon isotope ratios of dissolved carbonate, give qualitative indications of water 

residence times and flow paths. Enough information is 

available on the warm and cold-spring samples for this 

study to permit a detailed interpretation of the carbonate chemistry of most of the spring systems.

Table 5 shows both the analyzed PC02 and PC02 values 

calculated from the pH and other chemical analyses. In 

some samples, such as the 1975 samples 101, 301, 401, 

there is good agreement between the calculated and

analyzed PC02 values. In others, such as 1975 samples 

631 and 702 and 1976 samples 301, 701, and 602, the 

calculated PC02 values are lower than the analyzed 

values.

Saturation indices (SI) for the minerals calcite and 

dolomite were also calculated and are given in table 2. 

Both PC02 and SI values were calculated using 

WATEQF (Plummer, and others, 1976). SI values calculated from measured pH and alkalinity values range 

from great undersaturation (SIcaicite<-3) in hightritium, high-Pc02 recharge samples to apparent oversaturation (Slcaicite > + 0.2) in such warm-water samples as 631, 701, 702, and 801. In these apparently 

oversaturated samples, however, PC02 calculated from 

pH and alkalinity is lower than that analyzed.

No matter how carefully a field pH is measured, 

there is always a possibility that some CO2 outgassing 

will occur during collection and the measurement procedure, causing an erroneously high pH to be read. To 

test this possibility, the saturation index calculations 

were repeated assuming that the measured Pco2 values 

were correct (Pearson, and others, in prep.). The results 

of these calculations are also shown in table 5.

Consider the samples of 631 as an example. On the 

basis of measured pH, both samples are oversaturated 

with calcite and have lower calculated than analyzed 

PCo2 When recalculated from the analyzed PC02 ' S > however, both samples are just saturated with calcite and 

have lower calculated than analyzed pH's. The unchanging measured temperature and the identical 

(within measurement precision) PC02 and tritium values suggest that the chemistry of water from 631 

should also be constant. The recalculated SIcaicite and 

pH values better reflect a constant chemistry than do 

the measured pH values and the SI values calculated 

from them. Thus the Si's based on analyzed PC02 probably best represent the true state of the water chemistry.

As water recharging a ground-water system passes 

through the soil zone, it dissolves CO2 from the soil 

atmosphere. Soil CO2 is a product of plant respiration 

and decay, and it has 8 13C values near — 25°/0o- If this 

water then passes into a marine limestone or 

sandstone with carbonate cement of 8 13C~0°/00, the 

limestone or cement will dissolve in the soil CO2 - 

charged water and produce bicarbonate with 8 13C values as positive as -10 to -12°/00 - In waters in rock 

where marine carbonate is plentiful, further reactions 

between rock and water may make the 8 13C value even 

more positive. Thus the 8 13C content of a ground water 

is a measure of the extent of reaction between the 

water and carbonate rock.

The measured 8 13C values are given in table 2, and in 

Hobba, Chemerys, Fisher, and Pearson (1977) where
HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS OF THE APPALACHIANS E27

the method of collection and analysis is also described. 

If CO2 gas has been lost from a sample before or during 

the collection of the 8 13C sample, the analyzed values 

will not be a true measure of the 8 13C value of the 

source water before outgassing because of the isotopicfractionation factor of several permil between CO2 gas 

and dissolved carbonate species. If the PC02 of the 

source water is known, as from a dissolved gas 

analysis, the correct 813C can be calculated from the 

value measured on an outgassed sample (Pearson and 

others, in prep.). The 813C values of those samples on 

which the activities of carbonate species were calculated from analyzed PCo2 values were also adjusted for 

the effects of outgassing. The adjusted 8 13C values are 

given in table 2.

The SIcaicite values, based on analyzed PC02 where appropriate, and the 8 13C values reflect the types of rock 

through which the various waters have traveled. Some 

of the cold spring samples, including numbers 407, 633, 

and 905, have the chemical characteristics of soil-zone 

waters. They have high PC02 values, pH values lower 

than 6.0 and 813C values about -25°/00 . Their tritium 

contents show that they were recently recharged. As 

such waters move deeper into a ground-water system, 

they encounter and dissolve limestone and dolomite. 

This process lowers the PC02 , raises the pH, and makes 

the 8 13C values become more positive. Water long in 

contact with carbonate rock reaches saturation with 

calcite; however, water in contact with carbonate rock 

for only a short time, or in rock containing little carbonate, may be undersaturated.

The samples reflect this pattern. Thermal waters issuing from sandstone or quartzite in Perry County, 

Pa., Berkeley Springs, W. Va., and Warm Springs, Ga., 

are all undersaturated with calcite, though their low 

tritium contents suggest a sufficiently long residence 

time for saturation with calcite to have been reached if 

they had been exposed to carbonate rock. Samples that 

are associated with carbonate rock (such as 109, 602, 

631, and 801) and that contain little or no tritium are 

saturated with calcite. Those which have significant 

tritium contents and which represent either relatively 

recent waters or mixtures of waters of short or long 

residence times may or may not be saturated with calcite.

CHEMICAL GEOTHERMOMETERS

The measured temperature of a thermal spring or a 

hot-water well provides a direct indication of a subsurface heat source. However, thermal reservoir and 

discharge-water temperatures may not be equal; heat 

loss by conduction can lower the temperature of the 

thermal water considerably below that of the heat 

source. A number of geothermometers have been devised for estimating reservoir temperatures directly 

from the chemistry of thermal water samples (Truesdell, 1976). Successful application of geothermometers 

depends on the attainment of equilibrium in a temperature dependent water-rock reaction in the thermal reservoir, and no subsequent change in thermal water 

composition. The assumptions inherent in the use of 

geothermometers are discussed in Fournier, White, 

and Truesdell (1974).

Mixing of hot water with shallow, cold water often 

occurs in thermal spring systems. The effects of the 

admixture of cold water must be taken into account 

before estimates of reservoir temperatures can be 

made. In general, the temperature of the mixed water 

will be determined by the enthalpies of the hot and cold 

components (Fournier and Truesdell, 1974). At temperatures below boiling, the enthalpy of water (in 

calories per gram) is numerically equal to its temperature (Tobserved) of the mixed-water spring will be equal 

to the mean of the fractions of cold (Fcold ) and hot (Fhot) 

water times the respective temperatures, or:

-cold cold + Thot

_ m

hnt — I,observed

where

+ Fhot = 1.

(1)

(2)

The mixed water will also contain an intermediate 

concentration (C) of the geochemical indicator, so that:

'cold cold hot 'analyzed (3)

The indicator concentration and the temperature of the 

cold water can be estimated for favorable cases. The 

geothermometer relationship between temperature 

and indicator concentration of the hot water

(4)

then provides a unique solution of the equations for the 

mixing fractions and the temperature of the thermal 

reservoir in systems for which conductive heat loss is 

negligible. Alternatively, if the concentration of any 

conservative constituent of the hot water is known, and 

if this concentration differs significantly from that of 

the cold water component, then the mixing fractions 

can be determined and a hot water temperature may be 

calculated directly from the first three equations, and a 

reservoir temperature may be estimated from available geothermometers. An example of this alternative 

scheme for calculating mixing ratios is the use of the 

assumption that tritium in the hot component should
E28 GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

be negligible. Then the second term in equation 3 is 

zero, and the cold fraction is readily calculated.

The oldest and among the most generally useful of 

the chemical geothermometers make use of the silica 

content of the sample. Reactions between water and 

common alumino-silicate minerals are frequently of 

the form:

cation-Al-silicate + H2O

-» SiO2<tiiss<>ived> + cation + Al-Silicate(s)

where the reactant cation-Al-silicate may be a sodium, 

potassium, or calcium feldspar or clay mineral and the 

product Al-silicate is another clay mineral. The result 

of such reactions is to increase the dissolved silica content of the water until a silica mineral such as quartz 

(at relatively high temperatures) or chalcedony (at

lower temperatures) precipitates. The degree to which 

chemical equilibrium between dissolved silica and the 

precipitating mineral is approached depends on the 

temperature and the time available for the reaction to 

proceed. The silica geothermometer assumes attainment of equilibrium between dissolved silica and a particular silica mineral at the reservoir temperature, 

with no precipitation of dissolved silica after the water 

has left the heated reservoir. Determination of the 

silica content of the water sample will then indicate 

the thermal reservoir temperature, provided the 

geothermometer has been properly calibrated for the 

appropriate form of silica involved in the reaction. Figure 22 is a plot of dissolved silica concentrations versus 

observed temperatures for the Appalachian spring and 

well samples. The curves represent concentrations of 

dissolved silica in equilibrium with quartz and with 

chalcedony at the indicated temperatures. They are derived from silica solubility data determined by a

80 -

70-

60 -

50 -

GC An —

« 30 -

20-

10 -

EXPLANATION

. Spring

x Well, <25% modern water

A Well, > 25% modern water

10 20 30 40 50 60 70

TEMPERATURE, IN DEGREES CELSIUS

80 90 100

FIGURE 22.—Dissolved silica content and temperatures of Appalachian spring and well samples.
HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS OF THE APPALACHIANS E29

number of workers and reported by Fournier (1973).

From the plotted points of figure 22, all the Appalachian spring samples are seen to be supersaturated 

with quartz; 2 of them (the North Carolina warm 

spring samples) are supersaturated with chalcedony. 

Well water data illustrated in figure 22 were divided 

into 2 age groups according to sample tritium content 

relative to the expected tritium level of fresh recharge 

in the particular geographical area (from fig. 2IB.). 

Dissolved silica concentrations of 8 of the high-tritium 

well waters are above chalcedony saturation, indicating rapid low-temperature dissolution, by fresh recharge, of silica from shallow earth materials. Presumably little or no re-precipitation occurs because of 

slow equilibration rates for the more stable silica phases. On the other hand, the dissolved silicatemperature relationship for the older well waters is 

similar to that for the springs; a general increase in 

dissolved silica with increasing sample temperature is 

evident from the figure. Indicated quartz and chalcedony temperatures of the samples are included in 

table 4.

Another geothermometer that has been applied with 

fair success in thermal systems is the Na-K-Ca or cation thermometer (Fournier and Truesdell, 1973). Unlike the silica geothermometers, which are based on 

laboratory measurements of silica solubilities, the cation geothermometer is empirical and the particular 

mineral-water reactions that make it work are not 

known.

The formulation of this geothermometer for waters 

of less than 100°C and with a molar ratio VCa/Na 

greater than 1 is:

t(°C) -

1647

log(Na/K) + (4/3) log (VCa/Na) + 2.24 - 273.15

where Na, K, and Ca are the concentrations of sodium, 

potassium, and calcium, respectively, in moles per liter.

Cation temperatures are also listed in table 4. Temperatures calculated on samples with low Ca, Na, and 

K concentrations have large errors associated with 

them because of the imprecision of analysis at low concentrations. Temperatures calculated for some samples 

with higher Ca, Na, and K concentrations are negative, showing that the waters have not reached the 

minimum temperature necessary for the reactions that 

give rise to the geothermometer effect to become active. Because the cation geothermometer is empirical, 

this minimum temperature cannot be given unequivocally. Several of the warmest samples, including

those from Warm Springs, Ga. (no. 901), Hot Springs, 

N.C. (no. 801), and Warm Springs and Hot Springs, Va. 

(no's. 631, 607, and 602), have cation thermometer 

temperatures in reasonable agreement with both 

measured and chalcedony temperatures. The coolest of 

these are samples 607 A and B, which are about 20°C. 

Other samples, however, such as 301 and 702, with 

measured temperatures above 20°C have negative cation temperatures. Because all samples above 30°C do 

have reasonable cation temperatures, this might be a 

minimum temperature for reliability of this geothermometer.

At relatively low temperatures (<50°C), geothermometer estimates will not be particularly accurate 

because the basic premise behind their operation may 

not be valid, that is, equilibrium of the temperature 

dependent reaction may not be attained. However, fair 

success of the chalcedony and cation thermometers is 

indicated by the data for the warmest of the Appalachian spring systems. This agreement between chalcedony, rather than quartz, with measured and cation 

temperatures, together with the relatively low (<50°C) 

maximum temperatures suggests that chalcedony 

rather than quartz is controlling the silica concentrations in these systems.

DISCUSSION

SILICA-ROCK SPRING SYSTEMS

The Perry County, Pa., Warm Springs (201) and 

Berkeley Springs (301) and Minnehaha Springs (401), 

W. Va., issue from the Oriskany Sandstone; and the 

Georgia Warm Springs (901), from the Hollis 

Quartzite. Certain chemical characteristics of these 

springs are similar because of their similar aquifer 

rock type, and it is convenient to discuss them together. All have relatively low concentrations of dissolved solids—181 to 258 mg/L—as compared to the 

values of 266 to 848 mg/L in waters from carbonate 

rocks. A trace of dissolved helium was found in only 

one sample from Minnehaha Springs (401), and none 

was found in any of the other three silica-rock warm 

springs. Generally higher helium levels were found in 

samples from the carbonate warm springs and hot 

wells.

The four silica-rock thermal springs contain little 

(201) to no (901) tritium, and all but Minnehaha 

Springs (401) are significantly undersaturated with 

calcite and have 813C values of —130/00 or more negative. These undersaturated waters could not have been 

exposed for any extended period to aquifer rock containing carbonate minerals, and so they must have 

flowed from recharge to discharge through the lowcarbonate sandstones from which they issue. Although 

Minnehaha Springs (401) issues from the Oriskany
E30 GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

Sandstone, the water is saturated with calcite. This 

spring is in the vicinity of the Virginia thermal 

springs, all of which are associated with limestone. The 

fact that Minnehaha Springs is saturated with calcite 

suggests that, unlike the other sandstone springs, its 

water has been exposed to carbonate rock.

PERRY COUNTY, PA., AREA

Waters from warm spring 201 and samples 202 and 

216 are derived from the Devonian Oriskany 

Sandstone. Spring 218 issues from the Silurian Tonoloway Limestone. The sample from spring 218, with 

a tritium content of 78 TU (fig. 22 and table 4), is 

virtually entirely modern water, and the concentrations of chloride, nitrate, and sulfate suggest the 

possibility of human influence on the water composition. The Oriskany Sandstone samples are all dilute 

calcium bicarbonate waters. The only pronounced 

compositional differences between the warm spring 

and the other two waters are the slightly elevated 

silica content and the generally higher mineralization 

of the warm spring, suggesting a longer period of flow 

in the aquifer for this sample. The quartz geothermometer temperature for the 1975 sample from spring 

201 is probably a conservative estimate of maximum 

temperatures at depth in the aquifer.

BERKELEY SPRINGS, W. VA., AREA

The warm springs issue from the Oriskany 

Sandstone. Well 302 taps both the Oriskany and the 

overlying Marcellus Shale, whereas well 303 is entirely in the shale. Sample 307 is from a well completed 

in the Tonoloway Formation. The waters are of calcium 

bicarbonate and calcium-magnesium bicarbonate 

types, with relatively high sulfate in waters from the 

shale. The thermal-spring water is the most dilute 

water of those sampled in the area; sample 307, which 

from its tritium content of 72 TU appears to be all 

modern water, is the most highly mineralized sample.

Silica temperatures are highest for wells 302 and 

303 in the Marcellus Shale. These waters are supersaturated with chalcedony at the observed temperatures (fig. 22). However, recharge to the shale in this 

area is local and the measured tritium levels suggest 

that these samples contain about 50 percent recent recharge. Measured temperatures do not indicate 

significant local heating of the rock; therefore, the high 

silica contents in these samples very likely result from 

silica from weathering reactions having had insufficient time to equilibrate with the more stable chalcedony phase.

Indicated cation thermometer temperatures are all 

much lower than observed temperatures; however,

predicted temperatures for the warm springs are the 

highest in this group of samples. The relatively high 

potassium content of the spring water appears to be the 

only significant chemical indicator that may be associated with the source of heat. There is little likelihood 

that temperatures at depth in this aquifer system 

would exceed the indicated temperature of about 38°C 

for the warm spring.

MINNEHAHA SPRINGS, W. VA., AREA

The warm springs in this area flow from the Oriskany Sandstone. The calcium bicarbonate compositions of the waters are generally similar to those in 

the Perry County and Berkeley Springs areas, except 

that sulfate constitutes almost 30 percent of the anion 

equivalents in the Minnehaha warm spring samples 

401. Well 402 is completed in the Hamilton Formation, 

stratigraphically above the Oriskany. Samples 403 and 

407 from the Oriskany Sandstone contain high (>20 

mg/L) nitrate concentrations, and well 407, with a 

tritium concentration of 88 TU, is modern water. Compositions of these well waters are very likely affected 

by human activities. Spring 406 issues from the Tonoloway Formation. Its tritium level of 63 TU indicates 

that this calcium bicarbonate water is also recent.

Highest silica concentrations in the Minnehaha 

Springs area samples are found in well 402. The water 

is near equilibrium with chalcedony at the observed 

water temperature, whereas warm spring water 401 is 

considerably below chalcedony saturation. As is the 

case in the Berkeley Springs area, the stratigraphically and hydraulically higher formation contains 

young (high tritium) water, which according to the 

silica geothermometers, has been heated more than the 

thermal spring water in the area. Cation temperatures 

for spring 401 and well 402, although much lower than 

observed temperatures, do correspond with the observed trend in temperature. Levels of potassium in 

the warm-spring and well waters are similar to those 

noted at Berkeley Springs. Also the observed trace of 

dissolved helium in the 1975 warm spring sample suggests that the spring water has undergone some heating.

High temperatures at depth in the aquifer do not 

appear likely; the quartz thermometer temperature 

indicated for the warm spring water (about 34°C) is 

probably a conservative estimate of maximum temperature for water in the Oriskany Sandstone in this area.

WARM SPRINGS, GA., AREA

The warm spring (901) and cold spring (905) flow 

from the lower Paleozoic Hollis Quartzite. Hot well 907 

penetrates the younger Manchester Schist, but it may
HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS OF THE APPALACHIANS E31

also draw water from the underlying Hollis Quartzite. 

Well 909 is completed in the Carolina(?) Gneiss of 

former usage, which is older than the Hollis Quartzite.

Spring 905 yields extremely dilute, modern water 

with a composition unlike that of the warm spring. 

Chemically, the warm-spring and hot-well waters are 

similar. Both are waters of calcium-magnesium bicarbonate type, and relatively high in dissolved oxygen. 

The concentration of dissolved solids in the spring 

water is slightly higher than that in the hot well water 

(table 4). Well 909 differs from the warm spring and 

hot well in that sodium amounts to 25 percent of its 

cation equivalents, and dissolved oxygen is very low.

Tritium analyses indicate that spring 905 is almost 

entirely modern water, and that water from well 909 is 

approximately half young water. Warm spring 901 and 

hot well 907 do not contain significant components of 

modern water.

Chalcedony- and cation-thermometer temperatures 

are in fair agreement with each other and with the 

observed temperatures for the warm spring and hot 

well. The hot well probably draws water from the same 

source that supplies the spring. The agreement between geothermometer and observed temperatures for 

these waters indicates that the reservoir temperature 

in the warm spring system is approximately equal to 

the surface temperature of the water; little heat has 

been lost by conduction, and no evidence of mixing is 

apparent from the data.

Sample 909 geothermometer temperatures are also 

in good agreement with each other, but they are considerably higher than the observed temperature. The 

agreement between the chalcedony and cation geothermometers may be fortuitous; high dissolved silica 

may result from low-temperature surface weathering 

reactions, and mixing of two cold waters can result in 

mixed-water compositions from which high cation 

temperatures would be predicted. If well 909 is indeed 

a mixture of hot and cold water which has been cooled 

conductively, the hot reservoir temperature can be calculated as follows. From the tritium content of this 

sample (15 TU) and that estimated for shallow ground 

water in the area (35 TU, fig. 2LB), the sample appears 

to contain about 45 percent cool, shallow water. If the 

silica content of this water is 7.0 mg/L (from samples 

905 representing high-tritium, local, shallow ground 

water), the silica content of the hot water can be calculated to be about 50 mg/L, which corresponds to a chalcedony temperature of 70°C for the hot component. A 

similar calculation for the cation temperature of the 

hot component of sample 909 gives 53°C, which is 

below the chalcedony temperature. This suggests that 

deep waters hotter than 60-70°C probably do not exist 

in this area.

CARBONATE-ROCK SPRING SYSTEMS 

LEBANON SPRINGS, N.Y., AREA

The spring and wells sampled in the Lebanon 

Springs area all tap the Stockbridge Limestone of 

Early Cambrian to Early Ordovician age. Well 114 also 

draws water from the overlying Walloomsac Slate. 

Samples 101, 109, and 116 are calcium-magnesium 

bicarbonate waters; sample 114 is of the calcium sulfate type. The close similarity in chemical composition 

of waters from the thermal spring (101) and artesian 

well 109 suggest that they may be derived from the 

same source. Seasonal chemical variability in spring 

101 probably results from changes in proportions of 

modern water feeding the spring, as indicated by the 

tritium analyses (table 4). This temporal variability 

has diminished, tritium has decayed, and heat has 

been lost by conduction by the time the water reaches 

well 109.

Wells 114 and 116 were sampled to provide data for 

comparison with the spring analyses. These waters 

were thought to be typical of shallow wells in the 

Stockbridge Limestone. However, sample 114 is a calcium sulfate water, unlike that of the spring. Sample 

116 is chemically similar to the warm spring water, 

although relatively higher in total dissolved minerals 

and depleted in dissolved oxygen and in tritium; the 

chemical composition indicates that this water has 

been in the aquifer longer than the spring water. On 

the other hand, spring 101 and well 109 samples are 

higher in potassium, chloride, bromide, and strontium 

than is the well 116 water, and the samples contain 

detectable amounts of helium. These differences may 

be associated with the path of circulation and the 

source of heat for the spring. The amount of dissolved 

helium in the warm spring, - 1/^g/L, indicates that the 

water may have been heated by as much as 35°C.

Neither the silica nor the cation geothermometer indicates a high reservoir temperature for any of the 

samples from the Lebanon Springs area. If the spring 

water is assumed to consist of a mixture of a heated, 

quartz-saturated old component with 13 percent of a 

cold, modern water with negligible silica, then the 

silica content of the undiluted thermal water would be 

1 -h (l-0.13)x 12 = 14 mg/L SiO2 , and the corresponding 

quartz temperature would be 51°C. (The estimate of 13 

percent is based on the measured tritium content of 

11.6 TU, as compared with the content of 90 TU estimated for fresh recharge; see fig. 2LB.) This value of 

temperature is probably a reasonable upper limit on 

aquifer temperatures at depth in the area.

The carbonate chemical parameters are in keeping 

with the source rock of these samples. Stockbridge 

Limestone samples 109 and 116, with no tritium and 

hence an age greater than 20 years, are saturated with
E32 GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

calcite and have 8 13C values suggesting that their dissolved carbonate is about an equal mixture of soil-air 

CO2 (813C^O25°/00 ) and marine limestone (813OO°/00 ). 

Lebanon Springs itself (101) and sample 114 may be 

slightly undersaturated with calcite. Lebanon Springs 

contains tritium, and it is thus an admixture of water 

too young, perhaps, to have reached calcite saturation. 

Sample 114 draws water, in part, from the Walloomsac 

Slate, which may contain too little calcite to permit 

calcite saturation to be attained.

BOLAR SPRINGS, WARM SPRINGS, HOT SPRINGS, 

AND FALLING SPRING, VA., AREAS

The warm springs sampled in these areas flow from 

the Ordovician Lowville Limestone; sample 633 is from 

the Silurian rocks of Clinton age; the other cold springs 

and cold wells sampled draw water from the Lowville 

Limestone, or from the Trenton Limestone, and Martinsburg Shale or rocks of Black River age, which include the Lowville Limestone. Hot wells 607 and 702, 

completed in the Lower Ordovician Beekmantown 

Group, draw water from elevations below those of 

nearby springs in the Lowville Limestone. Sample 

607A represents a mixed water consisting of components from two or more producing zones. The February 

1976 samples from well 607, on the other hand, were 

obtained from isolated intervals; sample 607B represents the upper aquifer water and sample 607D is from 

the deepest water-bearing zone.

Except for the extremely dilute sample 633, dissolved solids of the waters sampled in the four areas 

ranged from 288 to 1,260 mg/L (table 4). Calcium and 

magnesium account for nearly all the dissolved cations; bicarbonate and sulfate in varying proportions 

are the major anions.

Small amounts of dissolved helium were found in one 

or more of the samples taken from the warm springs 

and hot wells in the four areas and in one sample from 

spring 619. The concentration of helium in hot well 

702, about 2 /zg/L, suggests that this water may have 

been heated by as much as 70°C. Tritium in warm 

springs 501 and 701 indicates that these samples contain one-third to one-half modern water; presumably 

the temperature of the older component may be substantially higher than the observed temperatures of 

the flowing springs.

The carbonate chemistry of these thermal springs is 

in keeping with their origin from a carbonate rock. All 

are saturated with calcite, and their 8 13C values are 

more positive than —10°/00 - Calcite saturation and 8 13C 

values in this range are typical of waters of moderate 

to long residence times in limestone.

Boiler Spring (602), which is saturated with dolomite 

as well as with calcite, has a 8 13C value more positive

than — 6°/00. 813C values in this range, with calcite and 

dolomite saturation, suggest either that the Boiler 

Spring water had a longer residence time than that of 

the other springs or that the other springs are mixtures 

of a Boiler Springlike water with a younger water.

Comparisons of observed temperatures with silica 

thermometer temperatures for the warmest springs 

(602 and 631) and for hot well sample 607D indicate 

slight undersaturation of the waters with chalcedony. 

Cation thermometer temperatures for these samples 

are in fair agreement with observed temperatures, and 

here, as in other warm spring areas, the higher cation 

temperatures are associated with relatively elevated 

potassium levels in the waters. Relatively high cation 

temperatures are indicated for spring 701; nearby hot 

well 702, with an observed temperature higher than 

that of the spring, has a predicted cation temperature 

of <0°C. On the other hand, indicated silica temperatures for these samples are in the reverse order. The 

modern water component of the hot-well sample, indicated by its tritium content, apparently influences the 

mixed-water chemistry enough to invalidate the cation 

thermometer results.

Samples 701 and 702 can be interpreted as mixed, 

conductively cooled waters, and by making certain assumptions, the maximum probable temperature of 

their hot component can be calculated. Thus, if it is 

assumed that the tritium content of the thermal component is 0 TU and that of the local cold ground water 

in the autumn of 1975 was 64 TU (fig. 2IS), the tritium 

contents of samples 701 and 702 suggest that the samples contain about 60 and 85 percent of a thermal component, respectively. If it is further assumed that the 

silica content of the cold ground water is 7 mg/L, the 

silica content of the hot component can be calculated 

using the expression

— 7 x proportion cold

proportion hot

The silica content of the hot component of both samples 

calculated this way is about 26 mg/L; this figure corresponds to a chalcedony equilibrium temperature of 

40°C. It is not possible to calculate a cation temperature of the undiluted thermal component because values for the calcium, sodium, and potassium contents of 

the cold component are not known. However, the cation 

geothermometer expression (above) includes the cation 

concentrations as ratios and so mixing with a presumably dilute water should not significantly change the 

calculated temperature. Thus, for samples 701, at 

least, the cation temperature supports the conclusion 

from the calculated chalcedony temperature that the
HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS OF THE APPALACHIANS E33

maximum temperature of the thermal component is 

about 40°C.

The occurrence of large warm springs issuing from 

the Lowville Limestone over a distance of more than 15 

miles, and the good yields of warm water from the two 

wells in the Beekmantown nearly 10 miles apart suggest the existence of an extensive warm reservoir in 

this part of the Appalachians. Boiler Spring (602) and 

the warm component of the mixtures 701 and 702 have 

concordant chalcedony, cation, and measured temperatures that suggest this reservoir is at about 40°C.

HOT SPRINGS, N.C., AREA

Waters sampled in the area are from rocks of Cambrian age. Warm spring 801 and well 810 are in the 

Shady Dolomite. Spring 805 flows from the overlying 

Rome Formation, whereas wells 803 and 804 are completed in the older Murray Shale and Hesse Quartzite, 

respectively. Although the warm spring is saturated 

with dolomite, its gross chemistry is predominantly of 

the calcium sulfate type, unlike the water from Shady 

Dolomite well 810, which contains equivalent amounts 

of calcium and magnesium as major cations and bicarbonate as the predominant anion. Spring 805 is similar 

in chemical type to the water from well 810. The warm 

spring water, in addition to being of a different chemical type, is also more highly mineralized than are the 

other samples from the area (table 4).

Well 803 water is unusual in that iron accounts for 

nearly 18 equivalent percent of the cations. Highly reducing conditions in the aquifer are indicated by the 

presence of measurable concentrations of dissolved 

methane in the water. On the other hand, measurable 

oxygen was present at the time of sampling, as shown 

by determinations in the field of dissolved oxygen, and 

the high tritium concentration indicates a modernwater component of about 80 percent. Apparently, 

when this well is pumped, a large amount of shallow, 

modern, oxygenated water is drawn up with the highly 

reducing older water in the shale, resulting in an unstable solution containing ferrous iron and dissolved 

oxygen. The oxygen is rapidly consumed, producing a 

voluminous precipitate of ferric hydroxide within a few 

minutes after sampling.

Tritium levels indicate that samples 804 and 810 are 

nearly entirely recent water. Spring 805 contains a lesser component of young water, while the warm spring 

(801) contains only relatively old water. Small quantities of dissolved helium were found in the warmspring water but not in any other samples from the 

area. The higher helium concentration (from the sample collected in December 1975) may indicate heating 

of the water by radioactive decay processes to 24°C 

above normal ground-water temperatures.

The young waters from wells 804 and 810 are undersaturated with chalcedony; older water from cold 

spring 805 is near saturation. Warm spring 801 and 

well 803 water, on the other hand, are supersaturated 

with respect to chalcedony. The relatively low flow 

(—0.5 L/s) of the warm spring, as well as the observed 

seasonal variability in its temperature (table 4 and fig. 

17), are consistent with the possibility of appreciable 

conductive cooling taking place between the heat 

source and the site at which the spring flows. If such 

conductive cooling occurs, then the water temperature 

near the heat source may be considerably greater than 

the measured outlet temperature, quite possibly as 

high as 50°C (equilibration with chalcedony). Cation 

thermometer temperatures suggest that the chalcedony temperature is the probable reservoir temperature.

If mixing rather than conductive cooling accounts for 

the observed chalcedony supersaturation of warm 

spring 801, then a hot reservoir temperature can be 

calculated from the equations described above, using 

appropriate values of temperature and dissolved silica 

for the cold component and assuming chalcedony saturation of the hot water component. Solution of the 

equations requires an analytical expression for the 

chalcedony-temperature curve. An equivalent graphical solution can be obtained, as illustrated in figure 22. 

Point A represents the December, 1975, warm spring 

sample. The cold component, at equilibrium with chalcedony at a temperature of 14.2°C, is indicated as point 

B. Tie-line BA which extends to the high-temperature 

intersection with the chalcedony saturation curve at 

point C, gives the temperature and dissolved silica content of the hot water component. This procedure yields 

a reservoir temperature of 77°C and silica content of 57 

mg/L for the hot water.

Interpretation of the thermal history of the water 

from well 803 is difficult. Quite possibly the high dissolved silica level derives from a lack of time for the 

silica from weathering reactions in the high-tritium 

cool water to reach equilibrium with chalcedony as 

seems to have occurred at the Berkeley Springs wells 

in the Marcellus Shale. The cation geothermometer 

could not be expected to predict accurate temperatures 

for the formation water because cation concentrations 

and ratios will be strongly influenced by the modern 

component of the mixed-water sample. The tritium 

content of water from well 803 can be interpreted as 

indicating that the 1975 sample contains about 20 percent hot and 80 percent cold water, and that the cold 

component contains 8 mg/L SiO2 . If this were the case, 

the undiluted hot component would have a silica content of about 90 mg/L, corresponding to a chalcedony 

temperature of 105°C. These temperatures exceed the
E34 GEOHYDROLOGY OF GEOTHERMAL SYSTEMS

concordant chalcedony and cation temperatures of 

sample 801 too much to be entirely credible, and it is 

thus likely that the high silica content of sample 803 is 

due to supersaturation with chalcedony at low temperature rather than to equilibrium at high temperature. 

Thus, in the Hot Springs, N.C., area, reservoir temperature predicted by the chalcedony geothermometer is 

77°C if the warm spring is a mixture of hot and cold 

waters, or about 50°C if the spring is conductively 

cooled. Cation geothermometer results suggest that 

the lower temperature is the more likely.

CONCLUSIONS

All evidence accumulated thus far indicates that the 

heat at the thermal springs investigated is derived 

from rocks at depth in a region of "normal" geothermal 

heat flow. The spring waters are mixtures of warm, 

deep ground water and shallow, cool ground water in 

variable proportions. In general, the warmest springs 

discharge the highest proportion of deep-circulating 

warm water.

Many thermal springs are in valleys on the crests or 

flanks of anticlines where steeply dipping beds of 

sandstone or limestone intersect faults or lineaments. 

Some thermal springs are in areas of complex geology 

and are associated with faulted limestone or sandstone.

Deep circulation of most of the ground water discharging in the thermal springs is in the general direction of the strike of the rock strata, along open tension 

fractures, bedding planes, and faults. Some deep circulation possibly occurs also along fault planes crossing 

the strike of the rocks.

Further work would be required to confirm hydrologic interpretations at the sites studied. The best sites 

for further study might be Berkeley Springs, W. Va., 

and Hot Springs, Va., because of the availability of 

much hydrologic information for those areas. The interpretations made from studies of those sites probably 

would be applicable to many other thermal springs in 

the East.

Further work would be required to determine the 

geothermal potential of a particular site. Probably the 

best sites to consider for this type of study would be 

Falling Spring or Bolar Spring, Va. or Hot Springs, N. 

C.; these Virginia springs have the highest energy output of the ones included in this investigation. Although 

the Hot Springs in North Carolina have a visibly small 

discharge, there are reliable reports of hot water discharging beneath the French Broad River near the Hot 

Springs, and in the alluvium near the mouth of Spring 

Creek. Thus, the area probably has a much higher energy output than indicated for the spring in table 3.

This study produced no evidence of very high reservoir temperatures in the aquifers. Oxygen-isotope

ratios of the waters fall within ranges expected for recharge at the latitudes of the springs; no positive shift 

in 8 18 0 is observed. Thus, temperatures at depth in the 

spring systems are almost certainly less than 150°C.

Geothermometers cannot be expected to provide very 

accurate results in systems with low reservoir temperatures. Observed high dissolved-silica levels of some 

cold-water samples (such as those from wells 302, 303, 

and 402, for example) which presumably have never 

been heated will obviously invalidate the estimates of 

temperature from the silica geothermometer and will 

also raise doubts about the indicated temperatures of 

nearby warm-spring waters. Nevertheless, a trend of 

increasing dissolved silica with increasing temperatures is observed in the Appalachian warm springs. 

Spring and hot-well samples at temperatures >25°C 

have indicated chalcedony temperatures within about 

10°C of the observed values, so that equilibration of 

dissolved silica with this mineral is clearly an important reaction associated with heating of the water. 

Cation-geothermometer temperatures of the warmest 

spring waters and of two of the three sampled hot-well 

waters are also in fair agreement with observed temperatures. The concordance of observed chalcedony, 

and cation temperatures for these warmest waters indicates that the chemical processes underlying the success of the cation geothermometer in other systems are 

also operative in the Appalachian thermal springs, and 

further, that reservoir temperatures in these systems 

probably do not exceed the chalcedony values of 30°- 

50°C.

Refinements can be made in geothermometer estimates of hot reservoir temperatures if the hot and cold 

components of a mixed warm water can be distinguished. Chemical analyses of cold-spring and well 

waters in the thermal spring areas proved inadequate 

for estimating the extent of mixing; the composition of 

the hot component could not be deduced from compositions of an assumed cold recharge water and the 

sampled spring water. However, tritium measurements distinguish clearly between modern and 

relatively old waters. Tritium determinations of fractions of cold (young) and hot (old) water, together with 

estimates of the silica content of the cold component, do 

provide improved silica geothermometer estimates of 

reservoir temperatures in some of the mixed spring 

systems.

Potassium and helium seem to be most significant 

among the qualitative geothermal indicators considered. Relatively high potassium levels (which are 

partly responsible for high cation geothermometer 

temperatures) are found in the warmer springs and in 

two of the three hot wells. Dissolved helium is present, 

in amounts from traces to several tenths of a /u,g/L, in
HYDROLOGY AND GEOCHEMISTRY OF THERMAL SPRINGS OF THE APPALACHIANS E35

most of the warm springs issuing from carbonate rocks. 

Only one sample from the silica-rock warm springs 

contained a detectable amount of helium. Arsenic may 

also be a useful geothermal indicator for application in 

the Appalachian warm spring systems if analytical 

sensitivity for arsenic can be increased so that a few 

hundredths of a /ug/L can be determined with confidence.

Several observations regarding the geochemistry of 

the spring waters are apparent from the tabulated 

data. The tritium analyses show that the warm spring 

waters are older than those of nearby cold springs. 

Also, in several of the warm springs the values for 8 13C 

are substantially higher than those in nearby cold 

springs; the high values indicate a longer or slower 

travel path in the formation of the warm water. 

Carbonate saturation indices for the thermal waters 

are higher than those calculated for nearby cold 

springs in the same formation. Three of the silica-rock 

warm springs are considerably undersaturated in calcite; the carbonate thermal waters and Minnehaha 

Springs, on the other hand, are near or above calcite 

saturation.

The technique of using analyzed carbon dioxide concentrations rather than field pH in calculations of carbonate equilibrium data helps considerably to reduce 

discrepancies between different samples from a spring 

of supposed constant composition and also to reduce or 

eliminate apparent carbonate supersaturation of a 

number of spring waters. The proposed method for determination of an adjusted 8 13C, using the gas analysis 

value for CO2 , also leads to more nearly constant compositions for a number of the Appalachian springs. The 

resulting improvements in the geochemical data lend 

support to the idea that outgassing of carbon dioxide 

during the preparation of a 13C sample, or in the course 

of measuring pH, may lead to erroneous conclusions 

about the carbonate chemistry of a water sample. The 

use of analyzed carbon dioxide concentrations may be 

generally helpful in geochemical studies of ground 

waters.

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