GEOCHEMISTRY OF GROUND-WATER IN TWO 

SANDSTONE AQUIFER SYSTEMS IN THE 

NORTHERN GREAT PLAINS IN PARTS OF

ft MONTANA AND WYOMING

i

\

U.S. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1402-C
Geochemistry of Ground-Water in 

Two Sandstone Aquifer Systems 

in the Northern Great Plains in 

Parts of Montana and Wyoming

By THOMAS HENDERSON

REGIONAL AQUIFER-SYSTEM ANALYSIS

U.S. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1402-C

A contribution of the

Regional Aquifer System Analysis Program

UNITED STATES GOVERNMENT PRINTING OFFICE, W AS H I N GTO N : 1985
UNITED STATES DEPARTMENT OF THE INTERIOR 

Donald Paul Hodel, Secretary

GEOLOGICAL SURVEY 

Dallas L. Peck, Director

First printing 1985 

Second printing 1985

Library of Congress Cataloging in Publication Data

Henderson, Thomas, 1957-

Geochemistry of ground-water in two sandstone aquifer systems in the northern Great Plains in parts of 

Montana and Wyoming.

(Regional aquifer-system analysis)(U.S. Geological Survey professional paper; 1402-C) 

Bibliography; C69-C72. 

Supt.ofDocs.no.: 119.16:1402-C

1. Aquifers Great Plains. 2. Water, Underground Great Plains. 3. Water chemistry.

I. Title. II. Series: Regional aquifer-system analysis (Series) III. Series: Geological Survey professional paper; 1402-C

GB1199.3.G74H46 1984 551.49'9078 83-600357

Rev.

For sale by the Branch of Distribution

U.S. Geological Survey

604 South Pickett Street

Alexandria, VA 22304
FOREWORD

The Regional Aquifer-System Analysis (RASA) program was started in 1978 

after a congressional mandate to develop quantitative appraisals of the major 

ground-water systems of the United States. The RASA program represents a 

systematic effort to study a number of the Nation's most important aquifer 

systems, which, in aggregate, underlie much of the country and which represent 

important components of the Nation's total water supply. In general, the boundaries of these studies are identified by the hydrologic extent of each system, and 

accordingly transcend the political subdivisions to which investigations have 

often arbitrarily been limited in the past. The broad objectives for each study are 

to assemble geologic, hydrologic, and geochemical information; to analyze and 

develop an understanding of the system; and to develop predictive capabilities 

that will contribute to the effective management of the system. Computer 

simulation is an important element of the RASA studies, both to develop an 

understanding of the natural, undisturbed hydrologic system, and any changes 

brought about by human activities, as well as to provide a means of predicting 

the regional effects of future pumping or other stresses.

The final interpretive results of the RASA program are presented in a series 

of U.S. Geological Survey Professional Papers that describe the geology, 

hydrology, and geochemistry of each regional aquifer system. Each study within 

the RASA Program is assigned a single Professional Paper number, and where 

the volume of interpretive material warrants, separate topical chapters that consider the principal elements of the investigation may be published. The series of 

RASA interpretive reports begins with Professional Paper 1400 and thereafter 

will continue in numerical sequence as the interpretive products of subsequent 

studies become available.

Dallas Peck 

Director

CONTENTS

Abstract Cl

Introduction 1

Acknowledgments 2

Sample collection and analysis 2

Selection and preparation of wells 2

Field parameters 2

Major and minor elements 3

Dissolved gases 3

Stable and radioactive isotopes 4

Mineralogic analysis 5

Judith Basin study area 5

Physiography 5

Topography 5

Drainage 6

Climate 6

Geology 6

Structure 6

Stratigraphy 8

Mineralogy of the Kootenai Formation 8

Hydrology 10

Data presentation and interpretation 12

Major and minor elements 12

Dissolved gases 23

Isotopic composition 24

Geochemical computer modeling 27

Saturation indices 27

Reaction simulation 31

Carbon isotopic evolution 35

Powder River Basin study area 37

 37

 37

 37

 38

 39

 39

 40

 40

Physiography 

Topography ____ _ ___

Drainage 

Climate 

Geology 

Structure 

Stratigraphy 

Mineralogy of the Lance-Fox Hills aquifer

Powder River Basin study area Continued 

Hydrology 

Data presentation and interpretation 

Major and minor elements - 

Dissolved gases 

Isotopic composition 

Geochemical computer modeling 

Saturation indices 

Reaction simulation 

Carbon isotopic evolution 

Conclusions 

References cited 

Supplemental data 

Geochemical principles 

Solution-mineral equilibria 

Law of mass action 

Equilibrium constants 

Activities and activity coefficients 

Complex equilibria 

Saturation indices 

Phase diagrams 

Aqueous solubility of gases 

Oxidation-reduction 

Mineral dissolution and weathering 

Ion exchange 

Environmental isotopes 

Stable isotopes 

Oxygen and hydrogen - 

Carbon 

Sulfur 

Radioactive isotopes 

Tritium 

Carbon-14 

Geochemical computer models 

WATEQ2 

PHREEQE 

Isotopic evolution equations 

Page

C42 

43 

43 

54

58

59

59

62

65

66

69

73

74

74

74

74

74

75

75

76

76

77

78

79

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83

84

ILLUSTRATIONS

FIGURE 1. Index map of the Judith Basin area, Montana, showing the outcrop of the Kootenai Formation and the locations of wells

sampled C6

2. Mapped faults in the eastern Judith Basin, Montana 7

3. Generalized geologic section of the eastern Judith Basin, Montana, through line A-A' in figures 1 and 6 8

4. Structure contours on Mississippian Madison Group showing the relationship between Big Spring Creek and an underlying synclinal structure, eastern Judith Basin, Montana 9

5. Potentiometric contours and inferred flow paths for the Kootenai Formation, eastern Judith Basin, Montana 11

6. Index map of eastern Judith Basin study area, Montana, showing locations of wells sampled in the Second Cat Creek sandstone of the Kootenai Formation and line A-A' of geologic section in figure 3 14

7. Dissolved solids concentration in waters from the Second Cat Creek sandstone of the Kootenai Formation, eastern Judith

Basin, Montana 15
VI CONTENTS

Page

FIGURE 8. Stiff diagrams of major element concentrations in waters from the Second Cat Creek sandstone of the Kootenai Formation, eastern Judith Basin, Montana C16

9. Trilinear diagram showing divergence in composition between water samples taken north and south of Big Spring Creek

from the Second Cat Creek sandstone of the Kootenai Formation, eastern Judith Basin, Montana 17

10. Log molar ratios of calcium plus magnesium to sodium concentrations in waters from the Second Cat Creek sandstone of

the Kootenai Formation, eastern Judith Basin, Montana 18

11. Sulfate concentrations in waters from the Second Cat Creek sandstone of the Kootenai Formation, eastern Judith Basin,

Montana 19

12. Variations in pH along line A-A' in waters from the Second Cat Creek sandstone of the Kootenai Formation, eastern

Judith Basin, Montana 20

13. Variations in total inorganic carbon concentrations along line A-A' in waters from the Second Cat Creek sandstone of the

Kootenai Formation, eastern Judith Basin, Montana 20

14. Stability diagram for the system Ca-Mg-C02 -H20 at 25°C, with compositions of waters from the Kootenai Formation,

Judith Basin, Montana, superposed 20

15. Stability diagram for the system Mg-K-Si02-Al203 -H20 at 25 °C, with compositions of waters from the Kootenai Formation, Judith Basin, Montana, superposed 21

16. Stability diagram for the system Mg-K-Si02 -Al203-H20 at 25°C, with compositions of waters from the Kootenai Formation, Judith Basin, Montana, superposed 21

17. Variations in strontium concentrations along line A-A' in waters from the Second Cat Creek sandstone of the Kootenai

Formation, eastern Judith Basin, Montana 21

18. Variations in calculated oxidation potential, nitrate plus nitrite, and dissolved organic carbon concentrations along line

A-A' in waters from the Second Cat Creek sandstone of the Kootenai Formation, eastern Judith Basin, Montana 22

19. Variations in iron and manganese concentrations along line A-A' in waters from the Second Cat Creek sandstone of the

Kootenai Formation, eastern Judith Basin, Montana 23

20. Isotopic composition of North American continental precipitation (Gat, 1981) with isotopic compositions of waters from

the Kootenai Formation, Judith Basin, Montana, superposed 25

21. Sulfate dMS values in waters from the Second Cat Creek sandstone of the Kootenai Formation, eastern Judith Basin,

Montana 26

22. Comparison of calculated total inorganic carbon concentrations with modeled concentrations for mineral equilibration 

with no C02 dissolution, following the reduction of 02 and NO", southern flow path of the Second Cat Creek sandstone, 

eastern Judith Basin, Montana 32

23. Comparison of field-measured pH values with modeled pH values for mineral equilibration with C02 dissolution, southern flow path of the Second Cat Creek sandstone, eastern Judith Basin, Montana 33

24. Calculated mass transfer for the calcite-dolomite-cation exchange-C02 dissolution model of ground-water evolution

along the southern flow path of the Second Cat Creek sandstone, eastern Judith Basin, Montana 34

25. Calculated mass transfer for the calcite-dolomite-cation exchange-Morrison Formation leakage model of ground-water

evolution along the northern flow path of the Second Cat Creek sandstone, eastern Judith Basin, Montana 35

26. Comparison of measured and modeled concentrations of selected constituents in waters from the Second Cat Creek

sandstone of the Kootenai Formation, eastern Judith Basin, Montana 36

27. Comparison of measured 6 13C values and the range of modeled 6 13C values for waters from the southern flow path of the

Second Cat Creek sandstone, eastern Judith Basin, Montana 37

28. Structural and physiographic features of the Powder River Basin, Wyoming and Montana 38

29. Index map of the Powder River Basin study area showing major towns and perennial streams, outcrop of the Fox Hills

Sandstone and the locations of wells sampled 39

30. Generalized geologic section of the Powder River Basin study area, Wyoming, through line B-B' in figure 29 40

31. Structural lineaments outuning the Belle Fourche Arch, Powder River B asin, Wyoming and Montana 41

32. Geologic map of the Powder River Basin study area, Wyoming 42

33. Potentiometric contours and inferred flow paths for the Lance-Fox Hills aquifer, Powder River Basin study area,

Wyoming 44

34. Index map of Powder River Basin study area, Wyoming, showing the locations of wells sampled in the Fox Hills Sandstone 48

35. Dissolved solids concentration in waters from the Fox Hills Sandstone, Powder River Basin, Wyoming 49

36. Stiff diagrams of major element concentrations in waters from the Fox Hills Sandstone, Powder River Basin, Wyoming 50

37. Log molar ratios of calcium plus magnesium to sodium concentrations in waters from the Fox Hills Sandstone, Powder

River Basin, Wyoming 51

38. Chloride concentrations in waters from the Fox Hills Sandstone, Powder River Basin, Wyoming 52

39. Sulfate concentrator in waters from the Fox Hills Sandstone, Powder River Basin, Wyoming 53

40. pH values of waters from the Fox Hills Sandstone, Powder River Basin, Wyoming 54

41. Total inorganic carbon concentrations in waters from the Fox Hills Sandstone, Powder River Basin, Wyoming 55

42. Stability diagram for the system Ca-Mg-C02 -H20 at 25°C, with compositions of waters from the Lance-Fox Hills aquifer, Powder River Basin, Wyoming, superposed 56

43. Stability diagram for the system Na-Si02 -Al203-H2O at 25°C, with compositions of waters from the Lance-Fox Hills

aquifer, Powder River Basin, Wyoming, superposed 56
CONTENTS VII

Page

FIGURE 44. Strontium concentrations in waters from the northern flow path of the Fox Hills Sandstone, Powder River Basin,

Wyoming - C56

45. Calculated oxidation potentials and dissolved organic carbon and nitrate concentrations in waters from the northern flow

path of the Fox Hills Sandstone, Powder River Basin, Wyoming - 57

46. Iron and manganese concentrations in waters from the northern flow path of the Fox Hills Sandstone, Powder River

Basin, Wyoming - 57

47. Isotopic composition of North American continental precipitation (Gat, 1981) with isotopic compositions of waters from

the Lance-Fox Hills aquifer, Powder River Basin, Wyoming, superposed 59

48. Comparison of calculated total inorganic carbon concentrations with modeled concentrations for mineral equilibration 

with no CO2 dissolution following the reduction of O2 and NO:, northern flow path of the Fox Hills Sandstone, Powder 

River Basin, Wyoming _____ 64

49. Comparison of field-measured pH values with modeled pH values for mineral dissolution with CO,, dissolution, northern

flow path of the Fox Hills Sandstone, Powder River Basin, Wyoming 64

50. Calculated mass transfer for the calcite-dolomite-cation exchange-CO2 dissolution model of ground-water evolution along

the northern flow path of the Fox Hills Sandstone, Powder River Basin, Wyoming 65

51. Calculated mass transfer for the calcite-dolomite-cation exchange-sulfate reduction model of ground-water evolution

along the northern flow path of the Fox Hills Sandstone, Powder River Basin, Wyoming - 66

52. Comparison of measured and modeled concentrations of selected constituents in waters from the northern flow path of

the Fox Hills Sandstone, Powder River Basin, Wyoming - 67

53. Comparison of measured 613C values and the range of modeled 6 13C values for waters from the northern flow path of the

Fox Hills Sandstone, Powder River Basin, Wyoming 68

TABLES

TABLE 1. Pertinent chemical reactions C3

2. Field analytical methods 4

3. Standard or relative deviations and detection limits of analyses 4

4. Minerals and other phases present in the Kootenai Formation, central Montana 10

5. Flow rates calculated for the Kootenai Formation in the Judith Basin, Montana 11

6. Concentrations of dissolved major elements in waters from the Kootenai Formation, Judith Basin, Montana -- 12

7. Concentrations of dissolved minor and trace elements in waters from the Kootenai Formation, Judith Basin, Montana 12

8. Chemical compositions of ground water from aquifers underlying the Kootenai Formation, Judith Basin area, Montana 20

9. Oxidation potentials calculated for waters from the Kootenai Formation, Judith Basin, Montana - 22

10. Concentrations and partial pressures of dissolved gases in water samples from the Second Cat Creek sandstone of the

Kootenai Formation, eastern Judith Basin, Montana 23

11. Comparison of field-measured pH values and pH values calculated from measured PCO in water samples from the Second

Cat Creek sandstone of the Kootenai Formation, eastern Judith Basin, Montana - - --------------------------- 24

12. Isotopic ratios or activities of isotopes in waters from the Kootenai Formation, Judith Basin, Montana 25

13. Saturation indices of silicate and aluminosilicate aquifer minerals in waters from the Kootenai Formation, Judith Basin,

Montana 28

14. Saturation indices of carbonate and sulfate minerals in waters from the Kootenai Formation, Judith Basin, Montana 28

15. Saturation indices of a 0.7 mole percent manganese calcite solid solution in waters from the Kootenai Formation, Judith

Basin, Montana 29

16. Saturation indices of iron oxyhydroxides and sulfides in waters from the Kootenai Formation, Judith Basin, Montana 30

17. Comparison of radiometric and hydrologic flow rates for ground waters of the Kootenai Formation, eastern Judith Basin,

Montana 37

18. Minerals identified by X-ray diffraction in samples of Fox Hills Sandstone from the Powder River Basin, Wyoming 43

19. Minerals of the Fox Hills Sandstone identified in previous studies, but not detected in X-ray diffraction analyses . 43

20. Concentrations of dissolved major elements in waters from the Lance-Fox Hills aquifer, Powder River Basin, Wyoming 45

21. Concentrations of dissolved minor and trace elements in waters from the Lance-Fox Hills aquifer, Powder River Basin,

Wyoming 46

22. Oxidation potentials calculated for waters from the Lance-Fox Hills Aquifer, Powder River Basin, Wyoming 56

23. Concentrations and partial pressures of dissolved gases in water samples from the Lance-Fox Hills aquifer, Powder River

Basin, Wyoming - 58

24. Comparison of field measured pH values and pH calculated from measured Pco for water samples from the Lance-Fox

Hills aquifer, Powder River Basin, Wyoming " 58

25. Isotopic ratios or activities of isotopes in waters from the Lance-Fox Hills aquifer, Powder River Basin, Wyoming 58

26. Saturation indices of silicate and aluminosilicate aquifer minerals in waters from the Lance Formation and the Fox Hills

Sandstone, Powder River Basin, Wyoming 60
VIII CONTENTS

TABLE 27. Saturation indices of carbonate, sulfate and phosphate minerals in waters from the -Lance Formation and the Fox Hills

Sandstone, Powder River Basin, Wyoming C60

28. Saturation indices of a 3.0-mole percent manganese calcite solid solution in waters from the Lance-Fox Hills aquifer,

Powder River Basin, Wyoming 61

29. Saturation indices of a 0.2-mole percent strontium calcite solid solution in waters from the Lance-Fox Hills aquifer,

Powder River Basin, Wyoming 62

30. Saturation indices of iron oxyhydroxides and sulfides in waters from the Lance-Fox Hills aquifer, Powder River Basin,

Wyoming 63

31. Comparison of radiometric and hydrologic flow rates for Lance-Fox Hills aquifer of the Powder River Basin, Wyoming 68

32. Cation exchange capacities of several clay mineral groups and iron and manganese oxyhydroxides 81

SYMBOLS AND DEFINITIONS

cKj isotopic fractionation factor between the ith and jth

species.

(3 reaction progress composite parameter. 

(3~ beta particle (electron). 

F ratio of carbon input to output rate. 

7; activity coefficient of the ith species. 

AC°R x heat capacity of reaction at constant pressure.

standard free energy of formation of the ith species.

standard free energy of reaction.

standard enthalpy of formation of the ith species.

standard enthalpy of reaction.

isotopic composition expressed as per mil differences in 

the measured isotopic ratios of sample and standard.

dissolving carbonate <5 13C.

solution <5 13C at the end of stage 1 evolution.

solution 6 13C at the end of stage 2 evolution.

additive fractionation factor between solution and 

precipite.

effective porosity.

standard deviation.

measured 14C activity, in percent modern carbon.

14C activity to be expected in the absence of radioactive

decay, in percent modern carbon. 

3j thermodynamic activity of the ith species. 

a; ion size parameter, 

(am) amorphous, 

atm pressure, in atmospheres, 

b saturated thickness of aquifer. 

(c) crystalline. 

CEC cation exchange capacity. 

CT total inorganic carbon. 

D Deuterium, 2H. 

DO dissolved oxygen. 

DOC dissolved organic carbon. 

dhldl hydraulic gradient. 

E° standard electrode potential. 

E h oxidation potential, log activity of e1 . 

EX ion exchange substrate. 

e1 electron, 

(g) gaseous.

AH°

<5 13C, 

513C, 

5 13C2 

eps

ND

I 

IAP

k,

Kgas

Ksp

%)

mi

mL

n 

N 

n 

PDB-1

PI 

P 

pH

PHz

R 

R;

r

SI

SMOW

SpC

(s) 

T 

T

zpc

t 

TU

t,,2 

v

X

Z.

ionic strength.

ion activity product.

denotes ith species.

thermodynamic equilibrium constant.

thermodynamic constant for the exchange of species i 

and j on a substrate.

Henry's Law Constant for the dissolution of a gaseous 

species.

thermodynamic solubility constant.

distribution constant.

molal concentration of the ith species.

molality of the ith element in the solution phase.

molality of the ith element in the solid phase.

neutron.

fractional abundance.

number of observations.

Peedee belemnite 13C isotopic standard.

partial pressure of the ith gaseous species.

negative log base 10 function.

negative log activity of H+ .

pH of the zero point of charge.

gas constant, 8.3147 Joules mole ~ 1 "Kelvin" 1 .

equilibrium isotopic ratio in the ith phase.

correlation coefficient.

saturation index.

Standard Mean Ocean Water 18O and 2H isotopic standard.

specific conductance.

solid.

transmissivity.

temperature.

reference temperature, 25 c

temperature of reaction.

time, in years.

tritium unit.

14C half life, in years.

ground water flow rate.

mole fraction.

charge on the ith species.

Celsius or 298.15°Kelvin.
REGIONAL AQUIFER-SYSTEM ANALYSIS

GEOCHEMISTRY OF GROUND-WATER IN TWO

SANDSTONE AQUIFER SYSTEMS IN THE 

NORTHERN GREAT PLAINS IN PARTS OF MONTANA AND WYOMING

By THOMAS HENDERSON

ABSTRACT

The Kootenai Formation in the Judith Basin, Montana, and the 

Lance Formation and Fox Hills Sandstone in the Powder River Basin, 

Wyoming, constitute two important sandstone aquifer systems in the 

Northern Great Plains region. Ground waters in each of these 

systems evolve from low dissolved-solids concentration, near-neutral 

pH, predominantly calcium and magnesium bicarbonate types in 

their recharge areas, to high dissolved-solids concentration, high pH, 

predominantly sodium-bicarbonate types in the basins. Oxidation 

potentials decrease as the waters flow downgradient under confined 

conditions. Calculation of the saturation states of aquifer minerals 

suggests several groups of mineral phases that could control groundwater chemistry. Mass transfer modeling indicates, however, that the 

observed behavior of major and minor dissolved species in both 

systems can satisfactorily be explained only by equilibration with calcite, dolomite, or calcite and dolomite. The geochemistry of these 

systems is probably controlled by the incongruent dissolution of dolomite to form calcite. This reaction appears to be driven by cation exchange and the dissolution of carbon dioxide. Plausible carbon dioxide sources include organic carbon oxidation and lignite coalification. 

Aluminosilicates influence major element chemistry primarily as subtrates for cation exchange, which, in combination with carbonate 

equilibria, buffer ground water pH at values of 8.5 to 8.9. Dissolvediron concentrations are controlled by equilibration with amorphous 

ferric oxyhydroxides in oxidizing waters, with amorphous ferric 

oxyhydroxides and siderite in moderately reducing waters, and with 

siderite and amorphous ferrous sulfide in strongly reducing waters. 

Measured variations in dissolved carbonate isotopic composition 

compare favorably with carbon isotopic evolution, calculated by 

assuming dedolomitization.

Recharge areas of the two systems are characterized by ground 

waters with high tritium and carbon-14 activities and relatively low 

dissolved-solids concentrations, with calcium and magnesium as the 

predominant cations. Recharge temperatures, calculated from 

dissolved-argon concentrations and 5 18O and 6D isotopic measurements, indicate that recharge is derived primarily from spring snowmelt rather than late spring and summer storms. Ground-water flow 

directions are generally parallel to trends of increasing dissolved 

solids concentrations and decreasing divalent to monovalent cation 

concentration ratios. However, these trends are sometimes obscured 

in areas of leakage or mixing. Further indication of leakage between 

aquifers is provided by abrupt changes in major element and isotopic 

chemistry, which are not characteristic of normally observed geochemical evolution. Ground-water flow rates, calculated by adjusting 

measured carbon-14 activities for carbonate mass transfer, are comparable to values calculated from aquifer tests and potentiometric 

data These carbon-14 flow rates average 1.6 meters per year for the 

Second Cat Creek sandstone of the Kootenai Formation, and 1.3 

meters per year for the Lance-Fox Hills aquifer.

INTRODUCTION

The accurate simulation of aqueous geochemistry 

through the use of computer models can serve as an invaluable aid in characterizing water quality and availability and in determining mineralogic controls on 

natural water chemistry. If the controlling influences 

for a particular aquifer system can be determined, the 

effects of man-made pertubations to that aquifer 

system can be evaluated, and plans to minimize deleterious effects can be considered.

This investigation of two confined sandstone aquifer 

systems of the Northern Great Plains region was initiated in order to: (1) Characterize the geochemical 

controls on ground-water chemistry; and (2) use groundwater chemistry as a hydrologic tool to locate areas of 

recharge, discharge, and leakage from adjacent 

aquifers, and also to determine ground-water ages and 

rates of flow.

Areas of the Judith Basin, Montana, and the Powder 

River Basin, Wyoming, were chosen for detailed study 

because of anticipated population growth in these regions and the need for greater supplies of water in the 

near future. The Kootenai Formation and Fox Hills 

Sandstone-Lance Formation were selected as study 

aquifers on the basis of aquifer geology and hydrology, 

as well as the availability and distribution of suitable 

existing water wells.

Available computer models have been used to 

calculate the status of gas-solution-mineral equilibria involving major- and trace-element species in ground 

water, and to simulate observed chemical and isotopic 

variations in the two flow systems, as a means of determining the influence of mineralogic controls on groundwater chemistry. The computer program WATEQ2 

(Ball and others, 1979; 1980) was used to calculate aqueous speciation of major and trace elements from detailed analyses of water samples collected during this 

study from existing domestic, stock, municipal, and industrial wells. WATEQ2 also computes the degree of 

saturation of a ground water with respect to minerals

Ci
C2 REGIONAL AQUIFER-SYSTEM ANALYSIS

present or likely to be present within the aquifer. In addition, a major emphasis was also placed on activity diagrams, which may serve to distinguish between stable 

and unstable phases that otherwise appear to be in equilibrium with a given ground water.

Simulation of expected changes in the solution chemistry, assuming water quality is controlled by reactions 

with different mineral phases, was accomplished by the 

use of the computer program PHREEQE (Parkhurst 

and others, 1980). Calculation of the resulting isotopic 

compositions associated with computed mass transfer, 

particularly in the carbon and sulfur systems, places an 

additional constraint on a proposed model, in that both 

chemical and isotopic compositions must be correctly 

predicted for the model to be considered valid. A U.S. 

Geological Survey computer program that incorporates 

the carbon isotopic evolution equations of Wigley and 

others (1978) was used to calculate the isotopic evolution associated with proposed mass transfer models, as 

well as to adjust apparent carbon-14 ages for carbonate 

dilution and precipitation effects. Thus, through the use 

of geochemical models, many pertinent chemical reactions can be simulated, and their relative effects on 

solution chemistry can be evaluated and compared to 

observed ground-water chemistries.

Previous investigations by Plummer (1977) and Thorstenson and others (1979) have used geochemical mass 

transfer and isotopic evolution models as a means of 

characterizing chemical reactions responsible for measured changes in dissolved constituents along hydrologic 

flow paths. Ground-water geochemistry of the limestone Floridan aquifer was determined by Plummer 

(1977) to be controlled by equilibration with calcite, 

dolomite, and gypsum, in addition to the oxidation of 

organic carbon by sulfate reduction. In the Fox Hillsbasal Hell Creek aquifer of North Dakota, Thorstenson 

and others (1979) concluded that equilibration with 

calcite and dolomite, oxidation of organic carbon by 

sulfate reduction, and cation exchange by clays control 

ground-water chemistry. Computed mass transfer between aqueous and solid phases, in combination with 

measured carbon isotopic data, was then used to calculate ground-water ages and flow rates.

The present investigation uses many of the techniques developed during these previous studies, applying them to two sandstone aquifers in the Judith and 

Powder River Basins of the Northern Great Plains 

region. The geochemical data collected are also used for 

information regarding the hydrologic characteristics of 

the two flow systems.

A list of chemical reactions pertinent to the discussion of ground-water geochemistry is presented in table 

1. In addition, the reader is directed to the supplemental 

data section for an analysis of the chemical theory used 

in this report.

The reference surface to which relief features and altitude data of the conterminous United States and 

Alaska are related is the National Geodetic Vertical 

Datum of 1929 (NGVD of 1929), formerly called mean 

sea level. The NGVD of 1929 is referred to as sea level in 

this report.

ACKNOWLEDGMENTS

The author would like to thank John F. Busby, 

Donald Langmuir, and Donald C. Thorstenson, for their 

many helpful suggestions and criticisms. H. W. Dodge, 

Jr., of the U.S. Geological Survey Uranium and 

Thorium Resources Branch was instrumental in obtaining core samples of the Fox Hills Sandstone and Donald 

W. Fisher of the U.S. Geological Survey Water Resources Division was responsible for the determination 

of dissolved gases. In addition, Jerome B. Krause of the 

Colorado School of Mines Research Institute was extremely helpful in allowing the use of his mineralogy 

laboratory facilities and X-ray diffraction data files. 

This work was supported in part by U.S. Geological 

Survey contract 14-08-0001-17999.

SAMPLE COLLECTION AND ANALYSIS 

SELECTION AND PREPARATION OF WELLS

Knowledge of the depth of a well and the position of 

screened and cased intervals is critical in determining 

which aquifer is being sampled. Therefore, wells to be 

sampled were chosen on the basis of available drillers' 

logs or other reliable information regarding well construction. This knowledge is especially important in the 

study of interbedded systems, where numerous sandstone aquifers are isolated by shale aquitards.

Prior to sampling, the selected wells were pumped or 

allowed to flow for periods of roughly 30 minutes to 1 

hour, until temperature, specific conductance, and pH 

measurements taken at approximately 10 minute intervals, had stabilized. This initial pumping period serves 

to remove waters influenced by contact with the metal 

well casing from the well bore and adjacent aquifer and 

allows for the collection of samples that are representative of waters present in the aquifer.

FIELD PARAMETERS

Several parameters, including temperature, specific 

conductance, pH, alkalinity, and dissolved oxygen were 

measured in the field to insure accurate determination.
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C3

TABLE 1 . Pertinent chemical reactions

Oxidation-reduction

02(g) + CH20 = C02(g) + H20

4NO ~ + 5CH2O + 4H + = 2N2(g) + 5CO2(?) + 7H9O

NO" + 2CH2O + 2H + = NH + + 2CO2(g) + H.O

4F<0H)s(am , + CH2° + 8H + - 4Fe + ' + C°2< g) + ! 1H2°

2MnO9 + CH2O + 4H + = 2Mn + 2 + CO } + 3H2O

S04~ - + 2CH2O + H + = HS + 2CO V) + 2H,O

CH.jCOOH = CH4(?) + C02(g)

CO,^ + 4H2(?) = CH^ + 2H2O

FeS J, + 7/20" (g) + H20 = Fe + 2 + 2SO; 2 + 2H +

(pyritc)

FeS + 14Fe + '! + 8H2O= 15Fe +2 + 2SO; 2 + 16H +

(pynte)

CH4< g, + 202, g)= H2CO" + H20

Eh redox couples

O,(g) + 4H + + 4e - = 2H2O ; E° = 1 .229 volts

F<0H)3(am, = Fe + 3 + 30H - ; pKFe(OH ,.!( im) = 37~^

FeCOS(c) = Fe + 2 + CO - 2 ; PKFeCOj =To.22

Fe +s + e' =Fe + 2 ; E(l = 0.771 volts

SO4~ 2 + 10H + + 8e ~ = H2S° + 4H2O ; E° = 0.301 volts

Dissolved carbonate equilibria

C02(i;) + H20 = H2CO;>

H2CO'^ = HCOS +H +

HCO,[ = CO - ~ + H +

Mineral dissolution and weathering

CaSO4 -2H,,O(0 = Ca + 2 + SO4' 2 + 2H,,O

CaCO.,(( , + H + = Ca + 2 + HCO -

(tdltltf)

CaMg(CO,,)Q(t) + 2H + = Ca + 2 + Mg f 2 + 2HCO.-

(doloin.tr)

CaM^COj)^) + 2H + = CaCO^(() + Mg + 2 + HCO~

(dolomite) (calcite)

(illiU-)

1.15Al,JSi2Or,(OH)4(c) + 0.6K + + 0.25Mg + ' + 1.2H4SiO4'

(kddmite)

3Na( , ,,A1, ,,Si, 670 K1(OH)2(c) + H + + ll.SH.O =

(Na-montmonllonile)

3.5Al.,Si,O5(OH)4(() + Na + + 4H4SiO(4

(kaolmitc)

3KAl2Si,,O]()(OH),(0 + 3H2O + 2H + = 3Al2Si9O5(OH)4(O + 3K +

(muscovite) (kdolinite)

NaAlSisO8(0 + 6.1H2O+ 1.1H + =

(dlblU.-)

,..J[Al(OH) ;]o7l S,Oj]n;(., ; ,, ) + Na^2.6H4SiO';

Reaction 

number

(1)

(2)

(3)

(4)

(5)

(6)

(7)

(8)

(9)

(10)

(11)

Reaction

number

(12)

(13)

(14)

(15)

(16)

Reaction

number

(17)

(18)

(19)

Reaction

number

(20)

(21)

(22)

(23)

(24)

(25)

(26)

(27)

TABLE 1. Pertinent chemical reactions Continued

Ion exchange

Ca + 2 + Na,7EX = 2Na + + CaEX* 

Mg + 2 + Na0EX = 2Na + 2 + MgEX*

Nuclear transmutation

14N + n = 12C + 3H

Radioactive decay

Reaction 

number

(28) 

(29)

Reaction 

number

(30) 

(31)

Reaction 

number

14C = 14N + )3- (32)

EX-Ion exchange site

The methods and equipment used to evaluate these 

parameters are listed in table 2.

MAJOR AND MINOR ELEMENTS

Major and minor element samples were filtered 

through 124-mm (millimeter) diameter, 0.45-/mi (micrometer) membrane filters using a GeoFilter variable 

speed peristaltic pump and backflush filter plate. Trace 

element samples were filtered through 124-mm diameter, 0.10-/mi membrane filters in order to decrease the 

amount of suspended matter in the sample (Kennedy 

and others, 1974; Jones and others, 1974). These 

samples for major, minor, and trace element analyses 

were then preserved and stored for later analysis by the 

U.S. Geological Survey National Water Quality Laboratory in Arvada, Colorado, in accordance with standard 

U.S. Geological Survey procedures outlined by Brown 

and others (1970) and Skougstad and others (1978). 

Precisions for these analyses at the 95 percent confidence level are 1.96 standard deviations (a) or the detection limit (table 3), whichever is greater.

DISSOLVED GASES

Dissolved-gas samples were collected at selected 

wells, using evacuated glass flasks of the type described 

by Hobba and others (1977). The samples were collected 

at the well discharge point by connecting the sample 

flask to the well head with plastic tubing and allowing 

several volumes of water to flow through the sample 

container to remove adsorbed gases. Stopcocks on the 

container were then closed to the flowing water and
C4 REGIONAL AQUIFER-SYSTEM ANALYSIS

TABLE 2. Field analytical methods (modified from Wood, 1976)

[All measurements made at the well site in non-flowing water, unless otherwise noted]

Parameter Analytical method

pH Orion Research Model 399A lonanalyzer with an 

Orion Model 91-09 combination pH electrode. 

Apparatus calibrated with NBS primary standard buffers (pH 4.01, 6.86, 7.41, and 9.18 at

25°C, corrected for temperature effects) using 

two buffers with pH's bracketing measured pH.

Temperature Mercury thermometer, calibrated from - 1° to 

(T], in degrees 51 °C in 0.1 °C increments. Measured in flowing 

Celsius ( L'C ) water at well head.

Specific conductance Yellow Springs Instruments Model 33 salinity- 

(SpC), in micro- conductivity-temperature meter. 

Siemens per centimeter (/xS/cm at 

25 °C)

Dissolved oxygen Yellow Springs Instruments Model 57 oxygen 

(DO), in milligrams meter. Calibrated in air at temperature of sample 

per liter (mg/L) water and in aliquot of sample water with 5 mL 

Na,,SO^ saturated solution. Measurement made 

with probe submerged in flowing sample water at 

well discharge point.

Alkalinity Potentiometric titration with 0.01639yV H^SO.. 

(bicarbonate, Sample filtered through 0.45 -micro meter memcarbonate), in milli- brane filter. Endpoints determined with inflecgrams per liter tion points in first derivative of titration curve. 

(mg/L)

1>\ the I' S C.-olnijKal SU.ACV

opened to the evacuated sidearm. Argon, helium, nitrogen, oxygen, carbon dioxide, and methane were later 

determined, using a gas-chromatographic procedure.

STABLE AND RADIOACTIVE ISOTOPES

Samples were also collected for the determination of 

the stable isotopic ratios of carbon, oxygen, hydrogen, 

and sulfur, as well as for the radioactive isotopes, 

carbon-14 and tritium. The following procedures, outlined by J. F. Busby and others (U.S. Geological Survey, 

written commun., 1981), were used in collecting these 

isotope samples: 

1. Carbon-13 (<5 13C): The reagent used to precipitate 

dissolved carbonate species from the sample was 

prepared in the following manner: 454 g of strontium 

chloride hexahydrate (SrCl2 -6H2O) were dissolved in 2 L 

(liters) of reagent-grade ammonium hydroxide. Dissolved carbon dioxide precipitates and settles out of 

solution as SrCO3 , allowing the supernatant to be decanted and stored for later use at the sampling site.

TABLE 3. Standard or relative deviations and detection limits of analyses performed by 

the U. S. Geological Survey National Water Quality Laboratory

[Data from Skougstad and others, 1978; fig/L, micrograms per liter; mg/L, milligrams per liter]

Standard 

deviation1 Relative Detection 

Parameter (X = parameter deviation limit Units 

concentration) (percent)

Aluminum 0.073X + 14.97 10 /ig/L

Barium 0.069X + 69.1 100 /xg/L

Boron 0.079X+17.1 20 /ig/L

Cadmium 0.192X + 0.09 1.0 /ig/L

Calcium 0.057X + 0.34 .1 mg/L

Chloride 0.057X + 0.25 16 .1 mg/L

Chromium 0.173X + 2.44 10 /ig/L

Copper 23 1.0 /ig/L

Fluoride 0.109X+ 0.01 .1 mg/L 

Iodide 88 .01 mg/L

Iron 0.056X + 23.90 10 /xg/L

Lead 0.097X + 1.45 1.0 /ig/L

Lithium 0.048X + 4.84 10 /xg/L

Magnesium 0.043X + 0.13 .1 mg/L

Manganese 0.049X+ 9.59 10 /ig/L 

Molybdenum 0.072X+ 0.45 10 /ig/L

Nitrate 22 0.1 mg/L

Nitrate + Nitrite 0.080X+ 0.01 .01 mg/L

Phosphate 12 .01 mg/L 

Potassium 0.092X + 0.11 .1 mg/L

Selenium 0.275X- 0.13 1.0 /ig/L 

Silica 0.031X+ 0.60 .1 mg/L

Sodium 0.039X+ 0.45 .1 mg/L 

Strontium 0.104X+15.4 10 /ig/L

Sulfate 0.001X+ 2.92 5 mg/L

Sulfide .1 mg/L

Vanadium 0.069X+ 0.422 1.0 /ig/L

Zinc 0.034X+16.6 10 /ig/L

Standard deviations calculated Irom above expressions and parameter concentration, in units 

specified

Samples were collected in 1-L polyethylene sample 

bottles; these bottles were rinsed twice with raw sample 

water, then filled to the brim. Approximately 50 mL 

(milliliters) of the sample was poured out and replaced 

with the strontium chloride reagent, so that the bottle 

was again filled to the brim. Caps were tightened twice 

and sealed with electrical tape at the sampling site. 

Later the same day, the caps of the bottles were retightened and dipped in melted paraffin to prevent 

evaporation and exchange of CO2 with the atmosphere. 

2. Oxygen-18 (<5 L8O), deuterium (6D), and tritium (3H): 

Raw-water samples were collected in 120-mL glass bottles for oxygen- 18 and deuterium determinations, and in 

1000-mL glass bottles for tritium. The bottles were 

rinsed twice with raw sample water, filled to the brim,
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C5

and capped, taking care to exclude any air bubbles. The 

caps were tightened twice and sealed with electrical 

tape. Later, the caps were retightened and dipped in 

melted paraffin to avoid any evaporation or atmospheric exchange.

3. Sulfur-34 (634S__2 , 634So_ 2): Samples for determinabO4 S

tion of sulfur isotopic ratios in dissolved sulfate and 

sulfide were collected in the Judith Basin, but a lack of 

appropriate equipment prevented similar sampling in 

the Powder River Basin. Where samples were collected, 

a 4- to 6-L raw-water sample was obtained and dissolved 

sulfide was precipitated by the addition of zinc acetate 

and sodium hydroxide. The suspension was filtered 

under nitrogen pressure through 0.10-/<im membrane 

filters, and both filtrate and precipitate were saved for 

later analysis of sulfate and sulfide isotopic ratios. Unfortunately, insufficient sulfide was present in Judith 

Basin waters for the determination of isotopic ratios.

4. Carbon-14 ( 14C): Two reagents were prepared for 

the collection and preservation of carbon-14 samples: 

(1) Carbon dioxide-free sodium-hydroxide solution was 

prepared by dissolving 150 g of sodium hydroxide 

(NaOH) in 600-mL distilled water. Fifteen milliliters of 

strontium-chloride solution were added to precipitate 

any dissolved carbonate. The precipitate was allowed to 

settle and the supernatant was decanted and stored in 

500-mL airtight bottles; (2) carbon dioxide-free 

strontium-chloride solution was prepared by dissolving 

454 g of strontium chloride hexahydrate (SrCl2 -6H2O) 

in 1-L water. One milliliter of sodium hydroxide solution was added to precipitate any dissolved carbonate. 

After settling, the supernatant was decanted and stored 

in airtight 500-mL bottles.

The preparation of a carbon-14 ( 14C) sample consists of 

precipitating and collecting all dissolved inorganic carbonate species present in a 100-L raw-water sample. 

This was done by carefully filling a 100-L capacity 

separatory funnel-like device with raw-sample water 

delivered through a garden hose, taking care to avoid 

any splashing that might result in exchange with atmospheric CO2. During filling, the top of the funnel was 

capped by a plastic cover to minimize exchange of CO2 

with the atmosphere. When the funnel was nearly full, 

500 mL of the CO2-free NaOH solution was added, 

followed by the addition of 500-mL of the CO2-free SrCl2 

solution. Several grams of ferrous ammonium sulfate 

[Fe(NH4)2(SO4 )2 -6H2O] were added as a flocculant and 

the hose removed when the funnel was filled to capacity. 

A 2-L wide-mouth plastic bottle was rinsed twice and 

filled with raw-sample water, then screwed onto a closed 

flap valve on the bottom of the funnel. The flap valve 

was opened at this point, allowing the precipitate to settle into the jar. Settling of the precipitate was generally 

complete within one or two hours, although occasionally

the first bottle had to be removed and replaced with a 

second bottle to collect all the precipitate. The sample 

bottles were removed, capped, tightened several times, 

and sealed with electrical tape to prevent atmospheric 

contamination.

MINERALOGIC ANALYSIS

Drill core and outcrop samples of the Fox Hills Sandstone were disaggregated by gently pulverizing in a porcelain mortar and pestle, and sulfide nodules present 

were removed by hand from the bulk sample. Fractions 

for clay mineral identification were isolated from the 

bulk sample by a Stokes Law gravity settling method 

using water, yielding particles of less than 2 pm in 

diameter. Bromoform liquid (specific gravity of 2.8) was 

used for the separation of heavy minerals.

Clay separates were mounted directly on glass slides, 

while whole rock, heavy mineral, and sulfide fractions 

were first pulverized in an agate mortar, sieved to 

remove particles greater than silt size, and then 

mounted in aluminum sample holders or on glass slides, 

depending on sample size. These fractions were subsequently examined by X-ray diffraction on a Phillips 

recording diffractometer employing nickel-filtered copper radiation. Interpretations of diffraction patterns 

were based on the file card data of the American Society 

for Testing and Materials (Joint Committee on Powder 

Diffraction Standards, 1981). Using these procedures, 

well crystalized minerals occurring in abundances of 

greater than about five percent were identifiable.

JUDITH BASIN STUDY AREA 

PHYSIOGRAPHY

The Judith Basin of central Montana is a topographic 

depression within an unglaciated section at the western 

margin of the Northern Great Plains. Locations of major physiographic features in the Judith Basin are presented in figure 1.

TOPOGRAPHY

Rolling, grass-covered plains at elevations of 1,050 to 

1,350 m comprise the greater part of the basin, and consist of gently sloping, gravel-covered terraces, dissected 

by steep-sided stream valleys. Gently dipping sedimentary strata form a belt of foothills separating the 

basin from the main flanks of the surrounding mountains. The Kootenai Formation outcrops in this belt,
C6

110° 30'

REGIONAL AQUIFER-SYSTEM ANALYSIS

110°15' 110°00' 109°45' 109° 30' 109° 15'

Highwood Mountains

" * > . 

J3»

47 C 15' iijNorth Moccasin 

Mountains

47° 00'

EXPLANATION

m- KOOTENAI FORMATION 

%%* (UPPER CRETACEOUS)

 MAJOR TOWNS

 J7 LOCATIONS OF KOOTENAI WELLS 

USED FOR THIS STUDY

v. TRACE OF SECTION SHOWN 

7\' IN FIGURE 3

MONTANA 

D

LOCATION OF 

JUDITH BASIN 

STUDY AREA

Big Snowy Mountains

15 KILOMETERS 

I

Geology from Feltis 1973, Zimmerman. 1966.

FIGURE 1. Index map of the Judith Basin area, Montana, showing the outcrop of the Kootenai Formation and the locations of wells sampled.

with a resistant basal sandstone forming the outermost 

of two persistent, pine-covered strike ridges that encircle the uplifts. The wooded Highwood, Little Belt, 

Big Snowy, Judith, and North and South Moccasin 

Mountains rise above the plains to elevations of 1,800 to 

2,750 m, effectively surrounding the basin on all sides 

but the north (Vine, 1956).

DRAINAGE

Runoff in the study area is characterized by high flow 

periods in the late winter and spring in response to rains 

and snowmelt from the nearby mountains. The Judith 

River and Arrow Creek, tributaries of the Missouri 

River, drain the Judith Basin. The main perennial 

streams in the eastern Judith Basin, Big Spring and 

Warm Spring Creeks, exhibit relatively constant flows; 

Big Spring Creek averaged 5.55 m3/s (cubic meters per 

second) and Spring Creek averaged 4.15 m3/s, during the 

1968 water year (Feltis, 1973). Other tributaries contribute considerable seasonal flow, but are commonly dry 

for extended periods.

CLIMATE

The climate of the Judith Basin is a modified continental or semiarid type. More than half of the seasonal 

precipitation falls during the months of May, June, and 

July, with annual averages of 40.8 cm (centimeters) at 

Lewistown and 39.2 cm at Hobson. Relative humidity is 

low except in the mountains, which also receive greater 

precipitation. The average annual temperature at 

Lewiston is 6°C (degrees Celsius), although extremes 

above 40 °C and below 40 °C are not uncommon (Vine, 

1956).

GEOLOGY

STRUCTURE

The structural Judith Basin is a large, northwardplunging syncline bordered by anticlinal folds, which 

form the Big Snowy, Little Belt, Judith, and Moccasin 

Mountains. The Highwood Mountains, consisting of a
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS

110°00' 109°45' 109°30'

C7

109°15'

47°15'

47°00'

46°45'

GEOLOGY BY:

1. CALVERT (1909)

2. REEVES (1931)

3. BLIXT (1933)

4 VINE AND HAIL (1950)

5. WALLACE (1953)

6. VINE (1956)

7. GARDINER (1959)

8. MILLER (1959)

9. FELTIS (1973)

EXPLANATION

KOOTENAI FORMATION 

(UPPER CRETACEOUS)

SECOND CAT CREEK WELLS

FIGURE 2. Mapped faults in the eastern Judith Basin, Montana.

highly dissected mass of Tertiary breccia, tuff, lava 

flows, and related intrusives, border the syncline to the 

northwest. Emplacement of these rocks did not greatly 

influence the regional synclinal structure; however, 

small faults and domes create local structural modifications (Zimmerrnan, 1966).

Major structural features within the study area are 

the Big Snowy, Judith, and North and South Moccasin 

Mountains. The Big Snowy Uplift is an east-westtrending asymmetric anticline, flanked by steeply dipping sedimentary strata to the south and gently inclined beds to the north. Dips in Cretaceous strata on 

the north flank range from 8 degrees to 10 degrees, 

becoming horizontal within 10 to 13 km (kilometers) of 

the mountains. The anticline plunges steeply to the 

northwest and southeast, disappearing a short distance 

outside the study area (Reeves, 1931).

The Judith and North and South Moccasin Uplifts 

consist of a complex series of domes, tilted and faulted 

sedimentary formations, and masses of intrusive rocks. 

Figure 2 shows the locations of major faults in the area 

Cretaceous strata outcrop along the flanks of the Judith
C8 REGIONAL AQUIFER-SYSTEM ANALYSIS

1600

1400-

w 1200-

1000-

800-

600-

400

VERTICAL EXAGGERATION 20X

COLORADO SHALE

UNDIFFERENTIATED JURASSIC 

AND PALEOZOIC SEDIMENTS

KOOTENAI FORMATION .15 KILOMETERS

MORRISON FORMATION National Geodetic Vertical Datum of 1929

FIGURE 3. Generalized geologic section of the eastern Judith Basin, Montana, through line A-A' in 

figures 1 and 6. Faults in the vicinity of the south Moccasin Mountains have been omitted.

and Moccasin Mountains, generally dipping basinward 

at angles of 20 degrees to 60 degrees, but are overturned in some areas. A geologic section through the 

study area demonstrates the relationships between Cretaceous strata and pertinent uplifts (fig. 3).

Structure contours and outcrop patterns in the Lewistown area indicate the presence of a northwest-trending 

synclinal structure on the north flank of the Big Snowy 

Uplift, perpendicular to the trend of the anticline (Gardner, 1959). This syncline parallels, and generally underlies, Big Spring Creek in the Lewistown area (fig. 4). Unfortunately, a lack of data near the South Moccasin 

Mountains prevents following this trend north and west 

of Lewistown; however, it seems probable that the position of Big Spring Creek is controlled by this structure.

STRATIGRAPHY

The sequence of sedimentary rocks exposed in the Big 

Snowy Uplift consists of approximately 3,350 m of 

strata, ranging in age from Precambrian to Holocene. 

The Upper Jurassic Morrison Formation was deposited 

in a fluvial or lacustrine environment, and is comprised 

of 75 to 130 m of variegated siltstones and mudstones 

interbedded with lenticular yellowish-gray to brown 

sandstones and minor freshwater limestones (Suttner, 

1969). At or near the top of the Morrison is the Great 

Falls coal bed discussed by Calvert (1909), which was 

locally exploited at one time. Calvert (1909) reported 

that the Great Falls coal beds, unlike younger Montana 

coals, contain up to 5 or 6 percent pyrite. Overlying an 

erosional unconformity at the top of the Morrison is the 

Kootenai Formation, a Lower Cretaceous fluvial and 

lacustrine sequence of crossbedded sandstones, siltstones, shales, and freshwater limestones, totaling 110 

to 170 m in thickness (Suttner, 1969). At the base of the

Kootenai is a 6- to 30-m-thick bed of crossbedded, gray 

or orange to white sandstone. Another relatively thick 

sequence of red to gray lenticular sandstones occur near 

the middle of the Kootenai. These two persistent sands 

are commonly referred to as the Third and Second Cat 

Creek sandstones. The remainder of the Kootenai is 

composed of red, purple, and gray shales, siltstones, and 

thinner interbedded sandstones and limestones. Individual sandstone beds vary greatly in thickness, and 

may thin and disappear over a short distance (Gardner, 

1959), making determination of Second and Third Cat 

Creek sandstones difficult in some instances.

Conformably overlying the red beds of the upper part 

of the Kootenai Formation are the dark gray to black 

marine sediments of the Colorado Shale. These shales 

are Early to Late Cretaceous in age and crop out at or 

near the surface over much of the basin. Where the entire section is present, the thickness ranges from 450 to 

600 m (Feltis, 1973). Several sandstone beds are present 

among the shales, the most important of which is the 

First Cat Creek sandstone of the Colorado Shale, which 

lies directly above the Kootenai Formation. This brown 

to black sand ranges from 6 to 25 m in thickness, and is 

interbedded with thin layers of fissile black shale (Gardner, 1959).

MINERALOGY OF THE KOOTENAI FORMATION

Drill cores of the Kootenai Formation from central 

Montana were unavailable at the time of this study, so 

no analyses of aquifer or confining bed mineralogy were 

performed. Instead, stratigraphic descriptions and mineralogic studies in the Judith Basin and southwest 

Montana were used to determine the mineral phases 

present within the aquifer system. A listing of these 

minerals is given in table 4.
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C9

110°00' 109°45' 109°30' 109°15'

47°15'

47°00'

46°45'

EXPLANATION

KOOTENAI FORMATION 

(UPPER CRETACEOUS)

SECOND CAT CREEK WELLS

2000 STRUCTURE CONTOUR 

Shows altitude of the top of 

the Madison Group. From 

Gardiner, 1959, Contour 

interval 500 feet. National 

Geodetic Vertical Datum of 

1929.

SYNCLINE Dashed where 

inferred.

10 15 KILOMETERS

Geology from Feltis. 1973

FIGURE 4. Structure contours on Mississippian Madison Group showing the relationship between Big Spring Creek and 

an underlying synclinal structure, eastern Judith Basin, Montana [1 foot (in contour interval)=0.305 meter].

A stratigraphic description by Miller (1959) of the 

Kootenai Formation in the South Moccasin Mountains 

indicates that the sandstones are composed of lightcolored detrital quartz and dark-brown to black chert 

with minor siderite fragments. Many of the sands are 

described as calcareous or limonitic, suggesting the 

presence of calcite and ferric oxyhydroxides, such as 

amorphous ferric hydroxide, goethite, and hematite. In 

addition, lacustrine dolomites and dark-gray to black 

carbonaceous shales were reported as minor occurrences 

interbedded with the sands and shales. Vine (1956)

described a similar mineralogy in a section of Kootenai 

Formation exposed at Skull Butte Dome near Stanford, 

but also mentions sandstones and conglomerates containing limestone pebbles.

Ballard (1966) has studied the petrography of basal 

Kootenai sandstones found at the north flank of the Little Belt Mountains. He classified the Third Cat Creek 

sandstone in the lower part of the Kootenai as moderate 

to coarse-grained, moderate to well-sorted, orthoquartzites or quartz arenites. The terrigenous material 

is almost entirely quartz and chert with quartz cement.
CIO REGIONAL AQUIFER-SYSTEM ANALYSIS

TABLE 4. Minerals and other phases present in the Kootenm Formation, central Montana

| From Ballard. 1966, Miller, 1954, Suttncr, 1968, 1969, Vine, 1956]

Mineral Idealized chemical formula

Quartz 

Siderite 

Dolomite - 

Hematite 

jllite _________

Kaolinite 

Zircon 

Tourmaline group 

Garnet group 

Leucoxene 

Organic matter 

SiO9

CaCO3

FeCO3

CaMg(CO.,),,

ZrSi

Secondary Ti oxides

Minor amounts of sandstone fragments and illite, and 

traces of hematite and organic material are also present. 

Heavy mineral separations show the presence of trace 

amounts of tourmaline, zircon, rutile, garnet, and leucoxene (mixed titanium oxides).

In southwest Montana, the mineralogy of the 

Kootenai and Morrison Formations was investigated by 

Suttner (1968; 1969) using petrographic and X-ray techniques. For the basal Kootenai sandstones, Suttner observed petrographies similar to those found by Ballard 

(1966) in the Little Belt Mountains. Calcite cement 

becomes common in the middle sandstones, and trace 

quantities of well-preserved sodic plagioclase and 

potassium feldspar crystals also become evident. The 

dominant clay minerals observed in Kootenai Formation shales are illite and kaolinite. Minor amounts of 

montmorillonite, chlorite, and mixed layer clays are also 

present, but only illite and kaolinite are present in samples taken nearest the Judith Basin.

HYDROLOGY

The surface- and ground-water resources of the Judith 

Basin have been previously evaluated by Zimmerman 

(1966) and Feltis (1973, 1977). J. F. Levings (U.S. Geological Survey, written commun., 1981) was investigated ground-water flow in the Kootenai Formation 

during the present study using hydrologic computer 

models.

The sandstone beds of the Kootenai Formation constitute the most widely used aquifers south of Big 

Spring Creek. The Second and Third Cat Creek sandstones constitute the main aquifers, supplying water of 

good quality and sufficient quantity for domestic and 

stock use. Isolated sandstone beds also provide water to

wells in the area. These bedrock aquifers are recharged 

by precipitation and infiltration through unconsolidated 

stream deposits on the outcrop. In the outcrop area the 

aquifers are unconfined, but the siltstones and shales of 

the Kootenai and Morrison Formations form confining 

layers within a short distance downdip. Wells located 

downdip from the aquifer outcrop commonly flow at the 

surface.

As indicated by the potentiometric contours plotted 

in figure 5 (J. F. Levings, U.S. Geological Survey, written commun., 1981), the inferred flow direction is generally northward from recharge areas on the flanks of 

the Big Snowy and Judith Mountains. Feltis (1973) 

reported that ground-water discharges from seeps in the 

Kootenai outcrop along the channel of Big Spring Creek 

near Hanover, although the magnitude of this discharge 

is uncertain. The perturbation of the potentiometric 

surface along Big Spring Creek may be explained by 

discharge into the creek, drawdown due to extensive 

utilization of the aquifer in the vicinity of Lewistown, or 

both. The elongation of the associated cone of depression suggests a zone of high permeability along the 

stream valley. Such a high permeability zone would be 

expected if fracturing accompanied the deformation 

that formed the synclinal structure underlying the 

valley.

Zimmerman (1966) reported transmissivity values (T) 

for the Kootenai Formation in the western and southern 

Judith Basin, ranging from 2.4X10~5 to 2.7X10"4 m2/s. 

Feltis (1977) measured an intermediate transmissivity 

value of 1.2X10-4 m2/s for a well (J25) located within the 

study area (fig. 5) drawing from the Second and Third 

Cat Creek sandstones and having a saturated thickness 

(b) of 71 m. Assuming no leakage, constant thickness, 

and constant effective porosity, Darcy's Law may be 

used to calculate ground-water flow rates (v) for flow 

paths perpendicular to the potentiometric contours:

where:

0=effective porosity (assume 0.20); 

dhldl=hydraulic gradient, determined from potentiometric contours.

Flow rates calculated in this manner using Feltis's intermediate transmissivity value are listed in table 5, 

and average 2.0 m/yr (meters per year), which is consistent with values given by Freeze and Cherry (1979) for 

confined sandstone aquifers.

According to Feltis (1973; 1977), wells completed in 

the sandstones of the Colorado Shale or the Morrison 

Formation sometimes yield water of suitable quantity
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS

110°00' 109°45' 109°30'

Cll

109°15'

47 1' 15'

47°00'

46°45'

3800 -

4000-

KOOTENAI FORMATION 

(UPPER CRETACEOUS)

SECOND CAT CREEK WELLS 

AQUIFER TEST WELL

 POTENTIOMETRIC CONTOURS 

Shows altitude at which water 

level would have stood in 

tightly cased wells, 1980 

(Levings, J., U.S. Geological 

Survey written communication, 

1981 J.Contour interval 100 

feet. National Geodetic 

Vertical Datum of 1929.

INFERRED FLOW PATHS

5 10 15 KILOMETERS 

_i_______I_______I

Geology from Feltis. 1973

FIGURE 5. Potentiometric contours and inferred flow paths for the Kootenai Formation, eastern Judith Basin, Montana

[1 foot (in contour interval)=0.305 meter].

Sample location

I1 1 J 11

J IJ

J15 

dh 

(meters)

120

14.H

9QO

dl 

(meters)

11,100 

1 Q nnn

OQ QAfl

Flow rate 

(meters per year)

1.7

1 Q

and quality for domestic or stock use. The First Cat 

Creek sandstone of the Colorado Shale often provides 

moderate amounts of water, as do the isolated sandstone, limestone, or coal beds of the Morrison Formation. Water from the Morrison is commonly high in 

dissolved sulfate, presumably due to oxidation of the 

abundant pyrite found in the formation.

Two deeper Mississippian aquifers, the Madison 

Group and the Kibbey Formation of the Big Snowy
C12 REGIONAL AQUIFER-SYSTEM ANALYSIS

( 

Sample

j-' n

J4 

J 6 

lu,) 8

J10 

Jll

.112 

J13 

J14 

j 15

J16 

J17 

J20 1 

J21' - 

J22' -

J23-J -

Calcium, 

Ca

55

84

120

83

200

(JO

39

44

8

7

4.9

51

3.1

2.4

43

.6

4.2 

4 

4.1

101

Magnesium, 

Mg

15

25

29

24

51

31

8.4

15

3.2

2.0

19

.9

t Q

.0 

.0 

0

39

, Sodium .

Na 

1

1 D

22 

34

20 

29

2.7

41

240 

170

62 

190

200 

100

230 

160

153

Potassium, 

K

Va + K

4.1

2.9

2.1

5.1

2.5

2.7

2.9 

2.2

3.0 

1.9

1.7

7.7

1.2 

3.0 

159

0

Chloride, 

Cl

1.7

3.0

2.2

2.2

2.1

2.4

2.3

1.6

7.4

2.9

1.4 

4.9

2.6

O Q

5.9 

.2 

4

12

Sulfate, 1

so4

24 

160

150 

510

210 

9.2

62

320 

45

QQ

64

4.2

72 

76 

51

A.P,

Bicarbonate, 

HCO3

275

241

269

303

248

135

273

291 

396

505

292

420 

.340 

335 

351

177

Carbonate, 

C03 PH

<1 7.59

< 1 7.62

^ 1 7 £Q

^ 1 7 Qt\

<1 6.87

<1 7.70

4 8.49

i) O 1 Q

^ 1 7 Q 1 )

4

O ^^

5 8.57

<1 7.60

21 8.8 

6 8.5 

12

O

n o

0

remperature, 

T

15.0

12.0

«

<>

8

1

8.4

1 ') Q

8

O

8

7

16.8

11.7

9.4

13.4

9.4

10.9

15 

14 

16 

11

Specific 

conductance, 

SpC (jtS/cm 

at 25°C)

454

678

925

715

1,250

745

273

1QQ

1,130

663

594

792

750

762

813 

748 

610 

630

750

Group, exhibit the highest hydrostatic pressures in the 

basin, and appear to discharge as numerous springs in 

the study area. Feltis (1973) concluded that faulting and 

fracturing have contributed to the localization of discharge from these aquifers, resulting in large springs, 

such as Big Spring, which is the source of the creek of 

the same name.

DATA PRESENTATION AND INTERPRETATION

MAJOR AND MINOR ELEMENTS

Analyses of the 21 well waters used to chemically 

characterize the Kootenai aquifer within the Judith 

Basin are presented in tables 6 and 7. The locations of

Aluminum, 

A]

Bromide,

Barium, Br Cadmium, Chrom mm, Copper 

Ba (mg/L) Cd Cr Cu

Dissolved

oxygen, Fluoridc. 

O, F

(mg/L) (mg/L)

Iodide. 

I

(mg/L)

Lead. 

Pb

«l
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C13

these sampled wells are plotted in figure 1. Waters from 

16 wells (J2-J17) were sampled and analyzed as part of 

this study; the remaining five (J20-J24) were obtained 

from previous investigations (Feltis, 1973; 1977). Correlation of ground-water chemistry with available driller's 

logs and well depths revealed significant vertical variations between samples from the Second and Third Cat 

Creek sandstones, despite the many similarities in the 

major element chemistry, mineral saturation indices, 

Eh and pH of these waters. The variability resulting 

from considering the whole formation as a uniform 

aquifer is much greater than if the two sandstones are 

treated as individual systems, suggesting that although 

similar geochemical processes are acting to control 

water chemistries in both aquifers, subtle differences in 

the stratigraphy, mineralogy, and permeability of the 

two systems result in the observed variability in the 

two chemistries. As such, geochemical modeling and detailed discussion of the evolution of ground-water geochemistry with time was restricted to the Second Cat 

Creek sandstone in the eastern Judith Basin, where the 

availability of analyses allows such a treatment. Locations of these twelve wells, determined to be perforated 

in the Second Cat Creek sandstone, are plotted in figure 

6. Analyses derived from the literature do not include 

accurate field pH and alkalinity determinations required for geochemical investigations, and were therefore used primarily as an aid to mapping major element 

variations.

Maps showing the chemical properties of ground 

waters in the Second Cat Creek sandstone reveal systematic variations in the concentrations of major 

elements. The dissolved solids distribution, plotted in 

figure 7, reveals a gradual increase from recharge areas 

in the Big Snowy and Judith Mountains north and 

westward into the basin, as would be expected if mineral

dissolution or leakage of highly mineralized waters was 

occurring along the flow path. The increase in dissolved 

solids is greater to the north of Big Spring Creek than to 

the south, suggesting that different mechanisms may 

control dissolved solids concentrations on either side of 

the creek. In addition, where the aquifer outcrops on the 

flanks of the South Moccasin Mountains, recharge 

would be expected, but low dissolved solids waters of 

the type sampled in the recharge areas to the south and 

east are not evident. The absence of low dissolved solids 

waters suggests that net recharge to the aquifer may be 

negligible near the South Moccasin Mountains.

Separating the dissolved solids data into its major element components, the relative contributions of major 

cations and anions to the general increase can be observed. Area! variations in major element concentrations are plotted in figure 8 in the form of Stiff diagrams 

(Stiff, 1951), which portray cationic and anionic concentrations on three horizontal axes extending to either 

side of a vertical zero axis. The Stiff diagrams show 

gradual, uniform increases in sodium plus potassium, 

chloride, and bicarbonate plus carbonate along the flow 

path. Calcium and magnesium decrease in concentration basinward. Only sulfate displays significantly different behavior on either side of Big Spring Creek, accounting for the high in dissolved solids near the South 

Moccasin Mountains. South of the creek, sulfate concentrations increase gradually but remain at low levels, 

whereas north of the creek, they increase dramatically 

over a short distance. A trilinear plot of major element 

concentrations (fig. 9) clearly shows the divergence of 

compositions between flow paths along the north and 

south sides of Big Spring Creek.

Uniform increases in sodium with concurrent decreases in calcium and magnesium concentrations are 

almost certainly indicative of ion-exchange phenomena.

waters from the Kootenai Formation, Judith Ba.\in, Montana

|K-r liu-r. <indi<.iti-, lonirmratiims Ixlnu dm i nun hum lui ih.it an. <l\

Nitrite plus

nitrate, Organic Phosphate, Sulfide

NO,, + NO carbon, PO4 Silica, (mg/L as 

Lithium. Manganese, Molybdenum, (mg/L, as C (mg/L as Selenium, SiO,, Strontium, hydrogen Uranium, Vanadium, Zinc,

Li Mn Mo nitrogen) (mg/L) phosphorus) Se (mg/L) Sr sulfide) U V Zn

4 < 1 < 1

7

4

'2

1

7

2

4

1 i

12

M

Ml

")(

41) < 1

HO < 1

H(l < 1

11(1 <1

'*<> < 1

< 1 < 1

10 <1

4 < 1

2 < 1(

20 < 1(

4 < 11

<()

<

<

<

<

<

?

<

<

<

<

'f < 10 <

14(1 20 <1() .(

0.8 0

7 <

2 h <

2 'i <

'i "> <

_> <;

7 4

'i ! <

1 1 <

1 (1

<t (

'1 (

1 <

<

<

<

<

<

<

<

<

2 <

1 <

 < 1 (I 'I <(l 111 < 1 <

17 ,101

i 4 .">!)(

ill .(01

it . !0(

!( , iOI

1 7'i(

4 1,U)I

! H 781

( 7 871

7 K 2,101

1 < 1 'Ml 'IK

3 h 02 < 1 71 2")(

.3 <.01 < 1 HI l.KX

1 K) <

2 11 <

<

<

<

<

<

<

<

<

17 <

1 ( <

)<l <

"0

1 1 <

1 i <

2 1 <

 <

i a <

1 1 <

4 14 <

84

1")

7

1 t

7

1 i

7

1

tl

<

t
C14

110° 00'

REGIONAL AQUIFER-SYSTEM ANALYSIS

109°45 P 109"30' 109°15'

47" 15'

47°00'

46°45'

J21

EXPLANATION

KOOTENAI FORMATION 

liiii (UPPER CRETACEOUS)

i J16 SECOND CAT CREEK WELLS

k A'

TRACE OF SECTION SHOWN 

IN FIGURE 3

10 15 KILOMETERS

Geology from Feltis. 1973

FIGURE 6. Index map of eastern Judith Basin study area, Montana, showing locations of wells sampled in the Second Cat 

Creek sandstone of the Kootenai Formation and line A-A' of geologic section in figure 3.

The log ratio of the molalities of exchangeable divalent 

and monovalent cations (log/ ([Ca]+[Mg])/[Na]2 ), an operational parameter plotted in figure 10, demonstrate the 

progressive nature of the exchange reacions (table 1, 

reactions 28 and 29) in the aquifer. The progression from 

positive to negative log ([Ca]+[Mg])/[Na]2 values serves 

to identify areas of recharge and directions of flow in the 

system.

Either illite or kaolinite may serve as the ionexchange medium. Because the cation-exchange capacity is several times greater in illite (10-40 meq/100 g 

(milliequivalents per 100 grams)) than kaolinite 

(3-15 meq/100 g), illite is probably the dominant exchanger, adsorbing a much greater reservoir of 

exchangeable sodium than kaolinite. Ferric oxyhydroxides, which are abundant in the Kootenai, have a pHzpc 

of approximately 6.7 to 8.5, or close to the ground-water 

pH, indicating that the surfaces of these particles are 

nearly neutral, and should have minimal exchange 

capacities.
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C15

110°00' 109°45' 109°30' 109°15'

47" 15'

47°00'

46°45'

500

500'

EXPLANATION

KOOTENAI FORMATION 

(UPPER CRETACEOUS)

SECOND CAT CREEK WELLS

LINE OF EQUAL DISSOLVED 

SOLIDS CONCENTRATION 

Interval 100 milligrams per 

liter. Dashed where inferred.

10 15 KILOMETERS

Geology from Feltis. 1973

FIGURE 7. Dissolved solids concentration in waters from the Second Cat Creek sandstone of the Kootenai Formation,

eastern Judith Basin, Montana.

Subtle increases in chloride concentration are probably due to leaching of confining layers or leakage from 

adjacent aquifers. Chloride minerals, such as halite 

(NaCl) and sylvite (KC1), are commonly associated with 

sedimentary rocks, but the absence of any marine 

evaporite sequences in the Kootenai preclude the use of 

these minerals as a significant chloride source. High 

chloride waters are commonly associated with shales, 

and leaching of shale confining beds could readily supply the required amounts of chloride (Johns and Huang,

1967; Billings and Williams, 1967). Similarly, upward 

leakage of high chloride waters through faults and fractures could transport chloride into the aquifer. Slightly 

larger increases in chloride concentrations north of Big 

Spring Creek may reflect differences in controlling 

mechanisms.

As shown in figure 11, sulfate concentrations also behave differently on either side of Big Spring Creek, suggesting some type of structural controls on groundwater chemistry. The most likely sources of sulfate
C16

11CTOO'

REGIONAL AQUIFER-SYSTEM ANALYSIS

109°45' 109°30' 109°15'

47 1 ' 15'

47°00'

EXPLANATION

KOOTENAI FORMATION 

(UPPER CRETACEOUS)

SECOND CAT CREEK WELLS

STIFF DIAGRAMS 

Na + K 

Ca

Mg , ^M^ , , 

10 5 0 5 10 

milliequivalents per liter

HCO-5+CO., SOA J 6 

Cl 4

46°45'

10 15 KILOMETERS

Geology from Feltis, 1973

FIGURE 8. Stiff diagrams for major element concentrations in waters from the Second Cat Creek sandstone of the

Kootenai Formation, eastern Judith Basin, Montana

within the aquifer would be the dissolution of trace 

amounts of gypsum (reaction 20) or the oxidation of 

pyrite (reactions 9 and 10), which can proceed in either 

the presence or absence of molecular oxygen. Either or 

both of these mechanisms could explain the gradual increases in sulfate observed to the south of Big Spring 

Creek, as both gypsum and pyrite are commonly found 

disseminated in shales, but the highly disparate sulfate 

concentrations on either side of the creek implies the influence of different controls to the north. The occurrence

of numerous faults and fractures in the area of the 

Judith and South Moccasin Mountains (as shown in fig. 

2), suggests the possibility of upward migration of high 

sulfate waters from underlying formations. Analytical 

data for Morrison and Madison Formation waters typical of the Judith Basin, (presented in table 8) indicate 

sulfate concentrations several times higher than those 

measured in waters from the overlying Kootenai Formation. Feltis (1973) commented on the importance of fractures in the Big Snowy, Judith, and North and South
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS

EXPLANATION

 J13 SOUTHERN FLOW PATH 

AJ12 NORTHERN FLOW PATH

DIRECTION OF FLOW

C17

Ca 

CATIONS EQUIVALENT PERCENT

Cl 

AN IONS

FIGURE 9. Trilinear diagram showing divergence in composition between water samples taken north and 

south of Big Spring Creek from the Second Cat Creek sandstone of the Kootenai Formation, eabtern 

Judith Basin, Montana.

Moccasin Mountains region in localizing discharge from 

these lower formations as springs. This type of fracturecontrolled discharge could also serve as hidden recharge 

to the Kootenai Formation, and would explain the 

presence of high sulfate waters in the fractured areas. 

Chloride concentrations are also higher in the underlying aquifers, so leakage would also explain slightly 

higher chloride concentrations north of the creek. 

Waters of the Kootenai Formation to the north and 

south of Big Spring Creek might be isolated from each 

other by the synclinal structure observed in outcrop and 

subsurface. Hinge faults and fractures would allow 

Kootenai Formation ground water to move upward and 

discharge into the stream bed alluvium, creating 

distinct flow systems on either side of the creek.

Significant increases in pH are also evident along the 

flow paths, as shown by the transect along line A-A' in 

figure 12. Two reactive mineral groups, carbonates 

(reactions 21-23) and aluminosilicates (reactions 24-27)

could interact with ground waters of the Kootenai and 

increase the pH. Assuming that the dissolution of these 

minerals is driven by the dissolution of CO2 and the subsequent dissociation of H2CO3 °, the hydrolysis of 

aluminosilicates will result in an increase in pH but no 

increase in CT beyond that contribution supplied by the 

addition of CO2 . However, when driven by the same 

mechanism, both congruent and incongruent dissolution of carbonates will result in an increase in pH accompanied by an increase in CT above and beyond that supplied by the addition of CO2 , thus serving to distinguish 

between these two groups of dissolution reactions.

Because the relative concentrations of H2CO3°, HCO~, 

and CO~2 are dependent on the pH of the solution (reactions 17-19) apparent increases in HCO~ and CO~2 may 

be due to increases in pH accompanied by dissociation of 

H2CO3°, rather than to actual increases in dissolved carbonate species. Calculation of total inorganic carbon (CT) 

clarifies this point, in that CT data reveal whether or not
C18

110° 00'

REGIONAL AQUIFER-SYSTEM ANALYSIS

109°45' 109° 30' 109°15'

47°15'

47°00'

46° 45'

EXPLANATION

KOOTENAI FORMATION 

(UPPER CRETACEOUS)

SECOND CAT CREEK WELLS

LINE OF EQUAL LOG MOLAR 

CALCIUM PLUS MAGNESIUM TO 

SODIUM SQUARED RATIO 

(LOG(Ca + Mg)/Na2 ) Interval 

is 1.0. Dashed where inferred.

10 15 KILOMETERS

Geology from Feltis, 1973

FIGURE 10. Log molar ratios of calcium plus magnesium to sodium concentrations in waters from the Second Cat Creek 

sandstone of the Kootenai Formation, eastern Judith Basin, Montana.

an increase in alkalinity is solely the effect of an increase 

in pH. The aqueous speciation model of WATEQ2 (Ball 

and others, 1980) was utilized to calculate the distribution of carbonate species from major element, pH, 

alkalinity, and temperature data. All of the carbonate 

species were then summed to determine CT. A transect 

along line A-A' in figure 13 demonstrates that CT 

generally increases downflow, although a small decrease 

may be noted north of Big Spring Creek. The observed

increases in CT reflect the presence of carbon sources in 

the aquifer. Carbon may enter the aqueous system 

through the oxidation of organic matter during the 

reduction of oxygen (reaction 1), nitrate (reactions 2 and 

3), and (or) sulfate (reaction 6), or during the fermentation 

of organic matter (reaction 7). Dissolution of carbonate 

minerals (reactions 21-23) and the migration of 

gaseous CO2 (reaction 17) or CH4 (reaction 11) may also 

increase -"r
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS

110°00' 109°45' 109°30'

C19

109°15'

47 1' 15'

47°00'

46°45'

200-

EXPLANATION

KOOTENAI FORMATION 

(UPPER CRETACEOUS)

SECOND CAT CREEK WELLS

LINE OF EQUAL SULFATE 

CONCENTRATION Interval 

50 milligrams per liter. Dashed 

where inferred.

10

Geology from Feltis. 1973

FIGURE 11. Sulfate concentrations in waters from the Second Cat Creek sandstone of the Kootenai Formation, eastern

Judith Basin, Montana.

Acetate fermentation is not a viable mechanism, 

because the Eh of the system is far too high for this 

reaction to proceed. The oxidation of organic carbon 

provides a more feasible carbonate source. Reduction of 

O2 and NO~ supplies significant quantities of CO2 

(>0.54 mmol/L) to oxygenated waters, but the rapid 

consumption of these oxidants and the absence of H2S 

in the Second Cat Creek sandstone suggests that the oxidation of organic carbon by O2 , NO~, or SO~2 does not 

generate significant quantities of CO2 far beyond the

outcrop area. Sulfate reduction or the coalification of 

lignite may produce high CO2 partial pressures (PC02) in 

Morrison or Kootenai strata adjacent to'the Second Cat 

Creek sandstone. The uncharged nature of CO2(aq) and 

H2CO° would allow these species to diffuse through 

microfractures and between clay particles of the shale 

confining beds into aquifers with lower CO2 partial 

pressures, whereas the bulk of any H2S produced by 

sulfate reduction would be precipitated in situ as ferrous sulfides. Well J16 appears to be screened in the
C20 REGIONAL AQUIFER-SYSTEM ANALYSIS

TABLE 8. Chemical compositions of ground water from aquifers underlying the Kootenai Formation, Judith Basin area, Montana

[Units in milligrams per liter unless otherwise noted, °C, degrees Celsius]

Formation

Calcium, 

Ca

Magnesium, 

Mg

Sodium

Na

i, Potassium, 

K

Na + K

Bicarbonate, 

HC03

Sulfate,

so4

Chloride, 

Cl pH

Temperature, 

T

( JC)

Dissolved 

solids Comments

Collected by the U S Geological Survey from wells located about 45 kilometers due east of Lewistown 

"Mean composition of eight samples taken within the Judith Basin Analyses from Zimmerman (1966) and Feltis (1973) 

3 Sitc 9 from Busbv and others (U S Geological Survey, written commun , 1981).

EXPLANATION

 SOUTHERN FLOW PATH 

A NORTHERN FLOW PATH

J13 J17 J12

0 10 20 30 40

DISTANCE ALONG LINE A-A', IN KILOMETERS

FIGURE 12. Variations in pH along line A-A' in waters from the Second Cat Creek sandstone of the Kootenai Formation, eastern 

Judith Basin, Montana.

EXPLANATION

 SOUTHERN FLOW PATH 

A NORTHERN FLOW PATH

0 10 20 30 40 

DISTANCE ALONG LINE A-A', IN KILOMETERS

FIGURE 13. Variations in total inorganic carbon concentrations 

along line A-A' in waters from the Second Cat Creek sandstone of 

the Kootenai Formation, eastern Judith Basin, Montana.

Third Cat Creek sandstone and draws water from a zone 

whereas H2S is evident. Sulfate is nearly depleted, and 

CT and organic carbon are much higher than in wells of 

the Second Cat Creek sandstone, suggesting the active

oxidation of organic matter by sulfate reducing bacteria. Similar CO2 partial pressures in this and surrounding samples, despite large differences in pH, may 

indicate that the aquifers in the formation are at a state 

of equilibrium with respect to dissolved gaseous species, 

and that CO2 might indeed be free to migrate from 

organic-rich to organic-poor zones within the formation.

Migration of CO2 into the aquifer would increase the 

acidity of waters present and allow for dissolution of carbonate minerals, which would further increase CT. The 

stabilities of calcite, dolomite, and magnesite (MgCO3 ) in 

ground waters of the Kootenai Formation may be investigated by calculating the stability fields of these 

minerals in the system Ca-Mg-CO2 -H20. An activity 

diagram constructed using the data of Ball and others 

(1980) is presented in figure 14, with compositions of 

Kootenai Formation ground waters superposed. Calcite 

is clearly the stable phase, indicating that any dolomite 

in contact with these waters should dissolve, possibly in 

an incongruent reaction to form calcite (reaction 23).

Potassium and silica concentrations were nearly constant in waters of the Second Cat Creek sandstone; 

potassium averaged 3.2 mg/L (milligrams per liter) and 

silica averaged 8.6 mg/L, while aluminum was below deMAGNESITE DOLOMITE CALCITE

LOG OF CALCIUM ACTIVITY MINUS LOG OF MAGNESIUM ACTIVITY (LOG (a CaJ /a Mg*2))

FIGURE 14. Stability diagram for the system Ca-Mg-CO2-H2O 

at 25 °C, with compositions of waters from the Kootenai Formation, Judith Basin, Montana, superposed.
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C21

tection (< 15 pg/~L) in all waters sampled. The silica concentration in sample J2 is anomalously low (0.0 mg/L), 

probably due to analytical error, and has been discarded. 

Despite the lack of aluminum data, the stabilities of pertinent aluminosilicates in Kootenai ground waters may 

be investigated using cation ratio activity diagrams. 

With thermodynamic data from Helgeson (1969) and an 

average silica activity of 10~387 , the stabilities of illite, 

kaolinite, montmorillonite, and chloride may be calculated in terms of log [(a +2)/au+)21 and log [(a )/(a )].

L Mg H J » A tt J

Superposition of Kootenai geochemical data on this 

stability diagram (fig. 15) shows that kaolinite is the 

stable phase in all waters sampled. The stability of 

kaolinite reported by Helgeson (1969) was several kilojoules per mole less than the preferred kaolinite stability 

value of Bassett and others (1979). Recalculation of the 

relative stabilities of these aluminosilicates using the 

reevaluated kaolinite stability results in shifting the 

illite-kaolinite stability boundary downward into the 

area of observed compositions, but also creates a stability field for Mg-montmorillonite between that of chlorite 

and kaolinite (fig. 16). Montmorillonite was not observed 

in Kootenai shales obtained near the Judith Basin. An 

increase in illite stability, or further decrease in kaolinite 

stability (which would still fall within the range of 

values presented by Bassett and others, 1979) would 

shift the illite-kaolinite stability boundary downward

LOG OF POTASSIUM ACTIVITY MINUS LOG OF HYDRONIUM ACTIVITY (LOG (a K + /a H4_) )

FIGURE 15. Stability diagram for the system Mg-K-SiO2 -Al2O3 - 

H2O at 25 °C, with compositions of waters from the Kootenai Formation, Judith Basin, Montana, superposed. Thermodynamic data 

from Helgeson (1969). aH SjOo=10~3-87 .

without enlarging the Mg-montmorillonite field. This 

final interpretation is consistent with field evidence, 

although the slight, but highly significant, variability in 

aluminosilicate stabilities due to variable crystallinity 

and composition of the pertinent phases renders any conclusions drawn for a particular system highly tentative, 

and subject to revision as better data for that system 

become available.

Strontium increases, then decreases in concentration 

along the Second Cat Creek sandstone flow path (fig. 

17). Similar behavior is observed for calcium and magnesium, suggesting similar controls for these alkaline

LOG OF POTASSIUM ACTIVITY MINUS LOG OF HYDRONIUM ACTIVITY (LOG (a <+ 3 H+ ) )

FIGURE 16 Stability diagram for the system Mg-K-SiO2 -Al2O3- 

H2O at 25°C, with compositions of waters from the Kootenai Formation, Judith Basin, Montana, superposed. Thermodynamic 

data for kaolinite from Bassett and others (1979). Remaining thermodynamic data from Helgeson (1969). aH SiOo=10~3 '87.

EXPLANATION

 SOUTHERN FLOW PATH 

A NORTHERN! FLOW PATH

£ 0 10 20

DISTANCE ALONG LINE A-A', IN KILOMETERS

FIGURE 17. Variations in strontium concentrations along line A-A' 

in waters from the Second Cat Creek sandstone of the Kootenai 

Formation, eastern Judith Basin, Montana.
C22 REGIONAL AQUIFER-SYSTEM ANALYSIS

earth elements. The importance of ion exchange in controlling calcium and magnesium levels has already been 

discussed, and similar controls would be equally effective in producing the observed variations in strontium 

concentrations (Wahlberg and others, 1965a; 1965b). 

Strontium might originate from the dissolution of 

disseminated strontianite (SrCO3) or celestite (SrSO4), 

which could be present in the aquifer or in confining 

shale beds, although there is little evidence for the existence of these phases. Strontium may coprecipitate 

with calcite (Holland and others, 1964), aragonite 

(Kinsman and Holland, 1969), gypsum or anhydrite 

(Usdowski, 1973), and would thus be released into the 

ground-water system upon dissolution of these phases. 

A positive correlation of greater than 99 percent significance (r=0.716, n=15) was calculated between sulfate 

and strontium concentrations, so it seems probable that 

strontium originates from the dissolution of gypsum, 

which could occur either within the aquifer or in adjacent strata where these phases are present.

Oxidation potential was estimated using the Nernst 

equation and appropriate redox couples (reactions 

12-16); Eh values calculated in this manner are presented in table 9. Eh values in the Second Cat Creek 

sandstone range from 850 mV in the recharge area to 

approximately -100 mV for wells farthest from the 

outcrop (J12, J15) with a minimum of -210 mV (J17) 

near the South Moccasin Mountains (fig. 18). A general 

decrease in oxidation potential is indicated, from oxidizing conditions in the recharge area to relatively uniform, 

moderately reducing conditions a short distance from 

the outcrop and into the basin. The one highly reducing 

sulfidic water sampled near the South Moccasin MounTABLE 9. Oxidation potentials calculated JOT waters from the Kootenai Formation, 

Judith Basin, Montana, using O2~H2O, Fe(OH)3 -FeCO3 or SO~ 2-H2S 

redox couples

[O2 , oxygen, H.,O, water, FefOH) , feme hydroxide; FeCOj, sidente. SO' 2 , sulfate; HjS, 

hydrogen sulfide; mV, millivolts]

Sample

J2

J3

J4

j5 ___________

J6 

TO

jlO __________

J12

T1 9

J 1 -3

J15

J16

J17

Calculated Eh 

(mV)

-230

-214 

-207 

44 

51 

- 1 

848

-101

23 

-77 

-283 

214

Redox couple

SO" 2 H S

O/^i - 2 T_T C

S04- 2 -H2S

Fe(OH)3([ -FeCO3

o2-H2d(am)

Fe(OH)3(^) -FeCO3 

Fe(OH)3(am)-FeCO3

SO*- 2 -H2S

°- s

LU rf i- §

£ 2

EXPLANATION

 SOUTHERN FLOW PATH 

A NORTHERN FLOW PATH

cc LLI

< tl

O -i 6.0-

J13J17 J12

0 10 20 30 40

DISTANCE ALONG LINE A-A1 , IN KILOMETERS

FIGURE 18. Variations in calculated oxidation potential, nitrate plus 

nitrite, and dissolved organic carbon concentrations along line 

A-A' in waters from the Second Cat Creek sandstone of the 

Kootenai Formation, eastern Judith Basin, Montana

tains appears somewhat anomalous. Strongly reducing 

conditions at J17 may be associated with the presence 

of faults related to the South Moccasin Uplift. Fluids 

containing reduced species such as H2S and CH4 would 

be free to migrate upward from lower strata along fracture zones, creating locally reducing conditions where 

they enter other formations. Alternatively, well J17 

may penetrate a zone particularly rich in organic material, which would allow for the biologically mediated 

reduction of sulfate (Goldhaber and Kaplan, 1974). Failure of the remainder of Second Cat Creek sandstone 

waters sampled to attain strongly reducing conditions 

is likely due to a paucity of organic matter suitable as a 

substrate for sulfate reducing bacteria (Dockins and 

others, 1980).

Dissolved organic carbon (DOC) concentrations support this conclusion, decreasing rapidly down the flow 

path and leveling off at an average value of 0.7 ± 

0.4 mg/L (fig. 18). Consumption of DOC during the 

reduction of O2 and NO~ accounts for the initial DOC 

decrease. A lack of solid organic matter in the aquifer 

would prevent further reduction from occurring.
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C23

EXPLANATION

 SOUTHERN FLOW PATH 

A NORTHERN FLOW PATH

J13 J17 J12

10 20 30 40 

DISTANCE ALONG LINE A-A', IN KILOMETERS

FIGURE 19. Variations in iron and manganese concentrations 

along line A-A' in waters from the Second Cat Creek sandstone of the Kootenai Formation, eastern Judith Basin, 

Montana

Iron and manganese determinations indicate initially 

low concentrations of these elements in oxygenated outcrop waters, increasing following the onset of anoxic 

conditions, then gradually decreasing again (fig. 19). 

The presence of ferric oxyhydroxides and siderite suggests control of iron concentrations by these phases in 

oxidizing and moderately reducing waters (Langmuir, 

1969; Nordstrom and others, 1979; Champ and others, 

1979). Where H2S is present, ferrous sulfides might 

similarly control aqueous iron concentrations.

Oxyhydroxide and carbonate solid phases are similarly observed to control manganese concentrations 

with decreasing Eh (Morgan, 1967). Regulation of manganese by mineral phases similar to those that control 

iron would explain the covariation in concentration of 

these two elements as Eh decreases, with oxyhydroxides controlling in oxidizing environments, siderite or

rhodochrosite (MnCO3) in moderately reducing environments, and ferrous sulfides in strongly reducing waters 

(Champ and others, 1979).

DISSOLVED GASES

Interpretation of dissolved gas concentrations was 

hindered by difficulties encountered during sampling. 

Many of the waters sampled were highly effervescent as 

they flowed from the well head, and gas bubbles invariably accumulated on the walls of the sample vessel. 

Although efforts were made to dislodge as many as possible, many samplers contained some attached bubbles, 

leading to analytical concentrations greater than the 

quantity of gas actually dissolved in the sample. Those 

samples displaying evidence of contamination of this 

type, such as anomalously high concentrations of 

N2(>40 mg/L), Ar(>l mg/L), and O2, or field observations of bubbles, were discarded. Because of this difficulty in sampling, the data presented in table 10 are 

considered only semiquantitative. Discussion of these 

analyses will be restricted to the observed concentration 

trends, and to comparisons with the same parameters 

calculated by other methods. The analyses presented in 

table 10 are all from the southern flow path of the Second Cat Creek sandstone and are arranged in order of 

occurrence from outcrop to basin.

Oxygen decreases down the flow path because of decreasing Eh and carbon dioxide decreases down the flow 

path because of increasing pH. Agreement between 

these dissolved oxygen values and those measured in 

the field is relatively good, differing by a maximum of 

only 7 percent; however, a comparison of analyzed Pco 

and that calculated by WATEQ2 from field-measured 

pH and alkalinity shows discrepancies of up to 40 percent. Although part of these differences are possibly due 

to the gas sampling problem, they also reflect errors in 

pH created by exchange with atmospheric CO2 during 

measurement. CO2 outgassing would tend to make the 

measured pH higher than the actual in situ value due to 

the redistribution of HCOj and COj2 required to maintain equilibrium among dissolved carbonate species

TABLE 10. Concentrations and partial pressures of dissolved gases in water samples from the Second Cat Creek sandstone of the Kootenai Formation,

eastern Judith Basin, Montana

[mg/L, milligrams per liter; atm, atmospheres]

Sample

jlO -

J 1J

T1^ _

Oxygen

mg/L

7 9

r>3

atm

0.12 

.009 

nnnfi

Carbon 

dioxide

mg/L

55 

3.4

9 7

atm

0.022 

.0016 

r>ni3

Nitrogen

mg/L

20 

31

31

atm

0.79 

1.33 

1 3Q

Argon

mg/L

0.80 

1.09 

1 13

atm

0.010 

.015 

017

Helium

mg/L

0.000 

.002 

003

atm

0.000 

.001 

.002

Methane

mg/L

0.00 

.02 

.02

atm

0.0000 

.0006 

.0006
C24 REGIONAL AQUIFER-SYSTEM ANALYSIS

(reactions 17-19). Using field-measured alkalinities, bicarbonate activity coefficients calculated using 

WATEQ2, and Kco and Kj [Henry's Law constant for 

CO2 dissolution (reaction 17) and the first dissociation 

constant for H2CO3° (reaction 18), respectively] for the 

appropriate temperature (Jacobson and Langmuir, 

1972), pH values may be calculated from the measured

pH=-log

*HCO.r'

Such values are given in table 11, and are as much as 

0.18 pH units lower than the field-measured pH. These 

observations are inconsistent with those of Pearson and 

others (1978), who concluded that field-measured pH's 

may be as much as 0.2 units higher than pH values in 

the aquifer, again suggesting that difficulties in sampling have influenced these results.

Concentrations of nitrogen (N2 ) indicate initial saturation with atmospheric N2 in the recharge area (PN =0.7 

atm); N9 values increase and then level off at higher concentrations down flow. Bacterially mediated denitrification (reaction 2), which should occur concurrently with 

O., reduction, would explain the decrease in NO~ plus 

NO^ and increase in N2 between J10 and J13, as well as 

a leveling off of N2 concentrations after nearly all of the 

NO~ has been consumed. Mass balance calculations indicate that the 0.24 mmol/L NO~ present in outcrop area 

waters is insufficient to account for the observed increase in N2 , however, suggesting that another 

mechanism may be responsible. Thorstenson and others 

(1979) postulate that high N9 concentrations in the Fox 

Hills-basal Hell Creek aquifer of the Dakotas may be 

the result of upward leakage or diffusion of N9 from gas 

pools at depth, but the presence of N2 gas pools in the 

Judith Basin area is uncertain.

Helium, argon, and methane increase in concentration 

down the flow path. Helium is a product of radioactive 

decay by alpha emission, and probably accumulates in 

the ground water as a result of decay of unstable 

isotopes in the uranium and thorium decay series. 

Methane is thermodynamically unstable in waters with 

Eh's above 100 mV, such as samples J13 and J15, and 

could not have formed in the Second Cat Creek sandstone. Therefore, methane most likely migrated into the 

aquifer system after being formed in more reducing 

strata. According to Mazor (1972) argon should remain 

at concentrations that are representative of atmospheric 

saturation at the temperature of the recharge waters, but 

only if the aquifer system is closed to influx and efflux of 

this gas. The presence of argon in concentrations exceeding its maximum solubility at 0°C (PA =8.12 X10~3

TAHLF. 11. Comparison of field-measured pH values and pH values calculated from 

measured PCQ in water samples from the Second Cat Creek sandstone of the Kootenai 

Formation, eastern Judith Basin, Montana

Sample

1 1(1

J13 

J15 

pH, 

field measured

6.87

8.48

8

CL C

pH, 

calculated

6.69

8.30

Difference

ft 1Q

.18 

.10

atm at 1370 m elevation, 4.3X10"4 cubic centimeters 

Ar/cubic centimeter H2O; Douglas, 1964) by nearly 40 

percent suggests either migration of argon into the 

aquifer or contamination from sampling difficulties. 

Methane generation at depth might displace noble gases 

into higher strata where they would be able to accumulate (D. Fisher, U.S. Geological Survey, written 

commun., 1980); however, the above-mentioned sampling difficulties preclude the drawing of such conclusions.

ISOTOPIC COMPOSITION

The isotopic compositions of ground waters in the 

Kootenai Formation are presented in table 12. As 

shown in figure 20, <5 18O and <5D measurements of 

ground waters from the Kootenai fall slightly below the 

worldwide meteoric water line of Craig (<5D=8<518O+10; 

1961b), but roughly correspond to the composition of 

North American continental precipitation as reported 

by Gat (1980) (<5D=7.95<518O+6.03), indicating the 

presence of only meteoric water in the aquifer. Gat also 

reports the data of Yurtsever (1975), who has compiled 

<5 18O values for continental precipitation and correlated 

them with mean surface temperatures, allowing a determination of recharge temperature to be made. Kootenai 

<5 18O measurements range from 20.1 to 17.6, indicating a precipitation temperature near 0°C, which is 

significantly lower than the mean annual temperature 

of 6°C for the Judith Basin. This temperature discrepancy suggests that present-day recharge is derived 

primarily from spring snowmelt, rather than from late 

spring and early summer rains, and that past recharge 

events must have occurred under similar conditions. 

Although significant variability in both <5 18O and 5D 

was observed, it does not seem readily attributable to 

differences in the aquifers sampled, and so is probably 

due to variations in the elevations of the Little Belt, Big 

Snowy, and Judith Mountain recharge areas.

Sulfate <5 34S measurements were obtained for all except two Judith Basin samples, but low concentrations 

of H2S precluded the measurement of sulfide 534S
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS

TABLE 12. Isotopic ratios or activities of isotopes in waters from the Kootenai Formation, Judith Basin, Montana

|pCi/L, picocurics per liter; < indicates concentrations below detection limit tor that analvsis]

C25

Sample

TO S 2-

11 .p

J4 

JJ

J6

TO J«

j,0 __ __

Jll 

T 1O J 1 * 

T 1 } J 1 -5

J14 

J 1J

J16 

j!7 

Oxygen

(6 180)

1Q Q

1 Q Q

19.1

1 Q Q

19.1

19.0

1Q 9

20.0

1 O A

1 Q O

20.1

i "7 a

 - 19.4

1 Q O

Hydrogen 

(6D)

ITJ

- 148

14C

-147

-141

1 *}Q

-149

ijj

Carbon

(6"C)

16.6

11.6

9.6

11.6

16.0

10.2

7.1

7.9

10.1 

9.4

14.4

9.2

Sulfur, SO; 2 

(6"S)

rtr\ f~

22.0

7.6

6.6

10.4

.4

3.3

11.9

8.6

2

f~

9.0

Tritium 

(pCi/L)

1.0

1.9

8

C

6.9

5.1

<2.9

410

6

r

<2.1

<4.3

s* n Q

6.0

<2.2

Carbon- 14 

(percent modern)

<1.3

1 Q

11.3

34.1

OO Q

1.3

 *^ OQ Q

18.4

^ Q

<.9

10.0

<.9

1.6

<1.4

o

CO

*

O to

DC

O

Q_ 

O

O

CO

CD 

O 

cc 

Q

D

LU 

Q

-130

-140

-150

-160

-21 -20 -19 -18 -17

OXYGEN-18/OXYGEN-16 ISOTOPIC RATIO (5 180. SMOW)

values. Sulfate S34S values exhibit considerable variations across Big Spring Creek, despite the close similarity of values for each group on either side of the 

creek, as shown in figure 21. A two-tailed F-test demonstrates that these two groups were derived from populations with a common variance at the 95 percent 

significance level, enabling a pooled standard deviation 

to be calculated. This calculation results in mean S34S 

values at the 95 percent significance level of 2.1 ±3.2 for 

the southern population, and 9.8±3.2 for the northern 

population. The difference in the means of the two populations is much greater than that predicted by a t-test 

for random variance at the 95 percent significance level, 

and demonstrates that the null hypothesis of no significant difference in the 634S values across Big Spring 

Creek must be rejected. Because S34S is dependent on 

the source of sulfate (Holser and Kaplan, 1966), it 

follows that sulfate is derived from significantly different sources on either side of the creek. Bacterial 

sulfate reduction, which would cause such increases in 

634S, might occur in only one well, J17, where H2S was 

present, and so would not account for higher 634S values 

in all three wells near the South Moccasin Mountains. 

Differences in both the 634S values and the sulfate concentrations on either side of the creek support the 

hypothesis of structurally controlled leakage into the 

Second Cat Creek sandstone occurring in the fractured 

region north of Big Spring Creek, but not in the more 

competent strata to the south.

Waters of the Second Cat Creek sandstone exhibit 

6 13C values that increase asymptotically from 16.0 at

FIGURE 20. Isotopic composition of North American continental precipitation (Gat. 1980) with isotopic compositions of waters from the 

Kootenai Formation, Judith Basin, Montana, superposed.
C26

1 icr oo'

REGIONAL AQUIFER-SYSTEM ANALYSIS

109°45' 109"30' 109°15'

47 1 ' 15'

47°00'

46°45'

EXPLANATION

KOOTENAI FORMATION 

(UPPER CRETACEOUS)

 2.6 SECOND CAT CREEK WELLS- - 

Sulfur 34/sulfur- 32 isotopic 

ratio (834S).

10 15 KILOMETERS

Geology from Feltis. 1973

FIGURE 21. Sulfate 634S values in waters from the Second Cat Creek sandstone of the Kootenai Formation, eastern Judith

Basin, Montana.

the beginning of closed (to a CO2 gas reservoir) system 

conditions (J10), to values ranging from 7.1 to 9.4 

in the basin. If CH4 was migrating into the aquifer and 

being oxidized by O, to form CO2 (reaction 11), 5 13C 

values would be expected to decrease basinward from 

the isotopically light nature of CH4 (Claypool and 

Kaplan, 1974). The oxidation of methane as a CO9 

source can therefore be considered negligible.

Wigley and others (1978) conclude from modeling 

studies that <5 13C should increase asymptotically in

systems where aqueous carbonate species are fractionated and removed from the system, as in the incongruent dissolution of dolomite (reaction 23). This 513C 

evolution model corresponds with 513C trends in the 

Second Cat Creek sandstone. As the limiting <5 13C solution value is dependent on input and output 513C values, 

it can be calculated from reaction simulation mass 

transfers and the isotopic evolution model of Wigley 

and others (1978) to substantiate the postulated controlling reactions.
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C27

Kootenai Formation tritium concentrations are below 

10 pCi/L for all samples except J10, which has a value of 

410 pCi/L. Low tritium concentrations reflect a lack of 

recent recharge to the confined system, and support the 

hypothesis that recharge to the Second Cat Creek sandstone is insignificant in the South Moccasin Mountains. 

The high tritium value for J10 indicates that recharge 

occurs primarily in the Big Snowy Mountains (and 

probably also in the Judith Mountains, where no 

samples were collected for this study), which is in agreement with conclusions drawn from the dissolved solids 

distribution and the divalent-monovalent cation ratio 

data.

GEOCHEMICAL COMPUTER MODELING

SATURATION INDICES

Saturation indices for Kootenai Formation aquifer 

minerals, calculated using the critically evaluated data 

and aqueous speciation model of WATEQ2 (Ball and 

others, 1980), demonstrate that, although the samples 

originate from at least two distinct aquifers, the 

equilibrium relationships between aqueous and mineral 

phases in each aquifer are sufficiently similar to allow a 

unified discussion of thermodynamic equilibrium in 

ground waters of the Kootenai Formation. WATEQ2 

saturation indices for Judith Basin samples are presented in table 13 for the silicates and aluminosilicates, 

and in table 14 for the carbonates and sulfates.

Quartz, the major aquifer constituent, appears to 

have little control on silica concentrations in solution. 

Highly ordered a-quartz is oversaturated, a condition 

often observed for crystalline phases that in lowtemperature systems do not precipitate directly from 

solution, but rather form initially as an amorphous or 

cryptocrystalline polymorph, such as amorphous silica 

or chalcedony. In most samples, these less crystalline 

phases are undersaturated, indicating that silica is not 

in equilibrium with the solution. An exception to this 

generalization is sample J10, in which chalcedony is at 

saturation. This well is situated on the Kootenai outcrop where most recharge occurs; judging from the 

negative Si's for many aquifer minerals, it is probably 

perforated in sandstones that have been leached of their 

more reactive carbonate phases, leaving only quartz, 

ferric oxyhydroxides, and perhaps kaolinite, illite, and 

montmorillonite to react with infiltrating ground water. 

As a result, silica concentrations have become high 

enough to enable chalcedony to reach saturation and 

possibly precipitate as a grain cement.

Because aluminum was present at concentrations 

below the detection limit of 0.015 mg/L, this value was 

entered into WATEQ2 as the aluminum concentration

to determine the maximum possible saturation indices 

of aluminosilicate minerals. Using this upper limit for 

aluminum concentration, many aluminosilicates are 

undersaturated, indicating that they do not control 

aqueous aluminum concentrations.

Albite, a detrital aluminosilicate, is undersaturated in 

all ground waters of the Kootenai Formation, while 

allophane, a poorly crystallized, variable-composition 

aluminosilicate phase, is saturated in all samples. 

Allophane is commonly encountered as an alteration 

product on the surfaces of weathered feldspar grains. 

The presence of this type of alteration is indicative of 

the final stage of feldspar weathering, which has been 

characterized in the experiments of Busenberg and 

Clemency (1976) by the very slow release of cations and 

silica. Illite and kaolinite are both saturated, as is montmorillonite. Although montmorillonite was not observed in Kootenai shales near the Judith Basin, its 

presence as an alteration phase cannot be ruled out. It 

seems probable that one or more of these saturated 

aluminosilicates controls aqueous aluminum and silica 

concentrations.

Saturation indices for other detrital silicates reported 

as trace constituents in Kootenai sandstones were not 

considered because of the high temperatures required 

for the formation of minerals such as garnet and 

tourmaline.

The carbonate minerals, calcite, dolomite, and siderite, which are abundantly represented in the aquifer, are 

at or near saturation. They appear to play an important 

role in controlling solution chemistry in most waters 

sampled except in the ground water from J10, where 

leaching of the outcrop has apparently eliminated these 

phases. Calcite is slightly oversaturated in many 

samples, indicating that determinate errors in field pH 

measurements, or calcite solid solutions involving 

manganese, strontium, and iron may be involved. These 

elements in solid solution would tend to decrease the 

stability of the solid phase slightly, and also decrease 

the saturation index of solid solution. For unknown 

reasons, dolomite saturation indices are rather widely 

scattered, but the mean dolomite SI shows that the 

mineral is roughly at saturation in the aquifer. Siderite 

is near saturation in all but highly oxidizing or strongly 

reducing samples (J2-4, 10, 16, 17), suggesting that 

some other iron phase may regulate iron concentrations 

in these.

Although rhodochrosite (MnCO3 ) is the thermodynamically stable Mn(II) phase in moderately reducing 

ground waters (Morgan, 1967), the mineral is undersaturated in many samples. The near equality of all but 

two of the saturation indices to a value of 1.3, 

however, is suggestive of a controlling phase of similar 

composition. Saturation indices for J2 and J10, which
C28 REGIONAL AQUIFER-SYSTEM ANALYSIS

TABLE 13. Saturation indices of silicate and aluminosilicate aquifer minerals in waters from the Kootenai Formation, Judith Basin, Montana

Sample Silicates Aluminosilicates

Quartz Chalcedony Amorphous

silica

Albite Allophane Kaolinite Illite Calcium 

montmorillonite

j 2 _______

J3

J^ 14

J TTJ 

Ifi

TO

JIO 

Jll

J12 -

J13 

J14 -

J15 

i it; J 1D

T 1 7

 0.42 ±0.08

.30 ± .10

.32 ± .10 

.44± .08

.37 ± .08

.50+ .08

.36± .10 

.29+ .08

.37+ .08

.37 ± .09

.36± .08 

.32+ .10

.36+ .09

-0.12±0.08

-.25± .10 

-,'24± .10

.11+ .08

.17+ .08

.05+ .08

.19+ .10

.22+ .08

.16+ .08

.18+ .09

.17+ .08

-.22+ .10 

.18+ .09

0.95 + 0.08

-1.09± .10

1.08+ .10

- .95± .08

-1.00± .08

- .90+ .08

-1.03± .10

-1.05± .08

-1.00± .08

-1.02± .09

-1.01± .08

-1.07± .10

1 . U-., + . Uj

-2. 57 ±0.89

-3.26+ .91

-3.34+ .91

-3.11+ .89

T\ J L, X ^'*-'

1.43+ .89

-1.41± .89

1.34+ .89

-1.55± .90

2.39± .89

-0.27 + 0.70

.28+ .67

 . £, 1 ~T_ UC*

-,26± .67 

-.29+ .69

-.36+ .59

-.26+ .35

.74+ .77

-.37+ .77

-.25+ .70

-.51± .78

-.29+ .78 

-.28+ .67

1.32 + 1.71

.17+ .72

.74+1.72 

1.01 ±1.71

1.29±1.71

1.05±1.71

.87±1.71

,48± .71

.77 ± .71

.95 ± .71

.65± .71

.52 ± .72 

1.14+ .71

1.36± 1.98

.01 ±2.00 

.04+1.99

.48± 1.98

1.24± 1.98

.31 + 1.98

.30± 1.99

.67 ±1.98

07 , OQ

.61 ± .98 

.71 ± .98

or , QQ

0.61 ±2.01

-.39 ±2. 02

,16±2.01

.51 ±2.01

.04 + 2.00

- .15 ±2. 01

- .42 ±2. 01

-.00 ±2. 01

.02 ± 2.01

-.23 ±2. 01

-.52 ±2. 02 

i Q , o n i

Mean and 

standard 

deviation 0.37 + 0.06 -0.17 ±0.06 - 1.01+0.06 -2. 53 ±0.93 -0.34 + 0.14 0.88 ±0.27 0.55 ±0.48

TABLE 14. Saturation indices of carbonate and sulfate minerals in waters from the Kootenai Formation,

Sample

.12 

j 3 _________

J* 14

IT

J6 _________

IQ

Jio - 

1 1 1 J 1 '

1 19 J lz 

1 1^ ,) 1J

J14 ________

J15 ________

J16 ________

.117 

Mean and 

standard 

deviation

Judith Basin,

Carbonates

Calcite

n i f, -L n

( >Q _i_

. /y ± 

'*n _i_

4.1 _i_

 .44+ .

1 t\f, _i_

 06 ± .

 .07± .

f)Q .

 -.01+ .

.08+ . 

.04+ .

0.19+ .

06 

06 

06 

06 

06 

06 

06 

06 

08 

12 

06 

14 

18 

06

18'

Dolomite

0.08 + 0.08 

.41 ± .08 

.03 ± .08 

.13± .08 

.29+ .06 

.52 ± .80

-.28± .08 

.07+ .12 

-.14+ .16 

.21+ .08 

-.42+ .24 

-.77± .39 

-.36+ .08

-0.03 ±0.36'

Siderite

-0.22±0 

.31± 

-.06+ . 

-.16± 

-.23 + 

-.07 + 

- 12. 13 ±2 

-.08± 

-.12±

- .31 ± 

.33±

-.16± 

-.66 ±2

-.72±

-0.16±0

.10 

.07 

07 

.09 

.08 

.09 

.08 

.09 

.39 

.73 

.07 

.73 

.08 

.19

.30'

Rhodochrosite

-2. 80 ±8. 20 

-1.13± .25 

- 1.11 ± .14 

-.94± .14 

-1.07+ .12 

-1.28+ .31

Q QQ i Q OA

1.81+ .86 

- 1.38 ±2. 08 

- 1.54 + 4.12 

1.36± .45 

- 1.11 ±2.08

1.21+2.76

1 . \J\J X ~J

- 1.30±0.25 U

Strontianite

4.09 +

.y.) ±

1.06 +

Q 1 ,

.90 +

2.31 +

.93 + 

- .32 ± 

- 1.12±

- J-> _L

- J i _L

-.55± 

1.11 +

6.95 

.10 

.10 

.10 

.10 

.10 

.10 

.10 

.12

4.12 

.10 

.14 

.14 

.10

-0.79±0.40'- 2

Montana

- 0.05 ± 0.35

Sulfates

Gypsum

2. 00 ±0.12 

1.10+ .06 

.75+ .06

1.12± .06 

- .41 + .06 

- .95 ± .06 

2.46± .27 

1.68± .08 

1.83± .08 

2.77± .12 

1.50± .06 

2.84+ .14 

4.11 ± .61 

1.36± .06

1.78 ±0.99

Celestite

'-* 4ft 4. fi

1.68+ .

- 1.33± . 

- .98 ± . 

.86+ .

1.52+ .

2.94+ .

1.90+ .

- 1.51 + . 

-2.22± .

- 1.61 ± .

-2.59+ . 

3.80+ .

1.65+ .

95 

10 

10 

10 

10 

10 

30 

10 

10 

12 

10 

14 

61 

10

- 1.89±0.82J

are strongly reducing and oxidizing waters, respectively, are exceptions. It has been determined that 

calcite is apparently oversaturated with respect to these 

ground waters, so it is postulated that rhodochrosite occurs as a solid solution in calcite. Bodine and others

(1965) reported an experimentally derived distribution 

constant (kD) of 17.4 for Mn+2 in calcite. The distribution constant relates the molalities of Mn and Ca in 

solution to mole fractions of these elements in the solid 

phase:
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C29

mMn

m;Ca m

where x and L denote solid and solution phases, respectively. From this relationship and average manganese 

(0.60 fxmol/L) and calcium (1.5 mmol/L) concentrations, it 

can be determined that calcite in equilibrium with these 

waters should contain an average of 0.7 mole percent 

rhodochrosite. The free energy of a calcite containing 0.7 

mole percent rhodochrosite was calculated to be 

 1128.1 kilojoules per mole by assuming ideal mixing 

between the two components (Berner, 1975), for the 

reaction:

Ca11 _x)MnxCO3 =(l-x)Ca+2+xMn+2+CO3- 2 ,

AG.0

ideal mix=xRTln x+(l-x)RTln(l-x),

and

f Mn calcite ""^^ideal mix ' ^^f calcite '

x(AG° +2 -AG° 2 -RTlnkD),

fMn

where:

X=mole fraction Mn in solid solution

R=gas constant

T= temperature (degrees kelvin).

Saturation indices computed from this free energy and 

the enthalpy of pure calcite are presented in table 15.

TABLE 15. Saturation indices of a 0.7 mole percent manganese calcite solid solution 

in waters from the Kootenai Formation, Judith Basin, Montana

[Free energy for the phase Mn0 ^-pa.^ ^^CO calculated by assuming an ideal solid solution 

(Berner, 1975)]

Sample

SI calcite 

(CaCO3)

SI Mn calcite

(Mn0.007Ca0.993CO3)

J3 - 

J4 -

J5 - 

J6 _______

J8 _______

Jll 

J12 

J13 

J14 

J15 

J16 

J17 

Mean and

standard

deviation

0.41

.29

.30

.36

.44

.06

.15

.07

.28

-.01

-.08

-.04

0.34

.21

.22

.33

.37

-.02

.09

-.01

.20

-.08

-.16

-.12

0.19 ±0.18 0.11+0.19

They demonstrate that apparent oversaturation with respect to pure calcite is largely explained when considering the manganese-calcite solid solution, and they 

suggest that an impure calcite may also control manganese concentrations in Kootenai ground waters.

Saturation indices for strontianite vary considerably, 

but are invariably less than zero. Although it is possible 

that this mineral exists as a solid solution with calcite, 

the inconsistency of strontianite saturation indices gives 

little support to the control of strontium concentrations 

by such a phase in these waters.

Despite the fact that no sulfate minerals were reported 

in the Kootenai Formation, their ubiquitous nature 

demands a discussion of sulfate saturation indices. In 

addition, the use of dissolved sulfate as a conservative 

tracer in evaluating flow directions and leakage in 

ground-water systems requires that any sources or sinks 

for this constituent be considered. Gypsum, the most 

common of the sulfate minerals, might occur in trace 

quantities dissimulated throughout the formation, but 

negative saturation indices indicate that an insufficient 

amount of this mineral is present to allow solid and 

aqueous phases to reach equilibrium. Celestite (SrSO4) is 

similarly undersaturated, and is probably not present in 

the formation above trace amounts.

In the oxidizing environment, manganese concentrations are probably controlled by the stability of manganese oxyhydroxides. Potter and Rossman (1979) have 

investigated naturally occurring manganese oxyhydroxides in dendrites, stream deposits, and desert varnishes, and have identified metastable phases, such as 

romanchite [(Ba, H2O)2Mn5O10], hollandite (BaMn8O16), 

cryptomelene (KMn8O 16 ), coronadite (PbMn8O 16 ), 

todorokite [(Mn, Ca, Mg)Mn3O7 H2O], and birnessite 

[(Na, Ca, K)Mn7O14 -3H2O]. Unfortunately, thermodynamic stabilities for these phases are unavailable except 

for birnessite, so equilibrium calculations are impossible 

in most cases. Using the detection limit of 1 /ig/L for 

the concentration of manganese in sample J10, saturation indices indicate oversaturation with respect to both 

birnessite, with a value of 4.29, and manganite 

(7-MnOOH), with a value of 1.05. Oversaturation with 

the stable and metastable manganese oxyhydroxide 

phases in oxidizing environments (Crerar and Barnes, 

1974), using the manganese detection limit as the 

manganese concentration, suggests that either phase 

may control manganese solubility in oxidizing ground 

waters of the Kootenai Formation.

Ferric oxyhydroxides are particularly common in the 

Kootenai, giving rise to the red or purple color seen in exposures of the formation. The saturation indices presented in table 16 show that hematite (Fe2O3) is grossly 

oversaturated in all samples. Goethite (FeOOH) is oversaturated in oxidizing and moderately reducing waters,
C30 REGIONAL AQUIFER-SYSTEM ANALYSIS

Sample

Hematite

Oxyhydroxides

Goethite 3(am)

Sulfides

Pyrite FeS

Sulfide absent

.15 

I Q

I 1 1

.1 ' '

J12

14

1 1 ~t

Mean and standard 

deviation

12.80 + 0.34 

12.68+ .32 

12.44+ .34 

14,I2±4.17 

12.99+ .34 

11.80+ .82 

11.98+ 1.48 

13.65+ .32 

12.09+ 1.48

12. 73 + 0.77

3.94 + 0 

3.88 + 

3.74 + 

4.60 + 2 

4.01 + 

3.42 + 

3 52 + 

4.36 + 

3.56±

3.89±0

.17 

.16 

.17 

.09 

.17 

.41 

.74 

.16 

.74

.39

-0.16 + 0. 

-.22+ . 

- .05 + . 

.49+2. 

-.07+ . 

.10+ .

.34+ .

.33+ .

1 Q j

-0.03 + 0.

17 

16 

17 

09 

17 

,41 

74 

16 

,74

27

 

 

 

Sulfide present

19

,4 _______________

n

<. 0

 '

3. 

5. 

3. 

3.

2

33 + 0. 

74+ . 

91 + . 

35 + 4. 

,44± .

,35 

32

17 

,48

-0.81 +0

.40 + 

- .51 +

- .79+2 

1.25 +

.18 

.16 

.16

.09

.24

4.46 + 0.

3.44+ .

-4.61+ . 

-4.82 + 2. 

-5.17+ .

,18 11.01+0.24 

,16 6.81 + 1.71 

16 10.78+ .86 

,09 8.75 + 2.16 

24 10.41 + .48

1.3h + 0 15

O 1 , (Ji-i

1.00+ .43 

.41 +2.10 

.55+ .29

-0.59 + 0.61

but undersaturated in strongly reducing waters where 

sulfide is present. Oversaturation is due to the highly 

ordered, crystalline structures of these minerals, which 

prevent them from being precipitated directly from solution. Instead, ferric iron precipitates from solution as 

amorphous ferric hydroxide [Fe(OH)3(am)] with crystallization to the more stable goethite or hematite occurring 

over a period of time (Langmuir and Whittemore, 1971). 

Using a Ksp of 10~385 for the reaction 13, which is 

characteristic of aged Fe(OH)3 precipitates, saturation 

indices for ferric hydroxide indicate near saturation in 

oxidizing and moderately reducing waters, and undersaturation in strongly reducing, sulfidic waters. Whittemore and Langmuir (1975) concluded that log IAP 

values near 37 are indicative of actively precipitating 

Fe(OH)3(am), whereas lower values of 39 to 44 imply 

dissolution of, or equilibrium with, more stable oxyhydroxides. The occurrence of log lAP's near 39, or 

Si's of zero or slightly less as calculated previously, in 

oxidizing and moderately reducing Kootenai ground 

waters, is indicative of equilibration with ferric oxyhydroxides of moderate crystallinity and stability, 

rather than active precipitation of these phases.

Equilibrium with a moderately crystallized phase would 

be expected in the Kootenai Formation, where previously deposited ferric oxyhydroxides are presently dissolving in modern ground waters.

Undersaturation with respect to Fe(OH)3(am| and 

goethite in strongly reducing waters is indicative of the 

control of iron concentrations by some other phase, most 

commonly iron sulfides. Pyrite was not identified in 

petrographic or X-ray analyses of the Kootenai Formation nor were any other sulfide minerals. However, 

because the samples analyzed were collected from the 

outcrop, pyrite originaly present would have been oxidized to ferric oxyhydroxides. Pyrite is ubiquitous in 

strongly reducing, aqueous environments (Berner, 1970), 

and should therefore be present in the Kootenai Formation where these conditions prevail. Saturation indices 

show gross oversaturation of this mineral in sulfidic 

waters, whereas amorphous ferrous sulfide (FeS(am|) is 

near saturation. Deviations from zero for FeS(am) saturation indices are most likely the result of determinate errors in the sulfide analyses. Ferrous sulfides have been 

observed to precipitate from solution in reducing marine 

(Berner, 1964a; 1964b) and lacustrine (Doyle, 1968b)
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C31

sediments, but not in the ground-water environment. 

Barnes and Clark (1969), however, have reported the 

presence of mackinawite (FeS), a crystallized ferrous 

sulfide, as encrustations on steel ground-water well casings. The precipitation of amorphous FeS, followed by 

aging and crystallization in the presence of S° to form 

pyrite (Berner, 1970) appears to be a feasible mechanism 

to explain the above saturation indices, which reflect 

equilibrium with amorphous ferrous sulfide.

REACTION SIMULATION

Consideration of major and trace element geochemistry, Eh and pH variations, and saturation indices in 

ground waters of the Second Cat Creek sandstone, in 

association with aquifer mineralogy and petrography, 

has led to the proposal of several mineral phases and 

chemical reactions that may control ground-water 

chemistry. The aluminosilicates illite, kaolinite, montmorillonite, and allophane may be near saturation and 

may control the chemistry of waters in the unconfined 

outcrop area. Following the onset of confined conditions, slight oversaturations with respect to calcite, 

phase relations in the Mg-Ca-H2O-CO2 system, and 

the presence of calcareous cement, suggest precipitation 

of calcite, whereas saturation with dolomite implies 

simultaneous equilibration with this carbonate. Carbon 

isotopic ratios also suggest control of dissolved carbonate concentrations by carbonate minerals.

Oxidation potentials decrease from initially oxidizing 

values, then level off at moderately reducing values. 

Consumption of O2 , NO~, and DOC within a short 

distance from the outcrop, coupled with increasing concentrations of dissolved N2 and CT, is indicative of 

biologically mediated O2 and NO~ reduction reactions. 

The paucity of organic carbon in the aquifer, as well as 

the absence of H2S and the simultaneous saturation 

with respect to both Fe(OH)3(am) and siderite in most 

waters of the Second Cat Creek sandstone, indicates 

that oxidation potentials may be poised by these two 

iron phases at levels too high to allow for either sulfate 

reduction or methanogenesis. One sample taken from 

the fractured South Moccasin Uplift shows the presence 

of H2S, which may indicate that reduced gases are free 

to migrate upward from lower aquifers along these fractures. Sulfur isotopic ratios give no evidence of sulfate 

reduction occurring at this sampling location.

Sulfate and chloride concentrations are significantly 

higher in the fractured region to the north of Big Spring 

Creek than to the south, and sulfur isotopic ratios indicate a significantly different source of sulfate on 

either side of the creek. The occurrence of high sulfate 

and high chloride concentrations in aquifers underlying

the Kootenai Formation suggests that the composition 

of waters in the Second Cat Creek sandstone north of 

Big Spring Creek may be influenced by upward leakage 

along faults and fractures.

Finally, uniformly decreasing log molar divalentmonovalent cation ratios are indicative of cation 

exchange. The exchange of monovalent sodium ions adsorbed on clay surfaces for divalent calcium and 

magnesium ions in solution influences solution-mineral 

equilibria, and therefore must be considered during 

reaction simulations.

The effects of the above mechanisms in controlling 

solution chemistry, both alone and in combination, have 

been investigated through reaction simulations using 

the computer program PHREEQE (Parkhurst and 

others, 1980). Attempts have been made to duplicate 

measured variations in chemistry from well to well 

along the flow path by equilibrating an initial solution 

with desired mineral phases. The varying degrees of 

success obtained using different controlling phases is interpreted as a measure of the likelihood that a particular 

reaction or mineral phase controls solution chemistry. 

Comparisons of pH and CT between measured and modeled chemistries are particlarly useful in the determination of controlling phases because both parameters vary 

over a wide range, depending on the choice of controlling reactions. However, final evaluation of controlling 

phases is based on agreement of all constituents considered in the computer equilibrium model with analyzed 

concentrations.

To consider the effects of mineral equilibrium in 

ground waters of the Second Cat Creek sandstone, ionexchange reactions must first be simulated. This was 

accomplished by fixing (aNa+ )2/(ac +2) and (aNa+)2/(a +2) 

molar activity ratios in the8modeled solutions to tneir 

calculated values for each sample along the flow path. 

In effect, this fixes the activities of monovalent and 

divalent cations on the exchange surface to a particular 

value for each sample, and implies an infinite reservoir 

of exchanger as can be seen from the exchange equilibrium expressions (reactions 28 and 29). Such a treatment eliminates the need to consider activities of ions 

on a solid surface, information that is currently 

unavailable.

As there is considerable evidence of leakage into the 

Second Cat Creek sandstone north of Big Spring Creek, 

areas north and south of the creek were treated as 

distinct flow systems. The southern flow path is 

discussed first, being the simpler of the two systems.

The initial modeling .solution is similar in major element composition to Montana rainwater (Junge and 

Werby, 1958), and has been equilibrated, using 

PHREEQE, with atmospheric partial pressures of O2 

and CO2 to approximate actual Eh and pH. Sulfate,
C32 REGIONAL AQUIFER-SYSTEM ANALYSIS

chloride, and nitrate were added to the solution as 

equivalent amounts of gypsum, halite, and calcium 

nitrate, in order to match measured concentrations, 

simulating the dissolution of trace amounts of these or 

similar compounds in the unsaturated zone. The resulting solutions were then equilibrated with different 

minerals to investigate the influence of these phases in 

controlling solution chemistry. The minerals chosen as 

plausible controlling phases are: (1) calcite; (2) dolomite; 

(3) calcite and dolomite; (4) allophane; (5) kaolinite and illite; (6) kaolinite, illite, and Mg-montmorillonite; and 

(7) kaolinite, illite, and Ca-montmorillonite. In addition 

to these phases, the model solutions were also equilibrated with Fe(OH)3(am) in the unconfined aquifer, and 

with Fe(OH)3(am) and siderite in the confined aquifer 

system, to allow simulation of iron concentration and 

Eh poising. A Ksp of 10~385 , which is characteristic of an 

aged precipitate, was chosen for Fe(OH)3(am).

Initial PHREEQE simulations demonstrate the effects of equilibrating the rainwater solution with these 

groups of mineral phases. Carbon dioxide was added to 

the solution prior to equilibrating with the minerals, 

simulating the dissolution of CO2 in the soil zone. To 

match the calculated CT concentration in sample J10, 

CO2 was added in concentrations of approximately 1.5 

millimoles per liter for the carbonate simulations and 

3.0 millimoles per liter for the aluminosilicate simulations. Dissolution of these amounts of CO2 would 

require equilibration of infiltrating waters with CO2 partial pressures of 10" 157 to 10" 128 atm in the unsaturated 

zone. Garrels and MacKenzie (1971, p. 141) report soil 

zone Pco values on the average of 10~20 and as high as 

10~05 atm, which is consistent with the model results.

Following equilibration with mineral phases and 

the appropriate ion-exchange constants for J10, 

0.537 mmol/L carbon with an oxidation state of zero 

was added to the solution to allow for the reduction of 

O2 and NOg~ present in the outcrop area but absent in 

samples taken downdip. The simulations then equilibrated the solution with the minerals and the appropriate 

ion-exchange constants for the downdip samples. As 

shown in figure 22, these simulations clearly demonstrate that insufficient dissolved carbonate is produced 

by the dissolution of soil zone CO2 , the reduction of O2 

and NO~ by organic carbon, and the equilibration with 

minerals and ion exchange to account for the observed 

increases in CT.

Subsequent simulations required an additional carbon source, if modeled and observed CT concentrations 

were to agree. Although the exact mechanism of this 

CO., production remains uncertain, it seems likely that 

the CO., is an oxidation product of organic carbon 

generated via sulfate reduction processes, or by the 

coalification of lignite, in strata adjacent to the aquifer.

& 7.0- 

in

d

^

_i 60-

EXPLANATION

CALCULATED CONCENTRATION 

MODELED CONCENTRATION

CALCITE + DOLOMITE CALCITE, DOLOMITE 

KAOLINITE + ILLITE

KAOLINITE + ILLITE +MONTMORILLONITE

0 10 20 30 40 

DISTANCE ALONG LINE A-A', IN KILOMETERS

FIGURE 22. Comparison of calculated total inorganic carbon concentrations with modeled concentrations for mineral equilibration 

with no CO2 dissolution, following the reduction of O2 and NOg~, 

southern flow path of the Second Cat Creek sandstone, eastern 

Judith Basin, Montana. All simulations include equilibration with 

calcium-sodium exchange, magnesium-sodium exchange and amorphous ferric hydroxide, as well as siderite in Jl 1, J13, and J15.

Quantities of CO2 were dissolved along the simulated 

flow path to match observed CT concentrations.

The model results, presented in figure 23, indicate 

that pH values for the confined system are successfully 

approximated only for the carbonate-ion exchange 

simulations. Although the kaolinite-illite-ion exchange 

simulation results in a pH value close to that measured 

in the sample taken from the unconfined aquifer, the 

aluminosilicates do not appear to control solution 

chemistry in the confined portions of the aquifer. Deviations between measured and modeled pH's for sample 

J10 using the kaolinite-illite-ion exchange model may be 

the result of the variable compositions and stabilities of 

these aluminosilicates. Refined geochemical simulations 

of the unconfined portion of the Second Cat Creek sandstone await an improved characterization of the compositions and stabilities of aluminosilicates present.

Both the calcite-ion exchange and the calcite-dolomiteion exchange models approximate pH to within 0.1 

units of" the field-measured value. The presence of 

dolomite in the aquifer and its instability with respect 

to calcite in the ground waters sampled (fig. 14) suggests that the calcite-dolomite-ion exchange simulations, which requires mass transfer characteristic of 

dedolomitization (reaction 23), is the more probable of 

these two models. In this calcite-dolomite-ion exchange 

model, dolomite dissolves in response to CO2 production 

and Na/Mg exchange, simultaneously precipitating 

calcite:
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C33

EXPLANATION

 MEASURED pH 

/ MODELED pH

KAOLINITE » ILLITE

0 10 20 30 40 

DISTANCE ALONG LINE A-A'. IN KILOMETERS

FIGURE 23. Comparison of field-measured pH values with modeled 

pH values for mineral equilibration with CO2 dissolution, southern 

flow path of the Second Cat Creek sandstone, eastern Judith Basin, 

Montana. All simulations include equilibration with calciumsodium exchange, magnesium-sodium exchange and amorphous 

ferric hydroxide, as well as siderite in Jl 1, J13, and J15.

CaMg(CO3 )2 +H2CO^+Na2EX 

=CaCO3+2HCO3-+2Na+ +MgEX.

In addition, a net excess of calcium in solution is exchanged for sodium on the exchanger surface (reaction 

28). The addition of illite and kaolinite to this model 

results in only minimal changes in mass transfer and 

final solution composition, suggesting that the importance of these aluminosilicates as ion exchangers far 

overshadows the influences of their solubilities on major 

element chemistry. Direct comparisons of simulated 

and analyzed concentrations for the Second Cat Creek 

sandstone, as well as the computed mass transfers required to attain these variations, are presented in 

figure 24.

At this point, it is important to emphasize that the 

mathematical solution to the set of simultaneous equations required to characterize mass transfer is nonunique. Other reaction schemes may be proposed that 

would adequately describe the geochemical evolution of 

ground waters in the Second Cat Creek sandstone. The 

mass transfer model proposed here is therefore only a 

reasonable hypothesis not a unique mathematical 

solution.

The northern flow path, which includes all samples 

taken from wells located north of Big Spring Creek, was 

modeled to simulate upward leakage from the underlying Morrison Formation through numerous faults and 

fractures in the area. Suitable samples of ground water

from the Morrison within the Judith Basin study area 

were unavailable; therefore, analyses of waters from two 

wells in the Morrison Formation located in Grassrange, 

Montana, 45 km east of Lewistown on the northern 

flank of the Big Snowy Uplift, were used. Although it is 

unlikely that these waters are identical to those in the 

area modeled, they are of sufficiently similar hydrogeochemical origin, prevalent chemical character, and 

dissolved solids content to approximate the composition of Morrison Formation waters within the study 

area. In addition, the pH and anionic composition of 

these Morrison waters is similar to that of springs and 

wells that discharge from the Madison Group (table 8). 

Therefore, simulated leakage of these Morrison Formation waters into the Second Cat Creek sandstone would 

also approximate the effects of similar amounts of 

leakage from the deeper Madison aquifer, although 

calculated mass transfers would differ because of 

dissimilar cationic concentrations in the waters of the 

Madison and Morrison Formations.

Sample Jll was chosen as the starting composition 

for the northern flow path simulations because of its 

similarity to the northern flow path outcrop sample J24, 

which was unsuitable for geochemical modeling. Using 

PHREEQE in the mixing mode, sufficient Morrison 

water was added to the simulated Jll water to match 

analyzed sulfate concentrations along the northern flow 

path. Sulfate was considered a conservative constituent 

in these simulations, based on evidence suggesting that 

sulfate was neither precipitating as a mineral phase nor 

being reduced to sulfide in the Second Cat Creek 

sandstone.

The mixed solution was then equilibrated with the 

same phases that adequately described southern flow 

path geochemistry, calcite and dolomite, as well as ion 

exchange, Fe(OH)3(am), and siderite. In addition, the solution was equilibrated with FeS)am), rather than siderite 

and Fe(OH)3(am), to simulate the chemistry at well J17. It 

was discovered that although 0.20 millimoles of CO2 per 

liter had to be dissolved to adequately describe the 

geochemistry at well J14, 0.26 and 0.65 millimoles of 

CO9 per liter had to be exsolved from the solution to approximate the chemistry at well J17, and 0.65 millimoles of CO2 per liter had to be exsolved from the solution to approximate the chemistry at well J12. Because 

the Morrison water composition is only approximate, interpretation of the apparent degassing is uncertain. It is 

entirely possible that the waters leaking into the 

Kootenai are higher in sulfate and lower CT than in the 

model; however, most of the analyses from lower 

aquifers suggest a predominance of pH values and 

bicarbonate and sulfate concentrations equal to or lower 

than those used in the model (table 8). The presence of 

waters with compositions that would both increase
C34 REGIONAL AQUIFER-SYSTEM ANALYSIS

DISSOLVE:

0.24 mmol/L 02 

0.05 mmol/L HALITE 

0.56 mmol/L GYPSUM 

0.12 mmol/L Ca(N03 ) 2

2.01 mmol/L C02 

0.54 mmol''L CH 2 O

2.02 mmol/L DOLOMITE 

5.34 umol/L Fe(OH) 3(am) 

4.14 umol/L SIDERITE

PRECIPITATE:

1.92 mmol/L CALCITE

ADSORB:

1.33 mmol/L Mg

DESORB:

0.42 mmol/L Ca 

1,82 mmol/L Na

SAMPLE J11

Ca

Mg

Na

Cl

so4

C T

Fe

PH

Eh

MEAS,

1.1

0.62

1.8

0.04

0.64

4.7

.0093

7.70

0.038

MODEL

1.2

0.69

1.9

0.05

0.59

4.7

.0095

7.76

0.032

 ^ 

DISSOLVE: 

0.03 mmol/L HALITE 

0.74 mmol 'L C02 

1.02 mmol/L DOLOMITE

PRECIPITATE: 

1.05 mmol/L CALCITE 

0.01 umol/L Fe(OH) 3(am) 

8.46 umol/L SIDERITE

ADSORB: 

1 .08 mmol/L Ca 

1.63 mmol/L Mg

DESORB: 

5.42 mmol/L Na

SAMPLE J13

Ca

Mg

Na

Cl

so4

C T

Fe

PH

Eh

MEAS.

0.12

0.08

7.4

0.08

0.47

6.5

.0005

8.48

-0.062

MODEL

0.12

0.08

7.4

0.08

0.59

6.5

.0010

8.53

-0.068

 »>

DISSOLVE: 

0.06 mmol/L HALITE 

0.08 mmol/L GYPSUM 

0.40 mmol/L C02 

0.42 mmol 'L DOLOMITE

PRECIPITATE: 

0.37 mmol/L CALCITE 

0.00 umol/L Fe(OH) 3(am) 

0.36 umol/L SIDERITE

ADSORB:

0.17 mmol/L Ca 

0.46 mmol/L Mg

DESORB: 

1,20 mmol/L Na

SAMPLE J15

Ca

Mg

Na

Cl

so4

C T

Fe

PH

Eh

MEAS.

0.08

0.04

8.3

0.14

0.67

7.3

.0005

8.55

-0.077

MODEL

0.09

0.04

8.7

0.14

0.67

7.3

.0007

8.64

-0.088

FIGURE 24. Calculated mass transfer for the calcite-dolomite-cation exchange-CO2 dissolution model of groundwater evolution along the southern flow path of the Second Cat Creek sandstone, eastern Judith Basin, Montana. Concentrations in millimoles per liter (mmol/L) and micromoles per liter (/xrnol/L); Eh in volts.

sulfate and decrease CT, with respect to the composition 

of water from well J14, appears unlikely. At the same 

time, correspondence of the aquifer outcrop with the 

area of CO2 exsolution (fig. 3) is an indication that such 

outgassing may indeed occur. In addition, recharge to 

the Second Cat Creek sandstone in the South Moccasin 

Mountains is negligible, which would allow for an upward rather than downward mass flux in the unsaturated zone. For the required degassing to occur, the Pco

of the unsaturated zone would have to be approximately 

10~2fs atm, as compared to the atmospheric Pco of

10~ 35 atm. Although the unsaturated zone Pco required

for degassing is somewhat low, it does not appear to be 

prohibitively so, and may allow CO9 outgassing to occur 

at the South Moccasin Mountains.

Further support for upward leakage of water similar 

to the composition modeled is given by the close agreement of another conservative constituent, chloride.

Modeled and observed chloride concentrations differ by 

a maximum of only 0.03 millimoles per liter in the three 

waters simulated in this manner, and indicate that the 

leaking waters have a m^ /mcl ratio identical to that of 

the Morrison waters.

Mass transfer calculated for the mixing simulations 

also shows the influence of dedolomitization (fig. 25), 

but instead of prevalent Na/Mg exchange, Ca/Mg exchange removes from solution the magnesium derived 

from the dissolution of dolomite, and supplies the 

calcium required for the precipitation of excess calcite 

due to CO2 outgassing:

CaMg(CO3 )2 +Na2EX+H2CO3' 

=CaCO3 +2Na+ +MgEX+2HCO2HCO3-+2Na+ +CaEX 

=CO2(g)t+CaCO3 +H2O+Na2EX.
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C35

MIX:

92% J11 water

8% Mornson water 

DISSOLVE: 

0.20 mmol'L C02 

0.96 mmol/L DOLOMITE 

0.01 umol/L Fe(OH)3(am) 

0.50 umol/L SIDERITE

PRECIPITATE:

1.76 mmol 'L CALCITE

ADSORB:

0.81 mmol 'L Mg

0.28 mmol 'L Na 

DESORB:

0.95 mmol 'L Ca

SAMPLE J11

Ca

Mg

Na

Cl

so4

CT

Fe

PH

Eh

MEAS.

1,1

0.62

1..8

0 U 04

0.64

4.7

.0093

7.70

0.038

1

MODEL

1.2

0,69

1.9

0.05

0.59

4.7

.0095

7.76

0.032

, 4

SAMPLE J14

Ca

Mg

Na

Cl

so4

CT

Fe

PH

Eh

MEAS.

1.3

0.78

2.7

0 04

093

5.2

0.016

7.82

0.023

MODEL

1.3

0.78

2.7

007

0.93

5.2

0.010

7.68

0.036

MORRISON

Ca

Mg

Na

Cl

so4

C T

PH

1

MEAS.

0.16

0.07

16.3

3.2

5.0

7.0

8.0

r

MIX: * 

81% J14 water 

19% Morrison water

DISSOLVE:

0.58 mmol/L DOLOMITE 

0.02 mmol/L H 2S

EXSOLVE:

0.26 mmol/L C02

PRECIPITATE:

1 .40 mmol/L CALCITE 

7.40 umol/L FeS(am) 

ADSORB; 

0.45 mmol/L Mg 

0.86 mmol/L Na

DESORB: 

0.88 mmol/L Ca

^

SAMPLE J1 7

Ca

Mg

Na

Cl

so4

C T

Fe

pH

Eh

MEAS.

1.1

0.74

4.4

0.11

1.7

5.1

.0025

7.60

0.214

MODEL

1.1

0.77

4.4

0.12

1.7

5.1

.0012

7.68

-0.235

MIX: * 

52% J17 water 

48% Morrison water 

DISSOLVE: 

0.18 mmol/L DOLOMITE 

0.06 mmol/L Fe(OH)3(arn) 

EXSOLVE: 

0.65 mmol/L COo 

PRECIPITATE: 

0.84 mmol 'L CALCITE 

0.06 mmol 'L SIDERITE

ADSORB:

0.48 mmoL L Mg

DESORB:

0.23 mmol 'L Ca

4

SAMPLE J12

Ca

Mg

Na

Cl

so4

C T

Fe

PH

Eh

MEAS,,

0.22

0.13

10.4

0.21

3.3

4.8

.0025

8.49

0.101

MODEL

022

0.14

10.6

0.22

3.3

4.8

.0017

8.42

-0.100

FIGURE 25. Calculated mass transfer for the calcite-dolomite-cation exchange-Morrison leakage model of ground-water evolution along the 

northern flow path of the Second Cat Creek sandstone, eastern Judith Basin, Montana. Concentrations in millimoles per liter (mmol/L) 

and micromoles per liter (/anol/L); Eh in volts.

Increases in calcium and magnesium concentrations of 

the upward leaking water would require increased 

desorption of sodium in combination with increased adsorption of magnesium and calcium to maintain equilibrium with ion exchange.

The quality of fit obtained using the dedolomitization 

model is shown in figure 26. In these plots, precisions of 

the analytical measurements at the 95 percent significance level (table 3) are superposed on the modeled concentrations, allowing comparisons to be made between 

analyzed and modeled values. The comparison is generally very good, although several deviations are evident 

for which no explanations can be offered. Once again, 

this mass transfer model is a nonunique solution to a set 

of simultaneous equations. The geochemical evolution 

of ground-water composition along the northern flow 

path may be adequately described without invoking upward leakage, but since leakage appears to be occurring, 

the mixing model is considered most plausible.

CARBON ISOTOPIC EVOLUTION

The evolution of carbon isotopic composition in 

ground waters of the Second Cat Creek sandstone was 

simulated through the use of a FORTRAN computer 

code of the model presented by Wigley and others 

(1978). This carbon isotopic evolution model considers 

the effects of mass transfer from an arbitrary number of 

carbonate sources and sinks in combination with equilibrium fractionation between the solid, gas, and 

aqueous phases in contact with the aquifer to calculate 

changes in the <5 13C value of ground-water carbonate. 

The same model may also be used to calculate the effects of mass transfer on the activity of 14C in ground 

waters, which then allows for adjustment of raw 14C 

ages, based on the dilution and precipitation of the 

radioactive isotope.

Other models of 14C evolution in natural systems, including those of Vogel (1970), Tamers (1975), Mook
C36 REGIONAL AQUIFER-SYSTEM ANALYSIS

EXPLANATION

RANGE OF MEASURED CONCENTRATION 

AT 95 PERCENT CONFIDENCE

 MODELED CONCENTRATION, 

SOUTHERN FLOW PATH

A MODELED CONCENTRATION, 

NORTHERN FLOW PATH

CONCENTRATIONS IN M1LLIMOLES PER LITER 

2.0-1

(J

co

LJJ

z

0.0

1.0-1

0.5-

0.0'

Q 

O

CO

12.0i

6.0-

J10 J15

o.o0 10 20 30 40 

DISTANCE ALONG LINE A-A', IN KILOMETERS

0.02-,

O 0.01-I

CJ 

CJ

0.00

QC

O

<

O

D

CO

LLJ 

Q

cc. 

O

I

CJ

4.0-

0.0-

4.0-1

2.0-

J_L

0.0

0.2n

0.1-

0.0

J10 J11 J14 J13J17 J12 J15

I i ^ i

0 10 20 30 40 

DISTANCE ALONG LINE A-A', IN KILOMETERS

FK iURE 26. Comparison of measured and modeled concentrations of selected constituents in waters from the Second 

Cat Creek sandstone of the Kootenai Formation, eastern Judith Basin, Montana.

(1976), and Fontes and Gamier (1979), neglect the effects of 14C precipitation in carbonate minerals, and 

thereby lead to adjusted 14C ages that are too old when 

applied to aqueous systems where incongruent dissolution of carbonates is an important mass transfer 

mechanism.

Unfortunately, samples collected from the northern 

flow path are unsuitable for isotopic modeling in that 

they appear to represent a mixture of waters with different evolutionary histories. These samples could be 

modeled with confidence only if the carbon isotopic composition of the upward leaking waters were known. 

However, a comparison of observed and calculated 6 13C

values for the southern flow path allows a check on the 

agreement of isotopic and mass transfer models. In 

addition, adjusted 14C ages for the southern flow path 

may be used to calculate flow rates and residence times 

in the unfractured aquifer.

The 6 13C values for carbonate sources were estimated 

from measurements reported in the literature for comparable materials. A 6 14C value of -28±4 (Craig, 1954) 

was chosen for sedimentary organic carbon and the CO2 

derived from it, and a freshwater carbonate <5 13C of 

 4.8±5.1 was used for Kootenai dolomite and calcite 

(Keith and Weber, 1964). Keith and Weber (1964) also 

reported 613C values with a mean of 2.50 for fresh-
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C37

0 EXPLANATION

 MEASURED ISOTOPIC RATIO

MODELED ISOTOPIC RATIO

o

co -10

J11 J13

£ 0 10 20 30 40 

<

o

DISTANCE ALONG LINE A-A', IN KILOMETERS

FIGURE 27. Comparison of measured <5 13C values and the range of 

modeled <5 13C values for water from the southern flow path of the 

Second Cat Creek sandstone, eastern Judith Basin, Montana The 

range of modeled <5 13C values reflects the uncertainty in the actual 

<5 13C compositions of aquifer carbon sources.

water Kootenai limestones and dolomites from several 

areas in Montana, which are consistent with the range 

of values used. Employing the temperature-dependent 

fractionation factors of Deines and others (1974), the 

equations of Wigley and others (1978) were applied to 

the <5 13C carbonate source values estimated above, the 

computed mass transfer between samples (fig. 24), and 

the initial observed <5 13C of the ground water (table 12), 

to compute the <5 13C values of three successive samples 

on the southern flow path, assuming the incongruent 

dolomite dissolution model. Agreement between predicted and observed <5 13C values is excellent within the 

uncertainty of the isotopic composition of the carbonate 

sources (fig. 27), and therefore supports the validity of 

the CO2 dissolution and dedolomitization model for the 

Second Cat Creek sandstone.

The equations of Wigley and others (1978) were applied to the measured carbon isotopic compositions 

(table 12) and the computed mass transfer (fig. 24) in an 

attempt to adjust raw 14C ages for the effects of dilution 

and precipitation that occur during ground wateraquifer interactions. Using this approach, an age of 

4,800 years was calculated for sample water Jll, and 

> 20,000 years for sample water J13. The 14C concentration in sample water J15 was too low to allow for a 

reliable determination of the raw age. Radiometric flow 

rates for the Second Cat Creek sandstone, calculated for 

a flow distance from the outcrop to the well measured 

perpendicular to the potentiometric contours, compare 

favorably to the previously derived hydrologic flow 

rates for the Kootenai Formation (table 17).

T'XBLK 17. Comparison of radiometric and hydrologic flow rates for ground waters of 

the Koutenm Formation, eastern Judith Basin, Montana

Sample

Hydrologic flow rate,

Kootenai Formation

(meters per year)

Radiometric flow rate,

Second Cat Creek sandstone

(meters per year)

J13 

J15

2.6 

1.7

2.3

POWDER RIVER BASIN STUDY AREA 

PHYSIOGRAPHY

The physical characteristics of the Powder River 

Basin are the result of several factors, the most important of which is the regional geologic structure. Dominant features of the topographic basin are the rugged 

Big Horn Mountains to the west, the elevated but more 

subdued Black Hills on the east, and the sprawling 

basin between them (fig. 28).

TOPOGRAPHY

The terrain of the basin floor is composed of hilly to 

rugged uplands, wide rolling valleys, and isolated badlands. In the study area, the basin floor lies at elevations roughly between 1,200 to 1,300 meters; the Black 

Hills Uplift reaches moderate elevations of approximately 1,400 to 1,800 meters above sea level. The uplift 

is characterized by tree and grass-covered peaks and 

dissected plateaus, the elevated areas giving way to 

lower sage-covered hills and the basin floor to the west 

(Wyoming Water Planning Program, 1972).

DRAINAGE

The study area (shown in detail in fig. 29) lies within 

the Missouri River drainage basin and is drained by 

tributaries of the Yellowstone, Cheyenne, and Little 

Missouri Rivers. In the three tributaries in the study 

area, the Little Powder, Belle Fourche, and Little Missouri Rivers, flow is extremely variable, ranging from 

virtually nonexistent at times during the winter to occasional floods of greater than 100 m3/s during the early 

summer. Keyhole Dam, located approximately 20 km 

northeast of Moorcroft, impounds runoff in the upper 

Belle Fourche River for controlled release to farms and 

ranches downstream later in the season (Whitcomb and 

Morris, 1964).
C38

108°

REGIONAL AQUIFER-SYSTEM ANALYSIS

107° 106° 105° 104°

OUTH DAKOTA 

NEBRASKA

0 20 40 60 KILOMETERS

43'

42°-

EXPLANATION

UNDIFFERENTIATED PRECAMBRIAN ROCKS

Modified from Robinson and others, 1964

SYNCLINE MAJOR TOWNS AND CITIES

% ANTIPI INF - OUTLINE OF ANTILLINb STUDY AREA

MONOCLINE 

FIGURE 28. Structural and physiographic features of the Powder River Basin, Wyoming and Montana.

CLIMATE

The study area has a semiarid climate with annual 

average precipitation at Moorcroft of 35 cm, at Devils 

Tower of 41 cm (Whitcomb and Morris, 1964), and at 

Gillette of 40 cm (Wyoming Water Planning Program, 

1972). Greater than 50 percent of this precipitation occurs between April and July. Rainfall in the spring and 

early summer occurs as brief but rather heavy showers,

and light snowfalls and occasional heavy snows are 

common throughout the winter (Whitcomb and Morris, 

1964).

Mean annual temperature at Moorcroft is 6.9 °C and 

at Devils Tower is 7.9°C (Whitcomb and Morris, 1964); 

at Gillette it is 7.3°C (Wyoming Water Planning Program, 1972), with the average monthly temperature 

during July being roughly 13 °C higher than that for 

January (Whitcomb and Morris, 1964).
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS

105° 30' 105° 15' 105°OO' 104°45' 104° 30'

C39

45° 00'

44° 45' -

44°30' -

44° 15' -

EXPLANATION

FOX HILLS SANDSTONE 

(UPPER CRETACEOUS)

MAJOR TOWNS

 P19 LOCATIONS OF LANCEFOX HILLS WELLS USED 

FOR THIS STUDY

TRACE OF SECTION 

SHOWN IN FIGURE 30

Keyhole 

Reservoir

44°00' -

Geology from Hodson and others. 1973.

FIGUKK 29. Index map of the Powder River Basin study area showing major towns and perennial streams, 

outcrop of the Fox Hills Sandstone and the locations of wells sampled.

GEOLOGY

STRUCTURE

The Powder River Basin is a northward-trending 

structural basin that has evolved as a result of the complex development of the Rocky Mountain region. A

trough lying roughly parallel to the Big Horn Uplift 

(fig. 28) identifies the axis of the basin, where roughly 

5,000 m of Phanerozoic sediments have accumulated. 

The study area occupies the northeastern portion of the 

basin adjacent to the Black Hills Uplift, a broad 

northward-trending anticline. The northwest-trending 

Black Hills monocline divides the basin to the west
C40 REGIONAL AQUIFER-SYSTEM ANALYSIS

B'

3000-

C/3

£ 1500-

K 

LJJ

S

z

~ 0 -

z

o

K

> -1500

-3000-

VERTICAL EXAGGERATION 4X

BLACK HILLS 

MONOCLINE

UNDIFFERENTIATED MESO2OIC 

AND PALEOZOIC SEDIMENTS

PRECAMBRIAN BASEMENT National Geodetic Vertical Datum of 1929

Modified from Breckenridge and others, 1974.

FIGURE 30. Generalized geologic section of the Powder River Basin study area, Wyoming, through line B-B' in

figure 29.

from the Black Hills Uplift to the east. The monocline 

consists of a belt of steeply dipping sedimentary strata, 

2.5 to 9.5 km wide, across which the structural relief is 

about 1,800 m. Dips are to the west or southwest along 

the monocline, ranging from 20 degrees to nearly vertical. The outcrop of the Fox Hills Sandstone lies parallel to the monocline, and delineates the western extent 

of steeply dipping strata. To the west, the dips of the 

Lance Formation and younger strata flatten abruptly, 

ranging from 2 to 5 degrees toward the axis of the basin 

(fig. 30) (Robinson and others, 1964).

The Belle Fourche Arch, a subtle northeast-trending 

paleo-arch, extends across the southern part of the 

study area. According to Slack (1981), the arch is the 

result of differential vertical uplift along numerous 

northeast-trending structural lineaments (fig. 31). 

Underlying zones of weakness in the Precambrian basement, thought to be shear zones of Precambrian age, 

allow for periodic movement along the lineaments, 

resulting in structural offsets in the Black Hills 

monocline, welldefined topographic escarpments, and 

linear drainage courses. Stratigraphic evidence suggests 

that these structural lineaments have been rejuvenated 

periodically throughout the Phanerozoic and that the 

subtle movements involved have affected depositional 

environments and, therefore, affected porosities and permeabilities along the crest of the arch (Slack, 1981).

STRATIGRAPHY

Sedimentary rocks exposed on the western flank of 

the Black Hills Uplift have a combined thickness of 

about 3,400 m and range in age from Mississippian to 

Quaternary.

The Upper Cretaceous Pierre Shale is a 625- to 

825-m-thick sequence of dark gray marine shales and

bentonites with minor interbedded sandstone units. 

Conformably overlying the Pierre Shale is the Fox Hills 

Sandstone, and Upper Cretaceous sequence of sandstones and shales that were deposited in an environment ranging from marginal marine to estuarine or tidal 

flat environment (Dodge and Spencer, 1977). The Fox 

Hills ranges from 38 to 60 m in thickness, with crossbedded sandstones at the top of the formation grading 

downward into sandy shales and culminating at the 

underlying Pierre Shale. The Fox Hills is poorly exposed in outcrop, generally appearing as low grassy 

ridges. Where exposed, the sandstones are gray to light 

brown, fine to medium grained, and soft and thin 

bedded. The rock has a slabby appearance, and locally 

contains numerous calcareous or ferruginous concretions (Whitcomb and Morris, 1964).

Conformably overlying the Fox Hills Sandstone are 

the freshwater, fluvial deltaic strata of the Upper Cretaceous Lance Formation (Dodge and Spencer, 1977). The 

Lance is comprised of 150 to 500 m of yellowish gray, 

fine- to medium-grained lenticular crossbedded sandstones interbedded with darker gray shaley siltstones 

and sandy claystones. Black carbonaceous shales are 

common at the gradational contact with the Fox Hills 

Sandstone.

The nonmarine sandstones, shales, coals, and tuffs of 

the Tertiary Fort Union and Wasatch Formations overlie the Lance Formation and crop out at the surface over 

much of the basin. A geologic map of the study area 

showing the outcrop areas of stratigraphic units pertinent to this study is presented in figure 32.

MINERALOGY OF THE LANCE-FOX HILLS AQUIFER

Mineralogy of the Fox Hills Sandstone (tables 18 and 

19) was determined by X-ray diffraction analysis of four 

drill-core and two outcrop samples obtained from the
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C41

108°

1 SOUTH DAKOTA 

I NEBRASKA

43°-

42°-

Modified from Slack. 1981 

EXPLANATION 0 20 '40 60 80 100 KILOMETERS

UNDIFFERENTIATED PRECAMBRIAN ROCKS

LINEAMENT 

- OUTLINE OF STUDY AREA 

FIGURE 31. Structural lineaments outlining the Belle Fourche Arch, Powder River Basin, Wyoming and Montana

study area and from a review of the geologic literature. 

Sandstone samples from the upper Fox Hills Sandstone 

were obtained from drill cores taken in the uranium 

district west of Oshoto "and from outcrop samples collected during this study from exposures between Moorcroft and Upton (sec. 9, T. 47 N., R. 67 W.).

The mineral phases identified by X-ray diffraction are 

listed in table 18. The absence of pyrite in the outcrop 

was the only difference observed between core and outcrop mineralogies.

Dobbin and Reeside (1930) provided a petrographic 

analysis by C. S. Ross of the Colgate Sandstone

Member of the Fox Hills Sandstone from north of the 

Powder River Basin near the Fort Peck Reservoir in 

Montana (T. 21 N., R. 32 E.). The Colgate Sandstone 

Member is the stratigraphic equivalent of the upper 

part of the Fox Hills Sandstone in Montana This petrographic analysis agrees with the results of the analyses 

by X-ray diffraction and also demonstrates the presence 

of minor amounts of biotite, tourmaline, garnet, zircon, 

and apatite. Ross also identified beidellite, a sodium 

smectite, in the clay matrix.

Stratigraphic descriptions of the Fox Hills Sandstone 

by Robinson and others (1964) pointed out the presence
C42

105°30'

45°00'

REGIONAL AQUIFER-SYSTEM ANALYSIS

105° 15' 105°00' 104°45' 104°30'

44° 45' -

44°30'

44°15'

44°00'

EXPLANATION

WASATCH FORMATION 

(TERTIARY)

LANCE FORMATION 

(UPPER CRETACEOUS)

FOX HILLS SANDSTONE 

(UPPER CRETACEOUS)

PIERRE SHALE 

(CRETACEOUS) AND 

OLDER STRATA

Hi

15 KILOMETERS

Geology from Hodson and others, 1973. 

FIGURE 32. Geologic map of the Powder River Basin study area, Wyoming.

of ferric oxyhydroxides as disseminated grains and 

ferruginous sandstone nodules, numerous calcareous 

beds and concretions, and siderite concretions. They 

concluded that although the lower part of the Lance 

contains an abundance of thin coal seams and carbonaceous shales, the sandstones of the Lance do not 

differ appreciably from those of the Fox Hills. Robinson 

and others (1964) also described selenite in the shales of 

the upper part of the Lance Formation.

HYDROLOGY

Based on substantial head differences between potentiometric surface maps constructed from data on wells 

screened in different geologic units, Anna and Croft 

(1977) concluded that three hydrologically distinct 

aquifer systems are present in the sediments overlying 

the Pierre Shale. The lowermost aquifer system, consisting of the upper part of the Fox Hills Sandstone and
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C43

Fraction Mineral

Micnxline 

Muscovite 

Calcitc 

Kaolinite 

Rutile 

Idealized 

chemical formula

SiO, 

NaAL.SijO,, 

KAlSi,OH 

KAl (Si f010(OH), 

CaCO , 

CaM^CO,)., 

Al,,Si. )O-(OH)4

Mg,Si4 0I()(OH)K 

FcS., 

TiO,

Idealized 

Mineral chemical formula

Fer ie oxvhvdroxides Fc(OH) r FeOOH, Fe.,Os 

Sit rite FeCO,

Be ellite - Na() .Al/Si, _A1() ,)O1()(OH).,

Bi, ,- - - K(Mg,Fc)^AlSip ]()(OH),

Gi n-t 

Apaiie - - Ca.(POJ, (F,C1,OH)

the lower part of the Lance Formation, is confined from 

below by the Pierre Shale and from above by the siltstones and claystones of the upper part of the Lance 

Formation. The two overlying Tertiary aquifers of the 

Fort Union and Wasatch Formations are therefore effectively isolated from the Lance-Fox Hills aquifer. The 

presence of interbedded siltstones and claystones in the 

lower part of the Lance decreases the vertical permeability in the aquifer, impeding vertical mixing between the 

Lance Formation and Fox Hills Sandstone.

A regional potentiometric surface map (D. H. Lobmeyer, U.S. Geological Survey, written commun., 1981), 

presented in figure 33, indicates the inferred directions 

of regional ground-water movement. Northward thinning of the aquifer sandstones in the southern part of 

the study area, in combination with enhanced vertical 

permeabilities along the Belle Fourche Arch (Slack, 

1981), suggests upward ground-water movement and 

subsequent discharge into the Belle Fourche River. An 

undetermined fraction of this flow probably continues 

northwest of the Belle Fourche River and out of the 

study area.

Recharge to the Lance-Fox Hills aquifer enters the 

outcrop area of these formations in the central and 

northern sections of the study area. Recharged water 

flows west toward the axis of the basin, then north into 

Montana, as shown by the northern flow line in figure 

33. Pump tests on the NGP-RASA (Northern Great 

Plains Regional Aquifer-System Analyses) test well in 

the northeastern Powder River Basin (sec. 11, T. 55 N., 

R. 77 W.) yield hydraulic conductivities of 3.5X10" 7 m/s 

for the Fox Hills Sandstone and 1.6X10"6 m/s for the 

Lance Formation, (D. H. Lobmeyer, U.S. Geological 

Survey, written commun., 1981). Assuming a constant 

effective porosity of 0.10 to 0.15, and an average hydraulic gradient of 2.2X10"3 , Darcy's Law may be used 

to calculate average flow rates of 0.16 to 1.08 m/yr for

the ground water in the Lance-Fox Hills aquifer in the 

study area.

DATA PRESENTATION AND INTERPRETATION

MAJOR AND MINOR ELEMENTS

Chemical and isotopic analyses of 21 ground-water 

samples were performed to characterize the Lance-Fox 

Hills aquifer system in the northeastern Powder River 

Basin. Major element determinations are presented in 

table 20; minor and trace element determinations are 

presented in table 21. The locations of the 18 wells 

sampled during this study, PI through PI8, are presented in figure 29, with three additional wells, P20, 

P30, and P41, sampled by Hodson (1971) and Chatham 

and others (1981). Drillers' logs and well depths were used to locate wells open to the upper part of the Fox Hills 

Sandstone to minimize the effects of vertical variations 

in the lithology of this lenticular system. Samples PI, 

P8, and PI2 represent confined ground waters of the 

lower part of the Lance Formation, and P13, P14, and 

P18 were sampled from wells located on the Lance outcrop. Samples from the lower part of the Lance were collected to characterize the part of the aquifer that lies 

above the Lance-Fox Hills contact and to evaluate the 

possibility of recharge to the Fox Hills occurring as 

downward leakage through the Lance in the area of the 

Lance outcrop. Additional information on the composition and geochemical behavior of ground waters from 

the lower part of the Lance Formation was derived from 

Chatham and others (1981). Discussion of ground-water 

chemistry is generally restricted to the upper part of the 

Fox Hills Sandstone because vertical variations in the 

chemistry of the Lance-Fox Hills aquifer system prevent consideration of this unit as a whole (Chatham and
C44 REGIONAL AQUIFER-SYSTEM ANALYSIS

105° 30 1 105° 15'

45° OO'

105° OO' 104°45' 104° 30'

44° 45' -

44°30' -

44°15' -

44°00' -

EXPLANATION

FOX HILLS SANDSTONE 

(UPPER CRETACEOUS)

FOX HILLS WELLS

POTENTIOMETRIC 

CONTOURS Shows 

altitude at which water 

level would have stood 

in tightly cased wells, 

1980 (Lobmeyer.D., U.S 

Geological Survey 

written communication, 

1981). Contour interval 

200 feet.

INFERRED FLOW PATHS

National Geodetic Vertical Datum of 1929

Geology from Hodson and others. 1973.

FIGURE 33. Potentiometric contours and inferred flow paths for the Lance-Fox Hills aquifer, Powder River 

Basin study area, Wyoming [1 foot (in contour interval)=0.305 meter].

others, 1981). Locations of wells sampled in the upper 

part of the Fox Hills are shown in figure 34.

The dissolved solids distribution of waters in the upper Fox Hills (fig. 35) reflects the hydrology of the aquifer system. As expected, dissolved solids concentrations 

increase downflow from initially low values in the northern recharge area A plume of intermediate dissolved 

solids content can be attributed to the mixing of recharge waters with older, northwest-flowing waters; 

waters high in dissolved solids identify the southern discharge area.

Two samples, P9 and PI 7, are representative of 

deeper basin waters and have intermediate dissolved 

solids concentrations. Similarities between waters recharging this area (Hodson and others, 1973) and recharge waters within the study area indicate that the
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C45

TXBI.F 20. Concentrations of dissolved ma/or elements in waters from the Lance-Fox Hill\ aquifer. Powder River Ba\m, Wyoming

n millmi.iins |x-i Inn unless uilu-rwisi sptiified. C. di-i^nes Cikms, jiS/un. IIIKIOMI-IIU-MS |x i icTiiiinrl. i Ondii.ur. come nii.iiinns Ixhiu ilcKHion limn loi ih.n ,,.,KMS|

Calcium, Magnesium, 

Sample Ca Mg Na

270 

100 

till 

Hi

140

24

24(1 

22( 1 

17 

liOd 

21 X) 

'(70 

70

170

Potassium, 

K

Na+ K

1 4

1 ! 

1 i

14

n 

i a

2 (1 ~i a

1 2

i i

i 2 

i '( 

7 4 

40 

r> 0 

'i H 

I,

Chloride, Sulfate.

ci so4

h7 120 

IB I'll! 

2( Till 

1 ! 7"i 

4 B 4i! 

2 4 (>li

1 'I 2N) 

21 '17 

1 1 1,000 

24 14(1 

2 2MI 

2. Ml 

1 1 ! 

 1 (Till 

14 74<) 

 «) 1 I)

2 no

'1 hid

Specific 

Bicarbonate, Carbonate. Temperature. conductance, 

HCO ( CO s pH T SpC (M S/cm 

( C) at 25 C)

Yi2 'i 

"ill i II 

(i2l> 12

tit i < 1

7'MI li 

4d4 < 1 

lill 111 

2dil 1 i 

4d! <1

"idO <1

V)7 20 

CM II

I '12 n B (til 

1 'II I'l h 1 Oil 

177 11! (I I. l"i( 

l"i 110 "i !l 

Hi 110 4W 

71 'Id ")»( 

"ill 'I (i 14( 

'« 11 li 721 

hi- 27 0 l.4d( 

71 '12 1 li'K 

'»2 2 ! (I 1 2IK 

21 in 7(« 

VI 10 (i 7'i( 

h! 10 2 ill

Oi 'III 1 .(!.'(

12 '1 "i 1 4H( 

II! ia 'I 1 Bl( 

2"i 10 4 d'K 

1 7 10 0 l.lihl 

i H") II (i 1,411

deep basin samples should be qualitatively similar to 

waters that eventually evolve from northern study area 

waters.

Stiff diagrams, plotted in figure 36, show that recharge waters contain variable amounts of sodium, 

calcium, and magnesium with bicarbonate, changing in 

cationic concentrations to dominantly sodium types 

within a short distance of the outcrop. Samples collected downflow from the recharge area also contain increasing concentrations of bicarbonate, sulfate, and 

chloride. The deep basin samples, P9 and PI7, are also 

predominantly sodium bicarbonate, but contain slightly 

more bicarbonate and chloride, as well as less sulfate, 

than samples collected to the north. The high dissolved 

solids concentrations in waters from the discharge area 

are apparently caused by additions of sulfate balanced 

by sodium, with waters there containing sodium, bicarbonate, and sulfate as their major components. In the 

area where recharge waters are thought to be mixing 

with more mature waters, an intermediate composition, 

predominantly sulfate water, is found. This composition 

has intermediate bicarbonate and low chloride with 

sodium as the predominant cation, followed by slightly 

lesser amounts of calcium and magnesium.

Ion-exchange reactions appear to play an important 

role in the evolution of ground waters in the Fox Hills, 

as shown by the uniform variation in log molar 

([Ca]+[Mg])/[Na]2 ratios, an operational parameter, 

plotted in figure 37. Only samples taken from the outcrop area have more than several milligrams per liter

calcium and magnesium, indicating that the divalent 

cations are exchanged for monovalent sodium ions present on clay surfaces (table 1, reactions 28 and 29) soon 

after recharge water enters the aquifer. Mixed layer 

smectite clays, such as the beidellite found in the Fox 

Hills Sandstone, are known for their high cation exchange capacities. Grim (1968) reported CEC values for 

smectites on the order of 80 to 150 meq/100 g, which 

are several times greater than the CEC's for illite 

(10-40 meq/100 g) and kaolinite (3-15 meq/100 g). The 

higher CEC of smectite, compared to that of illite and 

kaolinite, may explain the relatively rapid decrease in 

calcium and magnesium concentrations in the Fox Hills 

Sandstone, in contrast to the more gradual decreases 

observed in the Judith Basin where the latter two clays 

predominate.

The plot of log molar divalent/monovalent cation 

ratios (fig. 37) accurately describes the pattern of recharge and ground-water movement along the flow path 

in the northern half of the study area, but interpretation 

of the ratios is somewhat more difficult in the southern 

part. Ambiguity in the contour pattern is most likely 

the result of conflicting effects, because of the mixing in 

the outcrop area of older, northeastward moving waters 

having low log ([Ca]+[Mg])/[Na]2 ratios with younger 

recharge waters having high log ([Ca]+[Mg])/[Na]2 

ratios. In any event, the log molar divalent to monovalent ratios attain significantly larger values north of 

the Belle Fourche River than to the south and indicate 

the greater significance of recharge in influencing the
C46 REGIONAL AQUIFER-SYSTEM ANALYSIS

Sample Alumi 

A

 1

II 1

'': < 

4 1

JO 1 ...... ..

 fir 

H J 

num. NH4 Arser

(mg/L) As

">() <

27 <

1 4 <

 J» <

ic, Bari 

B

<

iim, Boro 

> B

L

1 "i l

Q 

 f

1J

n, Br 

(mg/

<

< 

<

ide, 

Cadm 

L) Cc

<

< 

<

um, Chrom 

Cr

<

<

< 

<

Dissolved 

oxygen, Fluoride. 

um. Copper, On F I 

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

I >'.><

< 1 2 i

< 1 J 4 < 1 < 

'I 1 < 01 JO

1 H 

 If "> 

< 1 1

lodid 

L) Fe

1 1 KM

i i .HI i

i i'Ji 

i

\ *i

1UI 

"ilK

11

Lead. 

Pb

< 1

\

chemistry and hydrology of the northern area.

Chloride concentrations, plotted in figure 38, increase 

along the flow path from the recharge area westward 

and are even higher in the deep basin samples. As in the 

Kootenai, increases in chloride concentration are 

thought to result from minor amounts of leakage 

through shale confining beds. A water sample taken 

from the underlying Parkman Sandstone Member of the 

Mesaverde Formation approximately 45 km southwest 

of Gillette shows the presence of 3,400 mg/L chloride 

(Wyoming Water Planning Program, 1972). To account 

for observed increases in chloride of roughly 25 mg/L, 

an upward leakage of less than 1 percent of this type 

water from the Pierre Shale would be required. Chloride 

may also be leached from shale layers within the aquifer, leading to similar changes in chemistry and making 

a determination of the exact source impossible with the 

data available.

In the discharge area, where chloride concentrations 

should be similar to those measured in the deep basin 

samples because of the similar origins of these two 

waters, the values are relatively low. Low to intermediate chloride concentrations in this area are another indication that, although water is being discharged to the 

Belle Fourche River, recharge may be entering the 

aquifer outcrop and diluting the northward moving 

water.

Sulfate concentrations, plotted in figure 39, behave 

differently in the northern and southern portions of the 

study area. Along the northern flow path, sulfate concentrations increase, then decrease slightly. Comparison with deeper basin samples suggests that sulfate 

will continue to decrease in concentration as the water 

matures. In the recharge area, sulfate is probably derived from the dissolution of gypsum (reaction 20) or the 

oxidation of pyrite (reactions 9 and 10). Gypsum is commonly present in alkali soils from evaporation of soil 

water in arid climates, but, based on mineralogic and 

stratigraphic descriptions, appears to be absent in the 

consolidated sediments of the Fox Hills Sandstone. 

Pyrite may be oxidized only in the presence of O2 (reaction 9) or ferric iron (reaction 10), but these oxidants are 

essentially absent after the appearance of H2S in the 

aquifer. Therefore, neither gypsum dissolution nor 

pyrite oxidation can adequately explain increases in 

sulfate concentration after the appearance of H2S. 

However, high sulfate waters, derived from gypsum 

dissolution or pyrite oxidation, are present in the overlying Lance Formation (up to 710 mg/L, Chatham and 

others, 1981). It seems likely, then, that increases in 

sulfate concentration in the Fox Hills Sandstone probably result from downward movement of high sulfate 

waters from the overlying, hydrologically continuous 

Lance Formation.

Subsequent decreases in sulfate concentration appear 

to be the result of bacterially mediated sulfate reduction 

(reaction 6). Sulfate-reducing bacteria (Desulfovibrio 

sp.) are ubiquitous in ground waters of the northern 

Powder River Basin in Montana (R. W. Lee, U.S. Geological Survey, oral commun., 1981) and should be expected in the Lance-Fox Hills aquifer of Wyoming. In 

the absence of sulfur isotopic compositions, however,
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C47

water*, from the Lance-Fox Hill\ aquifer. Powder River Ha.\m, Wyoming

H,S. hviliimi n Milfuli . <iiHlii. 1 iiMiiiiiiiiii.iiiinislKli 1 » ili u i nun liiiui foi ili.n ,iii.iKsi-,|

Organic Sulfide, 

Nitrate. Nitrite, carbon. Phosphate. Silica, S 

Lithium, Manganese, Molybdenum. NO^ NO2 organic C PO4 Selenium, SiO0 Strontium, (mg/L as Uranium, Vanadium, /inc. 

Li Mn Mo (mg/L) (mg/L) (mg/L) (mg/L) Se (mg/L) Sr H2S) U V Zn

II

U

2(

i.

i,

^

i

->

 >.

IK

i

1

11

(

'(

\

2

(

(

HI

121

IK

'!(

1 11

-1

!

4' II

Jl(

1(

_>K

1

"i!(

 

24(

171

<

<

<

<

<

<

<

<

<C

<

<

<

<

< 1

< 1

< 1

< 1

< 1

1) 1 i «>

1 i <

(14 <

n <

")« <

o<i <

'i > <

H <

O'l <

1'i <

II') <

O'l <

Ti <

7 (1

j 1

1 71 0 IB <

i 'i r> <

10 r> <

14

'i 7

") 1 <

IK <

H 4 <

14

(i <» <

n

1! 0

'2 K

1 1

7 i

'i'l Id h (I <

18 IS H (.

1 i < 01 <l 4

4 i)

b 

2 4

>(> <

I'i <

II <

11

)1 <

Ml <

11 <

il <

18 <

l! <

)". <

Id <

11 1 Mi <

10 <

'I '2 1

10 7

11 7

in '.MO

l"i l.Vl

'I 7 1,111

11 J.10

11 n

1'2 11

<l "l J 1(1

\'2 (i

10 7

K, 77

1 1 '10

H 'I di

1 10 "i.HK

17 I"!

l"i i.KH

I) 2 0 17 1

"i 2\

2 12

1 17

1 0-

i 1 (]

1 '1 1

i Hi

J 11,

2 21

2 07

1 01

", 01

< 1 \ 2

1 2 ~i

1 7 'i

J7I

i ;i

_>t

1 K

i ji

mi

i 'i

Mi'

r,i

2 1-' 1 <

1 7(1 1 'i

HO

I 2 --- - < 1 0 'I 1

) !( < 1 (1 J II 1

the small decrease in sulfate (10 mg/L) along the northern flow path cannot be unequivocally attributed to 

sulfate reduction. Similarly, mass transfer in the 

aqueous phase is not sufficiently well defined along the 

deep basin flow path to infer sulfate reduction. Therefore, the process of sulfate reduction must be considered 

as likely to occur only in the Lance-Fox Hills aquifer; it 

cannot be considered as an incontrovertible source of 

CO2 in the system.

The highest sulfate concentrations in the study area 

are found along the Belle Fourche River, where recharge 

waters are thought to be mixing with older, more 

mature waters in the aquifer outcrop. The introduction 

of variable amounts of oxygenated recharge water into 

a reduced ground-water flow system will undoubtedly 

result in fluctuations of the redox interface (Eh=OmV), 

and the oxidation of sulfide minerals, generating 

sulfate. Additional sulfate may be derived from the 

dissolution of gypsum in the soil zone, although this 

source would not readily explain high sulfate concentrations present only in the southern portions of the study 

area.

Large increases in pH are evident along the northern 

flow path from the recharge area basinward as shown in 

figure 40. As previously discussed, increases in pH are 

commonly observed in ground-water flow systems as a 

result of dissolution or hydrolysis of aquifer mineral 

phases. Comparison with the deep basin samples indicates that slight decreases in pH may occur as geochemical evolution continues in the aquifer. In the 

discharge area, pH's are at low to intermediate values,

similar to those measured in waters from the northern 

recharge area. Low to intermediate pH's in wellbuffered, high-bicarbonate waters might support the 

hypothesis that high sulfate concentrations in these 

waters are due to pyrite oxidation, although such a conclusion is tenuous.

Total inorganic carbon concentrations were calculated 

by summing the molalities of all carbonate species computed by WATEQ2. Calculated CT values, plotted in 

figure 41, indicate that increases in bicarbonate and carbonate concentrations are representative of increases in 

CT along the flow path. Comparison with the deep basin 

samples indicates that this trend should continue with 

further ground-water evolution. Concentrations of CT 

are highly variable in the discharge area, ranging from 

7.9 mg/L at P10 to 15.2 mg/L at P15, suggesting largescale local variations in ground-water chemistry.

Increases in the concentration of total inorganic carbon along the flow path in the Fox Hills Sandstone 

reflect an influx of carbon to the aqueous system. As in 

the Judith Basin, increases in CT concentration may be 

the result of oxidation of organic carbon by O2, NO~, or 

SO~ 2 reduction processes, coalification of lignite, 

methanogenesis, migration of CO2 or CH4 into the aquifer, or the dissolution of carbonate minerals. Although 

O2 and NO~ reduction processes undoubtedly produce 

CO2 in the outcrop area, these compounds are rapidly 

consumed and unavailable for reaction downdip. The 

presence of H2S in all but several outcrop wells suggests 

that sulfate reduction may be an active process, but a 

lack of appropriate data and the existence of open
C48

105°30'

45°00'

REGIONAL AQUIFER-SYSTEM ANALYSIS

105° 15' 105°OO' 104°45' 104° 30'

44°45' -

44°30' -

44°15' -

FOX HILLS SANDSTONE 

(UPPER CRETACEOUS)

44° 00' -

Geology from Hodson and others. 1973.

FIGURE 34. Index map of Powder River Basin study area, Wyoming, showing the locations of wells sampled

in the Fox Hills Sandstone.

system conditions with respect to SO^2 make sulfate 

reduction impossible to substantiate. Considering the 

occurrence of lignite beds at the Lance-Fox Hills contact, coalification may provide significant influxes of 

CO2. Migration of CO2 into the aquifer is also possible, 

although oxidation of migrating CH4 as a CO2 source is 

unlikely because CH4 would be expected to be relatively

stable in anoxic waters of the Fox Hills Sandstone. Acetate fermentation (reaction 7) may also be an active 

process and would be evident by a 70 per mil carbon isotope fractionation accompanying this reaction (Bottinga, 1969).

Acidity produced by the generation of CO2 would 

allow for the dissolution of saturated mineral phases.
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C49

105°30' 105° 15' 105°00' 104°45' 104° SC00'

44° 45' -

44°30' -

44°15' -

EXPLANATION

FOX HILLS SANDSTONE 

(UPPER CRETACEOUS)

FOX HILLS WELLS

LINE OF EQUAL 

DISSOLVED SOLIDS 

CONCENTRATION 

Interval 500 milligrams 

per liter. Dashed where 

inferred.

44°00' -

Geology from Hodson and others, 1973.

FIGURE 35. Dissolved solids concentration in waters from the Fox Hills Sandstone, Powder River Basin,

Wyoming.

Dissolution of carbonates would lead to further increases in CT (reactions 21-23); hydrolysis of aluminosilicates (reactions 24-27) would not. A phase diagram 

for the system Ca-Mg-CO2 -H2O is presented in figure 

42. The compositions of waters from the Fox Hills 

Sandstone having magnesium concentrations significantly greater than the detection limit plot primarily in

the calcite stability field, near the calcite-dolomite 

stability boundary. Therefore, dolomite should be 

unstable in these waters, possibly dissolving incongruently to form calcite.

It is interesting to note that the compositions of two 

Fox Hills Sandstone waters reported by Chatham and 

others (1981) clearly fall in the dolomite stability field.
C50

105°30'

45° 00'

REGIONAL AQUIFER-SYSTEM ANALYSIS 

105° 15' 105°00' 104°45' 1040 30'

44° 45' -

44° 30' -

44° 15' -

EXPLANATION

FOX HILLS SANDSTONE 

(UPPER CRETACEOUS)

FOX HILLS WELLS 

STIFF DIAGRAMS 

an.ons HCO +C0

10 5 0 5 10 

milliequivalents per liter

44° 00' -

Geology from Hodson and others, 1973.

FIGURE 36. Stiff diagrams of major element concentrations in waters from the Fox Hills Sandstone, Powder

River Basin, Wyoming.

These waters are thought to represent mixing between 

recent recharge and older, high sulfate, high-bicarbonate 

waters. The resulting mixture has a (a +2)/(a +2) ratio 

significantly greater than unity, implying that clolomitization of calcite may be occurring in these waters on a 

limited scale (Hanshaw and others, 1971; Folk and 

Land, 1975). 

Silica concentrations remain relatively constant along

the flow path, averaging 10.7 mg/L, while aluminum 

concentrations vary irregularly from 0.010 to 

0.140 mg/L. Reasons for these fluctuations in aluminum 

concentration are uncertain, although contamination by 

wind-blown dust, which was a problem at the well with 

the highest aluminum concentrations, may explain 

anomalously high concentrations of aluminum, iron, 

manganese, and silica in such samples.
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C51

105°30' 105° 15'

45° GO00' 104°45' 104° 30'

44° 45'

44°30'

44° 15'

44°00'

EXPLANATION

FOX HILLS SANDSTONE 

(UPPER CRETACEOUS)

FOX HILLS WELLS

LINE OF EQUAL LOG 

MOLAR CALCIUM PLUS 

MAGNESIUM TOSODIUM 

SQUARED RATIO

(LOG(Ca+Mg)/Na 2 ) ) - 

Interval variabje. 

Dashed where inferred.

Geology from Hodson and others, 1973.

FIGURE 37. Log molar ratios of calcium plus magnesium to sodium concentrations in waters from the Fox

Hills Sandstone, Powder River Basin, Wyoming.

Phase diagrams were calculated using the thermodynamic data of Helgeson (1969) and the kaolinite stability of Bassett and others (1979), to determine which 

alumino-silicates are stable in the waters sampled. A log 

(aH4Si0<j) log (aN +)/(aH +) activity diagram with compositions of waters from the Lance-Fox Hills aquifer superposed (fig. 43) best demonstrates water-mineral relationships in this system. Kaolinite is stable in low pH, low

sodium outcrop waters, whereas Na-montmorillonite, a 

pure end member, is stable in all other waters. Chlorite, 

which does not plot on this diagram, is unstable in all 

waters sampled, implying that aluminum and silica concentrations are controlled by kaolinite and a smectite 

similar in stability and composition to a Namontmorillonite. 

Strontium decreases in concentration basinward
C52

105°30'

45° 00'

REGIONAL AQUIFER-SYSTEM ANALYSIS

105° 15' 105°OO' 104°45' 104° 30'

44° 45' -

44°30' -

44°15' -

EXPLANATION

FOX HILLS SANDSTONE

s*:*:*a (UPPER CRETACEOUS) 

FOX HILLS WELLS

LINE OF EQUAL CHLORIDE 

CONCENTRATION 

Interval 10 milligrams 

per' liter. Dashed where 

inferred.

10

15 KILOMETERS

44°00' -

Geology from Hodson and others, 1973.

FIGURE 38. Chloride concentrations in waters from the Fox Hills Sandstone, Powder River Basin, Wyoming.

along the flow path (fig. 44). Variations in strontium 

concentration are probably related to adsorption on clay 

surfaces in a manner similar to the other alkaline earths 

(Wahlberg and others, 1965a; 1965b), but may also reflect equilibrium with respect to carbonates or sulfates 

present in the system. Celestite, strontianite, or a strontium calcite therefore seem to be the most likely mineral 

phases to control strontium concentrations.

Oxidation potentials in waters of the Fox Hills Sandstone were calculated using the Nernst equation and the 

O2 -H2O (reaction 12) or SO72 -H2S (reaction 16) redox 

couples, as appropriate. The Eh values calculated from 

these couples are listed in table 22 and plotted in figure 

45. Oxidation potentials vary from positive values in 

the recharge area to increasingly negative values along 

the flow path and in the deeper basin samples. Along
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C53

105°30' 105° 15' 105° OG45° 00'

104° 30'

44° 45' -

44°30' -

44°15' -

EXPLANATION

FOX HILLS SANDSTONE 

(UPPER CRETACEOUS)

FOX HILLS WELLS

100 LINE OF EQUAL SULFATE 

CONCENTRATION 

Interval 200 milligrams 

per liter. Dashed where 

inferred.

44°00' -

Geology from Hodson and others. 1973. 

FIGURE 39. Sulfate concentrations in waters from the Fox Hills Sandstone, Powder River Basin, Wyoming.

the flow path, O2 and NO~ are consumed (fig. 45) followed by the appearance of H2S immediately downgradient from the aquifer outcrop, suggesting an 

abundance of reduced phases such as pyrite or organic 

carbon in the aquifer sandstones. Dissolved organic carbon determinations (fig. 45) indicate much higher 

concentrations in the Lance-Fox Hills aquifer than in 

the Kootenai Formation, ranging in a random fashion

from 1.1 to 18 mg/L, with a mean of 7.8 mg/L. Waters 

sampled in the discharge area exhibit highly variable 

Eh, ranging from positive to negative values, supporting the suspected influx of oxygenated recharge water 

in this area.

Chatham and others (1981) have observed similar 

trends in major element concentrations, Eh and pH in 

their study of the Lance Formation. This similarity
C54

105°30'

REGIONAL AQUIFER-SYSTEM ANALYSIS

105° 15' 105°00' 104°45' 104°30'

45° 00'

44° 45' -

44°30' -

44°15' -

44° 00' -

 8.5 FOX HILLS WELLS 

pH values

Geology from Hodson and others, 1973. 

FIGURE 40. pH values of waters from the Fox Hills Sandstone, Powder River Basin, Wyoming.

suggests that similar geochemical mechanisms control 

ground water in both the upper part of the Fox Hills 

and lower part of the Lance Formations, as would be expected from their similar mineralogy and hydrogeology. 

Iron and manganese are present in low concentrations 

in both oxidized and reduced waters (fig. 46), with maxima occurring near the redox interface as for several 

flow systems described by Champ and others (1979) and 

Langmuir and Whittemore (1971). Such behavior suggests controls on the aqueous concentrations of these 

metals similar to those observed in the Judith Basin, or 

control by oxyhydroxides in oxidized waters and by carbonates or sulfides in reduced waters.

DISSOLVED GASES

As in the Judith Basin, considerable difficulty was encountered in attempting to collect uncontaminated gas 

samples. Bubbles adhering to the walls of the sampling
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C55

105°30' 105° 15'

45° 00'

105° 00' 104°45' 104° 30'

44° 45' -

44°30' -

44° 15' -

44° 00' -

EXPLANATION

FOX HILLS SANDSTONE 

(UPPER CRETACEOUS)

FOX HILLS WELLS

LINE OF EQUAL TOTAL 

INORGANIC CARBON 

CONCENTRATION 

Interval 5 millimoles 

per liter. Dashed where 

inferred.

Geology from Hodson and others. 1973.

FIGURE 41. Total inorganic carbon concentrations in waters from the Fox Hills Sandstone, Powder River

Basin, Wyoming.

flask were a constant problem when sampling effervescent waters, and some samples were contaminated by 

the leaking of air into the evacuated sidearm. Analyses 

with anomalous N2 , Ar, and O2 concentrations were 

therefore discarded, leaving those listed in table 23.

Dissolved-oxygen concentrations agree with those 

measured in the field to within 0.3 mg/L or less for three 

of the wells sampled, but the analyses for sample PI3

differ by 2.0 mg/L, and the analyses for sample PI5 differ by 0.6 mg/L. Field notes indicate that the pumps 

mounted on these two wells were at times drawing air 

into the sample water, which would account for the 

widely disparate analyses. Other dissolved gases measured at these wells do not seem anomalous in comparison with remaining samples from the Lance-Fox 

Hills aquifer, suggesting that the oxygen concentration
C56 REGIONAL AQUIFER-SYSTEM ANALYSIS

9 -2

5 -3

MAGNESITE DOLOMITE CALCITE

t 

LOG OF CALCIUM ACTIVITY MINUS LOG OF MAGNESIUM ACTIVITY (LOG (aCa+2,a +2))

FIGURE 42. Stability diagram for the system Ca-Mg-CO2 -H2O 

at 25°C, with compositions of waters from the Lance-Fox 

Hills aquifer, Powder River Basin, Wyoming, superposed.

LOG OF SILICIC ACID ACTIVITY ( LOG aH S|O0 )

FIGURE 43. Stability diagram for the system Na-SiO2 - 

A12O3 -H2O at 25°C, with compositions of waters from 

the Lance-Fox Hills aquifer, Powder River Basin, Wyoming, superposed. Thermodynamic data for kaolinite 

from Bassett and others (1979). Remaining thermodynamic data from Helgeson (1969).

measured using the gas sampler may reflect aquifer conditions more accurately than that measured using the 

dissolved-oxygen probe and meter (table 2).

Comparison of measured CO2 partial pressures with 

Pco values computed from field-measured pH and 

alkalinity values using WATEQ2 show relatively good 

agreement, as shown in table 24. Differences between 

measured and calculated values lie within 0.1 pH units, 

which is the accuracy reported by Langmuir (1971) for 

field pH measurements in ground waters from noncarbonate rock terrains. Improved agreement between rev re P2

0 10 20 30 40- 50 60 

DISTANCE FROM OUTCROP ALONG NORTHERN FLOW PATH, IN KILOMETERS

FIGURE 44. Strontium concentrations in waters from the northern flow path of the Fox Hills Sandstone, Powder River Basin, 

Wyoming.

r\Hll 22. O\idatwn potentials calculated fur waters from the Lame-Fox Hill\

aquifer. Powder River Bcum, Wyoming, Ming O,,-H,O or SO4~~-H,S redox

Sample

PI ____________

P2 ______ 

PJ.

pf , 

P7

P8 ____________

PIO ___________

i) i i

P12 

pn 

PI 4 

i>r, ___________

PI i^

PI 7 

P1H 

Calculated Eh 

(mV)

307

317

302

809 

203

-219 

801 

302

314

')f tQ

-322 

-321 

- 201 

788 

769 

808 

321

- 187

Redox couple

SO---H.S 

SO4-*-H,S 

SO; J H,S

O,-H,cT

s64 J-H,S

SO;-'-H,",S 

O.,-H,0~ 

SO } J-H,S 

SO; -'-R,S 

SO; -'-H JS 

so; J-H> 

so4 J-H",S 

SO;--H") S

0,-H,0"

O^-H",O

0>H~,O SO-'-H,S

SO; 2 H.>

suits obtained by these two methods over that observed 

for samples from the Judith Basin is probably due to 

the higher buffer capacity of the Powder River Basin 

samples. Average buffer capacity of the pertinent 

Judith Basin samples is 0.63 meq/L per pH unit and of 

the powder River Basin samples is 1.45 meq/L per pH 

unt. With a higher buffer capacity, similar amounts of 

CO., could escape (or enter) the solution with less of a 

change in pH.

Nitrogen partial pressures for most samples are 

higher than would be expected, assuming atmospheric 

equilibration with recharge waters (PN atm =0-69), suggesting that Lance-Fox Hills ground waters become enriched in N2 after isolation from the atmosphere. Nitrate
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C57

£ <

t en

z g

d

'7P6 P2 P3

0 10 20 30 40 50 60

DISTANCE FROM OUTCROP ALONG NORTHERN FLOW PATH, IN KILOMETERS

FIGURE 45. Calculated oxidation potentials and dissolved organic 

carbon and nitrate concentrations in waters from the northern flow 

path of the Fox Hills Sandstone, Powder River Basin, Wyoming.

may be reduced to N2 by denitrifying bacteria (reaction 

2), thereby increasing N2 while simultaneously decreasing NO~ concentrations, as observed. The data are insufficient to test this hypothesis, however.

The temperature of recharge waters may be calculated from the concentration of argon present in the 

ground water (Mazor, 1972). Based on an atmospheric 

abundance of 0.934 percent, an atmospheric pressure of 

0.885 atm at 1,220 m elevation, and measured solubility 

coefficients for argon (Douglas, 1964), the temperature 

dependence of argon solubility has been calculated and 

compared to measured argon concentrations. Excluding 

sample PI 7, which appears to have lost argon and 

nitrogen as a result of displacement by methane (D. W. 

Fisher, U.S. Geological Survey, written commun., 

1980), the calculations yield recharge temperatures 

ranging between 0°C and 6°C, with a mean of 2.75°C. 

This temperature is significantly less than the annual 

average of approximately 7°C, and suggests that spring 

snowmelt provides a higher proportion of recharge to 

the Fox Hills than warmer spring and early summer 

storms and that paleo-recharge has been derived from 

similar events. This interpretation is consistent with 

the observation that precipitation from spring and summer storms results in runoff and increased streamflow 

rather than infiltration into bedrock aquifers (Whit-

>7P6 P2 P3

go 10 20 30 40 50 60

DISTANCE FROM OUTCROP ALONG NORTHERN FLOW PATH, IN KILOMETERS

FIGURE 46. Iron and manganese concentrations in waters from the 

northern flow path of the Fox Hills Sandstone, Powder River 

Basin, Wyoming.

comb, 1965). Another noble gas, helium, is absent in detectable quantities in all but one sample, PI7. This is 

presumably the most mature ground water sampled, so 

it seems likely that helium has accumulated through the 

radioactive decay of alpha emitting radionuclides, as in 

the Judith Basin, and has reached detectable levels in 

only the oldest ground waters.

Methane is present in trace quantities in samples P13 

and PI 5, where oxidation potentials appear too high to 

allow methanogenic bacteria to function. The methane 

in these samples probably originated in more reducing 

environments and subsequently migrated to these 

waters. Methane occurs in greater concentrations in the 

two deep basin samples: P9 (not previously discussed 

due to bubbles in the gas sampler) and PI7. Methanogenic bacteria generate CH4 by either acetate fermentation (reaction 7) (Carothers and Kharaka, 1980) or CO2 

reduction (reaction 8) (Nissenbaum and others, 1972; 

Claypool and Kaplan, 1974). Although insufficient data 

are available to determine which of these processes 

dominates in the deeper waters of the Fox Hills Sandstone, low oxidation potentials and the absence of 

sulfate in quantities greater than 1 millimole per liter 

suggest that the methane was produced in situ, rather 

than migrating from even deeper strata Based on
C58 REGIONAL AQUIFER-SYSTEM ANALYSIS

Sample

P4

PI:', 

PI 5

PI 6

pi 7

Oxygen

mg/L

n \

________ .04

1 '<

 3.6

.3

atm

0.009 

.0008 

.024 

.065 

.010

Carbon 

dioxide

mg/L

47 

23 

13 

33

7.2

atm

0.021 

.010 

.006 

.014 

.007

Nitrogen

mg/L

22 

25 

21 

21 

10

atm

0.93 

1.05 

.86 

.85 

.70

Argon

mg/L

0.75 

.83 

.74 

.67 

.38

atm

0.010 

.011 

.010 

.009 

.008

Helium

mg/L

0.000 

.000 

.00(1 

.000 

.001

atm

0.000 

.000 

.000 

.000 

.001

Methane

mg/L

0.00 

.03 

.03 

.00 

<>./

atm

0.000 

.001 

.001 

.000 

.56

T\Hl.f. 24. Comparison of field-measured pH vahm and pH calculated from measured 

P,.() for water .\ample*, from the Lance-Fox HilL aquifer, Powder River Basin,

Sample

P4

P13 

P15 

Plb 

pi 7

PH, 

field measured

7.15

7.59

8.03

7.42

8.18

PH, 

calculated

7.23

7 SO

o AC:

7.47

8.22

Difference

0.08

f\f\

- .05

.04

Sample

>1

'3 

H -

>5 -

J6

>7

>K

»)

> i n

>i) _ >n? -

> ]:5 .

>14

M5 -

>K, _

>17 .

>1 W

Oxygen

(6 180)

20 5

19.4

-19.1

I Q "7

1 O (1

1 Q ')

1 Q X.

1 U '^

19.6

t O Q

19.2

-20.7

1 Q Q

17.5

19.2

J8.:>

19 0

1 Q 7

Hydrogen 

<6D)

i CQ

149

147

- 145

140

143

143

1 "ii

151

147

144

136

114

141

145

143

Carbon 

(6 nC)

10.60

11.05

- 10.10

- 11.15

12.35

12.90

11.40

12.80

- 5.95

n

o R

11.25

- 13.45

i o QC;

10.35

1 ') SO

3.75

13.65

Tritium 

(pCi/L)

<5

<5

<5

<5

6

<5

58

<5

22

<5

<5

<5

50

25

<5

<5

Carbon-14 

(percent 

modern)

>6.2

1.1

40.9

81.4

61.6

16.3

<2.7

42.4

<2.1

3.4

52.7

97.0

26.4

42.9

<1.0

observations of natural systems, Nissenbaum and 

others (1972) suggested that, despite the stability of 

CH4 over CO2 at oxidation potentials less than 250 

millivolts, catalytic methanogenic bacteria are inhibited

by the presence of more efficient sulfate reducers 

(Desulfovibrio sp.). Once sulfate has been nearly consumed, methanogenic bacteria become more efficient in 

utilizing the available nutrients than the sulfate reducers, and the methanogenic bacteria become the 

dominant anaerobes (Claypool and Kaplan, 1974).

ISOTOPIC COMPOSITION

Isotopic compositions of ground waters in the 

Lance-Fox Hills are presented in table 25. Oxygen and 

hydrogen isotopic measurements in ground waters of 

both the Fox Hills Sandstone and the Lance Formation 

correspond to the compositon of North American continental precipitation as reported by Gat (1980) 

(5D=7.95518O+6.03), indicating the presence of only 

meteoric waters in the aquifer (fig. 47). Values for 618O 

range from 20.7 to 17.5, and for 6D, from 158 to 

 136. Comparison with the continental precipitation 

isotopic data of Yurtsever (1975), as reported by Gat 

(1980), suggests that recharge is derived from precipitation events occurring at temperatures of roughly 0°C. 

This temperature is consistent with dissolved argon recharge temperatures and supports the conclusion that 

the major portion of recharge to the Lance-Fox Hills 

aquifer is derived from spring snowmelt or similar 

paleo-events, rather than from warmer late spring and 

summer storms. Differences between Lance and Fox 

Hills 6 18O and 5D values appear insignificant, indicating similar means of recharge for the two formations.

Carbon isotopic compositions remain relatively constant for all but the two deepest samples from the Fox 

Hills Sandstone, ranging along the flow path from 613C 

values of 12.9 and 11.4 in the recharge area to 

values of 11.0, 10.1, and 11.2 downflow. As discussed in the Supplemental Data section, 613C values 

near 15 or higher suggest control of ground-water 

chemistry by carbonate rather than aluminosilicate 

solid phases, assuming the system is closed to exchange 

with a gas phase CO2 reservoir but receiving CO2 inputs;
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C59

-130

O

CO

O

CO

CD 

O

DC 

Q

-140

-150

Q

-160

-21 -20 -19 -18 -17 

OXYGEN-IS/OXYGEN-16 ISOTOPIC RATIO (6180, SMOW)

FIGURE 47. Isotopic composition of North American continental 

precipitation (Gat, 1981) with isotopic compositions of waters from 

the Lance-Fox Hills aquifer, Powder River Basin, Wyoming, 

superposed.

exchange with a CO2 gas phase would result in <5 13C 

values of 23 to 15. The observed carbon isotopic 

compositions therefore imply geochemical control of 

dissolved carbonate chemistry by carbonate mineral 

phases.

The deepest, and presumably the most mature, 

ground waters sampled, P9 and PI 7, exhibit 613C values 

of -6.0 (P9) and -3.8 (PI7), which are much higher 

values than found in any of the other waters sampled 

fromthe Lance-Fox Hills aquifer. The presence of 

significant quantities of methane in these deep basin 

samples strongly suggests enrichment of 13C in dissolved carbonate species as a result of isotopic fractionation 

during the bacterial production of methane. Sackett and 

Chung (1979) have shown that isotopic equilibrium between coexisting CH4 and CO2 does not occur at aquifer

temperatures; therefore, migration of methane would be 

unable to produce these isotopic variations. Similarly, 

oxidation of migrating CH4 to CO2 , which is not known 

to occur in the absence of molecular O2 , would tend to 

lower <5 13C values from the addition of the isotopically 

light CO2 produced.

The similarity in geochemical environments and isotopic evolution between known methanogenic waters 

(Rosenfeld and Silverman, 1959; Oana and Deevey, 

1960; Nissenbaum and others, 1972; Carothers and 

Kharaka, 1980) and the deep basin Fox Hills samples is 

therefore taken as evidence of active methanogenesis in 

the deeper portions of the aquifer. Because mass 

balance information regarding the compositions of 

these waters before the onset of methanogenesis is lacking, a determination of which process actually produces 

the methane CO2 reduction or acetate fermentationis impossible.

Tritium concentrations for the majority of Lance-Fox 

Hills ground waters are below 10 pCi/L, indicating that 

the confined system is not open to recharge by, or mixing with, recently recharged surface waters. Higher tritium concentrations in the outcrop area, ranging from 

22 to 560 pCi/L, indicate the presence of waters less 

than 25 years in age. These younger waters are also evident in the aquifer outcrop near the Belle Fourche 

River, proving the existence of recent recharge, and substantiating the hypothesized mixing of mature waters 

with recent recharge in this area.

GEOCHEMICAL COMPUTER MODELING

SATURATION INDICES

Mineral saturation indices, computed using the 

WATEQ2 aqueous speciation model (Ball and others, 

1980), indicate that solution-mineral equilibria in both 

the Lance Formation and the Fox Hills Sandstone are 

sufficiently similar to allow a discussion of the aquifer 

as a whole. Saturation indices for the Powder River 

Basin samples are presented in table 26 for silicates and 

aluminosilicates, and in table 27 for carbonates, 

sulfates, and phosphates.

Quartz, a major constituent in the Fox Hills Sandstone, is oversaturated; amorphous silica is undersaturated. Chalcedony is undersaturated or near saturation, but petrographic evidence (Dobbin and Reeside, 

1930) suggested that this mineral is not an important 

diagenetic phase. Albite, another important clastic constituent, is undersaturated, but its alteration phase allophane is saturated in most samples, reflecting groundwater equilibration with this feldspar alteration phase 

rather than with unaltered albite. Muscovite, a
C60 REGIONAL AQUIFER-SYSTEM ANALYSIS

TABLE 26. Saturation indices of silicate and aluminosilicate aquifer minerals in waters from the Lance Formation and the Fox Hills Sandstone, Powder River Basin, Wyoming

Sample Quartz Chalcedony

Amorphous 

silica Albite Allophane Muscovite Chlorite Beidellite Kaolinite

Lance Formation

po

P12

Mean and 

standard 

deviation

0.31 ±0 OH 

.46+ 07 

.i5± 08 

66 ± 06 

.58+ .06 

64 ± 06

0 50 ±0.15

- 0 2 1 ± 0 08

.12± 06 

.04 ± 06 

.10± .06

-0 40 ±0.14

-.59± .07

- 48+ .06

-0 55 ±0 14

- 1 13 + 0 72

- 1 36 ± 1 29

- 2 28 ± 1 46 

- 1 56 ± 0 55

- 80+ 1.32

01 + 38

-0.60 + 0 45

- 1 8h±2 12

- 96 + 2 74

- 1 23 + 2 74 

- 1 48 ± 1.07 - 3 99 ± 3 3Q-

- 0 18+1 66

1 14 + 2 14

99 + 2 14

0 32 + 1 22

- Kl ± > 73

2 10 ± 97 

1 '! 1 ± 2 90

(1 Mi + 1 10

Fox Hills Sandstone

P5 -

P6

P17

Mean and 

standard 

deviation

TABLE 27. Saturation

Sample Calcite

0 28 + 0.08 

36 ± .07 

71 + 05 

63+ 06 

47 ± 08 

52 + 07 

25 + 07 

46 ± .08 

30 ± 07 

42 ± 08 

48 ± 08 

24 ± .Ob

0 43 ±0 15

indices of carbonate,

Dolomite

09+ .06

- 03+ 07

-.08± .08

- 19+ 06 

-0 09 ±0 12

- 66 ± .07 

- 34 + 05

- 59 ± 08

- 73 ± 07

- 70 ± 07

-0 60 ±0 12

- 1 22 ± 0 72

- 2 08 ± 1 29

- 2 93 ± 1 38

- 1 65 + 1 38

-.65± .41

- 1.58±0.87

- 15+ 15

- .38± 1 09

-1.56+ ^

- 40 + 97

- 0 68 ± 0 66

- 2 37 + 2 12

- 74 ±2 74

50 + 99

- 1 87 + 1 Ib 

- 1 10 + 0 96

3 74 + 72

- 4 84 + 2 10

sulfate, and phosphate minerals in waters from the Lance Formation and the Fox Hills Sandstone,

Siderite Rhodochrosite Strontianite Gypsum Celestite F-apatite

1 6 + 214

31 + 'i li

'iH + 2 14

-lt.48±0 88

Powder River Basin,

F-CO3 apatite

-0 26 + 1 M 

1 98 ± .34 

1 52 + 2 7'i 

1 % ± 2 56

% ± 2 73

78 ±2 73 

42 ± 56

88 ± 56 

06 ± 78

0 87±0 76

Wyoming

MnHPO4

Lance Formation

Mean and 

standard 

deviation 0 11 ±0.17

- 74 ± 08 

02 ± 07 

.07 ± 07

-024 ±0.32

0 40 + 0.56 

- .54 ±2 08 1 

-.30 ±2.08*

- 11.76± 1.06 

.20+ .06

0.10±0.392

-0.51 ± 1 40

- 14 ± 06

- 33 ± 06 

-0 34±0.143

-0 70 ±0.28 

-.68± 19 

-,74± 28 

-1.17± 10 

-1.13± 60 

-.80± 09

- 0 87 ± 0 22

- 2 92 ±0 21 

-2 15+ .11 

-254+ 19 

- 1 44 ± .06 

- 2 30 ± 20 

- 1 14 ± 05

- 2 08 ± 0 67

- 2 96 ± 0 28

- 1 62+ 10

- 2 28 ± 0.66

- 1 13 + 1 11

- 1 31 +0 34

 ? 15 + 2 13 

2 96 ± 1 01

2. 00 ±2 48 

2 70 ± 0 62

- n± 55

72 ± 23 

0 12 ±0 52

Fox Hills Sandstone

P5 00 + 0*)

Mean and 

standard 

deviation 0 Ib ±0 Ib

Calculated assuming minor 

Calculated without oxidized

- 1 00± 1 13 1

- 21 ± .07 

-.30+ 07 

.02 ± 08 

14± .07

44 ± .07

58+08 

78 ± .06

-009 ±0.53

0.35 ±0.73 

.66 ± .30

- 46 ± 13 

.43 ± .07 

- 11 44 ± .56 

70 ± 39 

43 ± 06

- 11 26+ 39 

-1030± 09 

12 ± 1 06

0 28±0.382

- 0 53 1 67 

- .58 1 40

- 52 06

-.56 .12

-.79 2 76 

-.81 .12 

-.55 2.08 

-1.21 86

-0 67 ±0 12 2

- 0 65 ± 0 28 

- 76+ 28 

- 1 06 ± 09 

- 98 ± 10 

-.81± 10 

-.82± 10 

- 51 + 21 

- 69 ± 10 

- 64 ± 31 

- 58+ 11 

- 48 ± 09 

- 53+ 16

-0 71 +0 18

- 2 85 ± 0 25 

- 2 97 ± 26 

- 1 41 ± 06 

- 1 80 ± 08 

- 1 88 ± 07

-3 12+ 22 

- 57 ± 05 

- 3 1 1 ± 29 

- 1 19+ 06 

- 49 ± 05 

- 4 70 + 1 56

-2 15 ± 1 23

-2 88 + 0.28 

-290± .28 

- 1 74 ± .10

- 2 03 ± .11 

- 1 85 ± 10

-299± .31 

- 1 48+ 11

- 2 22 ± 1 06

- 1.40± 1 07 

- 2 38 ± 1 33

-2.53± 1 .38

- 1 40 + 73

-2 08 + 0 75

2 12 + 2 05 

42 + 2.56 

2 31 + LSI 

41 ± 2.48

- 1 79 + 2 42

1 49 + 2 9(1 

.( 48 ± 1 27

0 65 ± 2 22

- Sh± 72 

8(1 ± 12 

14 + .23

~

- 79 ± 7'i 

- 0 34 ± 0 60

element concentration equals analytical detection limit, 

samples P4, P7, P14, P15, and P16

potassium mica, is near saturation, indicating possible 

ground-water equilibration with this phase. Apparent 

undersaturation may indicate that muscovite, as well as

sericitic alterations of a similar composition, are 

unstable under current geochemical conditions and 

might be alterating to more stable phases such as
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C61

hydromuscovite, illite, or kaolinite. Chlorite is highly 

undersaturated in most of the samples analyzed, indicating that this detrital phase does not control solution chemistry. Beidellite, a sodium smectite, is 

saturated in waters of the Lance-Fox Hills aquifer, as is 

kaolinite, suggesting that equilibrium between kaolinite 

and beidellite may control silica and aluminum concentrations in solution. This conclusion is in agreement 

with petrographic descriptions that describe kaolinite 

and beidellite as authigenic, and quartz, albite, 

muscovite, and chlorite as detrital phases (Dobbin and 

Reeside, 1930).

Saturation indices for carbonate minerals (table 27) 

are remarkably similar to their counterparts in the 

Kootenai Formation. Calcite is saturated or slightly 

oversaturated in waters of the Lance-Fox Hills aquifer, 

suggesting equilibrium with, or possible precipitation 

of, this phase. Apparent oversaturation may be the 

result of cationic substitutions in the calcite lattice or 

may reflect small errors in the field pH measurements. 

Dolomite is saturated in most samples, although those 

from the discharge area are clearly oversaturated with 

respect to dolomite. These oversaturations may indicate 

limited dolomitization of calcite in the outcrop area due 

to mixing of high CT, high pH waters with recharged 

waters containing high concentrations of magnesium. 

Siderite saturation indices indicate saturation with 

respect to this phase in all but oxygenated waters.

Fox Hills rhodochrosite saturation indices cluster 

near a value of 0.67, with significant deviations occurring only in oxygenated waters, where manganese oxyhydroxides are thought to control aqueous manganese 

concentrations. Using the distribution coefficient 

measured by Bodine and others (1965) for partitioning 

manganese between the aqueous phase and a coexisting 

calcite phase, and the mean concentrations of calcium 

(0.28 mmol/L) and manganese (0.00049 mmol/L) in anoxic ground waters of the Lance-Fox Hills aquifer, it has 

been calculated that a calcite in equilibrium with these 

ground waters should contain an average of 3-mole percent rhodochrosite. Saturation indices computed for a 

percent Mn calcite ideal solid solution (Berner, 1975) 

with a calculated stability of 1,121 kilojoules per mole, 

presented in table 28, indicate saturation with respect to 

this phase in most samples and suggest possi- ble control of manganese concentrations by such a solid solution in all but oxygenated waters.

Saturation indices for strontianite clearly indicate 

undersaturation, but, unlike Judith Basin samples, 

cluster near a value of 0.76, again suggesting possible 

mineralogic control by a carbonate solid solution. The 

distribution coefficient of 0.14 measured by Holland 

others (1964) for the partitioning of strontium between 

an aqueous phase and calcite, and the mean concentrations of strontium (0.013 mmol/L) and calcium

I 28. Saturation indicc\ of a 3. 0-moie percent mans>ane^e calcite \nlid solution in

waters from the Lane?-Fox ////A aquifer. Powder River Basin, \Vyuming

Sample Saturation indices

Calcite Manganese calcite

Lance Formation

PI

PO

PI 1 '

Pl'-i

PI 4

PI a

0

1 0

. 30

. -14

.21

 .13

 .16

0.07 

.23 

.08

.06 

.09

Mean and standard deviation - 0.11 ±0.17 0.04 ±0.16

Fox Hills Sandstone

P')

P3

P4

Pf)

P6

P7

PH)

PI 1

PI 5

P16

P17

Me-an and standard deviation -

0.20

.04

.04

nn

.11

.09

.13

.33

.07

1 Q

.16

0.16 ±0.16

0.15

- .03

.07

.02

.05

.07

.24

.02

.37

.29

.09

0.08 + 0.15

(0.94 mmol/L), indicate that a calcite in equilibrium with 

ground waters of the Lance-Fox Hills Sandstone 

should contain 0.2-mole percent strontianite. The free 

energy of a S^^Ca^^CC^ ideal solid solution (Berner, 

1975) is 1,130 kilojoules per mole, and saturation indices calculated from this value demonstrate saturation 

with respect to this solid solution in most samples (table 

29). Strontium, derived from the dissolution of impure 

carbonates and sulfates, may be controlled in the 

aqueous phase, then, by sorption effects and equilibrium with a Sr-calcite solid solution.

Sulfate mineral phases also demonstrate relationships 

similar to those observed in the Kootenai Formation. 

Both gypsum and celestite show widely fluctuating, but 

invariably negative, saturation indices. Control of strontium, calcium, or sulfate concentrations by these phases 

is therefore unlikely, as is the presence of substantial 

quantities of these phases.

Apatite saturation indices (table 27), computed for 

both a pure fluorapatite [Ca5 (PO4 )3F] and a cryptocrystalline carbonate fluorapatite [Ca974Na021Mg005(PO4)319

(CO F1 OTT 3'0.8lr 2.32^-"-0.28' pK =-120.7] (Chien and Black,

1975), suggest saturation with respect to an apatite
C62 REGIONAL AQUIFER-SYSTEM ANALYSIS

T\BI I 2C'. Saturation mdue\ of a (I 2-mole berient strontium calcite wild solution in

ivattrs from the Lance-Fox Hill\ aquifer, Powder River Basin, Wyoming

Sample

Lance

PI

P8

Pr>

PI 3

PI 4

pi Q

Mean and standard deviation -

Calcite

Formation

0.13

.30

.14

.21

.16

o.n ±o.:

Saturation indices

Strontium calcite

0.06

.21

.06

- .30

.04

.07

17 0.02 + 0.17

Fox Hills Sandstone

P2

P3

P4

P ^

P6

P7

Ptn

PI 1

PI 3

P16

PI 7 

 0.20

.04

 .04

f\f\

 .13

 .33

 .07

1 Q

OQ

\ft

0.13

.11

.06

.09

01

(11

.06

.23

nn

on

r\n

AQ

Mean and standard deviation - 0.16 + 0.16 0.08 + 0.15

having a composition and stability intermediate between these two phases. The two deep basin samples, 

P9 and PI7, are undersaturated with respect to both of 

these apatites, indicating that these samples may be in 

equilibrium with a more crystallized apatite or one of 

slightly different composition. The occurrence of apatite 

in the Fox Hills Sandstone implies that phases similar 

to these are present in the aquifer, and are probably in 

equilibrium with the waters sampled. The phase manganese biphosphate (MnHPO4) is also saturated in most 

samples from the Lance-Fox Hills aquifer, indicating 

that aqueous manganese and phosphorous concentrations may also be influenced by such a phase.

Although ferric oxyhydroxides have been identified in 

neither X-ray nor petrographic analyses, their presence 

in the Lance-Fox Hills outcrop is implied by numerous 

references to ferruginous concretions m the stratigraphic descriptions of Robinson and others (1964). 

Control of aqueous iron concentrations by amorphous 

ferric hydroxide, rather than by goethite or hematite, in 

oxidizing waters is apparent from the saturation indices 

presented in table 30. Oversaturation with respect to 

Fe(OH)3)am)(pK=38.5) is indicative of active precipitation of this phase (IAP=-37.1 to -38.1) (Whittemore 

and Langmuir, 1975), possibly due to the oxidation of 

pyrite by O0 (reaction 9), followed by the hydrolysis and 

precipitation of ferric iron. Oxidation of pyrite would 

also account for increases in sulfate concentration in the 

outcrop area, as well as the localization of ferric hydroxides in concretions, similar to the occurrence of pyrite in 

the reduced environment.

Although manganese oxides were not detected by 

X-ray diffraction in samples from the Fox Hills outcrop, 

it is possible that these common phases have escaped 

detection by their poorly crystallized nature. In the five 

oxidizing ground-water samples of the Lance-Fox Hills 

aquifer, both manganite and birnessite are highly oversaturated, with mean saturation indices and standard 

deviations of 4.24±0.76 and 7.61 ±0.74, respectively, indicating that these phases probably do not control 

aqueous manganese concentrations in the oxidizing 

environment.

In sulfidic waters, amorphous ferrous sulfide, rather 

than crystallized ferrous sulfides such as pyrite, controls dissolved iron cncentrations (table 30). Slight oversaturations with respect to FeS(am) are most likely due to 

interferences from other reduced species during the 

determination of dissolved sulfide. Although goethite 

saturation indices in sulfidic waters approach zero in 

many cases, the absence of hematite (a well crystllized 

oxyhydroxide) in core samples, as well as the occurrence 

of pyrite in the aquifer, suggests control of iron concentrations by amorphous ferrous sulfide in reduced 

waters, rather than by a ferric oxyhydroxide phase.

REACTION SIMULATION

The discussion of ground-water chemistry, 

mineralogy, and petrography of the Lance-Fox Hills 

aquifer has led to the proposal of several mineral phases 

and chemical reactions that may control ground-water 

chemistry in this aquifer. Saturation with respect to the 

aluminosilicates beidellite, muscovite, kaolinite, and 

allophane, in addition to the carbonates calcite and dolomite, suggests that one or more of these phases may 

control the ground-water chemistry of the Lance-Fox 

Hills aquifer. Carbon isotopic ratios indicate possible 

control of dissolved carbonate concentrations by carbonate minerals.

Oxidation potentials decrease within the aquifer outcrop area from highly oxidizing to strongly reducing 

values. Consumption of O2 and NO~ in combination 

with the production of additional N2 and CT is indicative 

of biologically mediated O2 and NO~ reduction reactions. Saturation indices for Fe(OH)3(am) in oxidized 

ground waters imply active precipitation of this phase, 

possibly as a result of the oxidation of pyrite. An abun-
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C63

Sample Hematite Goethite Fe(OH)3 m a Pyrite FeS(ar b

n)

Sulfide absent

P4

P7

P15 

P16 

Mean and standard deviation -

18. 89 ±0.32 

16.41 ± 1.15 

16.01 ±2.15 

17.05+ .82 

18.36 ± .34

17.34 + 1.24

6.98 + 0.16 

5.74± .57 

5.54± 1.07 

6. 06+. .41 

6.71+ .17

6.21 ± 0.62

3. 06 ±0.16 

1.73+ .57 

1.57+ 1.07 

2.06+ .41 

2.70+ .17

 

 

 

2.22 + 0.64

Sulfide present

P6 

P8

IX)

Pit) 

PI 7

pi u

Mean and standard deviation -

6.79± 1.15 

6.70± 1.48 

6.99 ± .67 

2.57± .39 

4.36± .32 

5.34 + 4.17 

7.02± .82 

4.88 ± .31 

6.21 ±4.17 

6. 46 ±4. 17 

3.45+ .36 

4.94 + 2.15 

3.31 + .31

5.31 ± 1.53

0.91 

.86 

1.01 

- 1.18 

-.28 

.20 

1.00 

.02

.61

.75 

.75

.07

0 1

0.17

± 

± 

± 

±

± 

± 

± 

± 

± 

± 

± 

± 

+

+

0.57 

.74 

.33 

.20 

.16 

2.09

.16

2.09 

2.09 

.18 

1.07 

.16

0.76

2.68 + 0.

2.48+ .

-5.10+ . 

4.34+ .

3 68 + 2,

- 1.87+ . 4.06+ .

2.51 +2.

2.95 + 2.

4.69+ . <> ')o + i

- 3.37+ 1.

57 

71- 

33 

20 

16 

.09 

41 

.16 

.09 

.09 

.18 

.07 

.16

.11

8.60+ 1. 

8.87 ± . 8.81 + . 

9.18+ . 

11.03+ . 8.85 + 2. 

8.06 + , 10.32+ . 

7.42 + 2 

8. 46 ±2. 

8.20 + 

6.64 + 1

10.20+ 1

8.82+ 1

,02 

.81 

.91 

,87 

58 

13 

.94 

,86 

25 

,13 

.59 

.37 

.71

.19

1.00 + 

1.22 + 

1.15 +

.11 +

1.3( t +

.74 +

.89 + 

.86 + 

.33 ± 

.86 + 

.62 +

02 + 

.49 ±

0.63 +

0.70 

.75 

.52 

.45 

.30 

2.09 

.58 

.43 

2.12 

2.09 

.31 

1.15 

.86

0 58

dance of organic carbon both in the aquifer and in the 

ground water, the presence of sulfate reducing bacteria 

in the Powder River Basin, and the appearance of dissolved sulfide following the reduction of O2 and NO~ are 

suggestive of sulfate reduction. However, isotopic and 

mass transfer data are insufficient to prove this point. 

In the deeper portions of the basin, the disappearance of 

sulfate, with a concurrent increase in methane concentrations and <5 13C values, indicates the activity of 

methanogenic bacteria. In the reducing ground waters, 

saturation with respect to both siderite and Fe(OH)3(am) 

suggests control of dissolved iron concentrations by 

these phases.

Increasing sulfate concentrations along the flow path 

are most likely due to the dissoluton of gypsum present 

in the upper part of the Lance Formation. Increasing 

chloride concentrations are probably the result of minor 

upward leakage from the underlying Pierre Shale, or 

from the leaching of interbedded shales. Finally, uniformly decreasing log molar ([Ca]+[Mg])/[Na]2 ratios are 

indicative of cation exchange.

Simulation of the above reactions and concurrent 

equilibration with appropriate minerals has allowed for 

a determination of possible controlling phases along the

northern flow path of the Fox Hills Sandstone. As in the 

Second Cat Creek sandstone models, ion exchange was 

simulated by fixing (aca+ 2 )/(aNa+)2 and < aMg+2)/(a +)2 

ratios to their observed values for each sample. Minor 

additions of chloride from shale confining beds were 

simulated by dissolution of appropriate amounts of 

halite. Increases in sulfate concentration, originating 

from either gypsum dissolution or pyrite oxidation, 

were modeled by dissolving appropriate amounts of 

gypsum in the outcrop area or natrona (Na2 SO4 ) 

downdip. The choice of these two phases to represent 

sulfate dissolution is based on the predominant cation 

in solution, calcium in the outcrop area and sodium 

downdip, and does not directly imply the presence of 

these minerals in the aquifer. Dissolved iron concentrations were controlled through equilibration with FeS(am) 

and siderite in reducing waters, and with Fe(OH)3(am) 

(pKsp =37.0) in oxidizing waters.

A water similar in major element composition to 

Wyoming rainwater (Junge and Werby, 1958) was equilibrated with atmospheric partial pressures of CO2 and 

O2 using PHREEQE and was used as the initial solution for the following reaction simulations. In order to 

attain CT values identical to those calculated for outcrop
C64 REGIONAL AQUIFER-SYSTEM ANALYSIS

area waters, the unsaturated zone Pco before equilibration with mineral phases were required to be on the 

order of 10~ 12 atm for the carbonate simulations and 

10~09 atm for the aluminosilicate simulations. Thorstenson and others (1979) measured a Pco of 10~07 atm in 

the unsaturated zone at Gascoyne, North Dakota, indicating that although these modeled CO2 concentrations appear somewhat high, they are realistic for 

liqpiitic sediments of the type found in the study area. 

The O2 and NO~ present in the simulated recharge 

water were reduced by the addition of an appropriate 

amount of organic carbon. For the reaction simulations, 

the stability of beidellite was assumed to be similar to 

that of another smectite, the fixed-composition Namontmorillonite phase calculated by Helgeson (1969). 

As shown in figure 48, subsequent equilibration with 

neither calcite, dolomite, calcite and dolomite, kaolinite 

and Na-montmorillonite, kaolinite and muscovite, 

kaolinite, Na-montmorillonite and muscovite, nor allophane results in sufficiently high CT concentrations for 

the downdip samples, indicating that additional CO2 

sources must be present. As a result, CO2 was added to 

the simulated ground waters in order to match the calculated CT concentration for each sample along the 

northern flow path. The pH values resulting from equilibration with the phases listed above are shown in figure 

49.

The calcite and dolomite model predicts the fieldmeasured pH values more precisely than the remaining 

models, and mass transfer calculated using this model is 

indicative of dedolomitization driven by cation exchange and dissolution of CO2 . A problem arises at this 

point, however, regarding the choice of CO2 sources. If 

CO9 is derived from a maturing high-volatile organic 

carbon, the associated mass transfers for siderite and 

ferrous sulfide will be relatively low. Alternatively, if 

CO2 is produced within the aquifer during oxidation of 

organic matter by sulfate reducing bacteria, H2S will be 

generated in quantities of roughly half the amount of 

CO2 produced. In order to maintain equilibrium, siderite 

will be required to dissolve, releasing enough ferrous 

iron to precipitate the sulfide produced. The dissolution 

of siderite will consume a portion of the CO2 acidity, 

thereby decreasing the solubility of dolomite and altering carbonate mass transfer. This effect is demonstrated in figures 50 and 51, which present calculated 

mass transfers for lignite coalification (fig. 50) and 

sulfate reduction (fig. 51) CO2 sources. The differences 

between these two models appear minimal, but should 

be considered in the adjustment of carbon-14 ages and 

the calculation of flow rates.

Calculated mass transfers indicate controls on 

ground-water chemistry similar to those proposed for 

the Second Cat Creek sandstone, namely dedolomitization driven by lignite coalification or sulfate reduction:

O 8.0 -

o

* 7.0 H

EXPLANATION

 CALCULATED CONCENTRATION 

^MODELED CONCENTRATION

C A LCITE +DOLOMITE

KAOLINITE + Na MONTMORILLONITE

37P6

KAOLINITE+MUSCOVITE, KAOLINITE+MUSCOVITE+Na MONTMORILLONITE

P2 P3

0 10 20 30 40 50 60 

DISTANCE FROM OUTCROP ALONG NORTHERN FLOW PATH, IN KILOMETERS

FIGURE 48. Comparison of calculated total inorganic carbon concentrations with modeled concentrations for mineral equilibration 

with no CO9 dissolution following the reduction of O2 and NO~, 

northern flow path of the Fox Hills Sandstone, Powder River 

Basin, Wyoming. All simulations include equilibration with 

calcium-sodium exchange and magnesium-sodium exchange. 

Simulations P6, P2, P3, and Pll also include equilibration with 

amorhous ferrous sulfide and siderite; simulation P7 includes 

equilibration with amorphous ferric hydroxide.

EXPLANATION

 MEASURED pH

-^'MODELED pH

KAOLINITE+MUSCOVITE, KAOLINITE+MUSCOVITE + Na-MDNTMORILLONITE

LlNITE+Na-MONTMORILLONITE

ALLOPHANE

P3

0 10 20 30 40 50 60 

DISTANCE FROM OUTCROP ALONG NORTHERN FLOW PATH, IN KILOMETERS

FIGURE 49. Comparison of field-measured pH values with modeled pH values for mineral dissolution with CO2 dissolution, 

northern flow path of the Fox Hills Sandstone, Powder River 

Basin, Wyoming. All simulations include equilibration with 

calcium-sodium exchange and magnesium-sodium exchange. 

Simulations P6, P2, P3, and Pll also include equilibration with 

amorphous ferrous sulfide and siderite; simulation P7 includes 

equilibration with amorphous ferric hydroxide.
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C65

RAIM WATFR ^

UEAS DISSOLVE:

Ca 0.02 ^4m >l'L02

 0.11 mmolA HALITE 

Na 9:°_L n,p7 mmr,|/| rjYPSUM

Cl 0.002 0.03 mmol/L Ca(NO? ) ?

SO^ 0.01 3 51 mmol/L C02

C 6.04 0-!8 mmol/L CH 2O 

 L- 2.20 mmol/L DOLOMITE

0.45 umol/L Fe(OH)3(g .

PRECIPITATE: 

1.31 mmolA CALCITE

ADSORB: 

0.09 mmolA Ca 

0.38 mmol/L Mg

DESORB: 

0.96 mmol/L Na

SAMPLE P7

Ca

Mg

Na

Cl

so^

CT

NOj

°2

Fe

PH

Eh

MEAS.

1.3

1.7

1.0

0.15

0.58

6.8

0.056

0.062

.0007

7.50

0.801

MOD.

1.4

1.8

1.1

0.11

0.58

6.8

0.056

0.062

.0004

7.41

0.850

^

DISSOLVE: 

0.10 mmol/L GYPSUM 

0.55 mmol/L C02 

0.14 mmol/L CH 20 

0.90 mmol/L DOLOMITE 

5.7 umol/L SIDERITE 

1.1 umol/L FeS (gm)

PRECIPITATE: 

0.92 mmol/L CALCITE

ADSORB: 

0.67 mmol/L Ca 

1.96 mmol/L Mg

DESORB: 

5.26 mmol/L Na

DISSOLVE: 

0.14 mmol/L HALITE 

0.38 mmol/L C02 

0.26 mmol/L DOLOMITE 

0.59 umol/L FeS(gm)

PRECIPITATE: 

0.14 mmol/L CALCITE 

0.58 umol/L SIDERITE

ADSORB: 

0.13 mmol/L Ca 

0.27 mmol/L Mg

DESORB: 

0.79 mmol/L Na

SAMPLE F

Ca

Mg

Na

Cl

so4

CT

2H2S

Fe

PH

Eh

MEAS.

0.72

0.66

6.1

0.07

0.69

8.3

0.009

0.018

7.71

-0.219

vj

MOD.

0.78

0.71

6.3

0.11

0.69

8.3

0.002

0.007

7.65

0.228

^

SAMPLE P3

Ca

Mg

Na

Cl

so4

CT

IH2S

Fe

PH

Eh

MEAS.

0.04

0.00

14.4

0.65

1.6

10.3

0.006

.0014

8.77

 0.302

MOD.

0.03

0.00

14.0

0.65

1.5

10.1

0.001

.0003

8.76

 0.308

^

DISSOLVE:

0.40mmol/L HALITE 

0.82 mmol/L Na2S04 

0.13 mmol/L C02 

0.62 mmol/L DOLOMITE 

1.4umol/L SIDERITE

PRECIPITATE: 

0.34 mmol/L CALCITE 

8.3 umol/L FeS (gm)

ADSORB: 

1 .02 mmol/L Ca 

1 .32 mmol/L Mg

DESORB: 

4.68 mmol/L Ma

 *

DISSOLVE:

0.03 mmol/L HALITE 

0.12 mmol/L DOLOMITE 

0.04 umol/L FeS(gm)

PRECIPITATE: 

0.11 mmol/L CALCITE 

0.15 umol/L SIDERITE

ADSORB: 

0.02 mmol/L Ca 

0.12 mmol/L Mg

DESORB:

0.28 mmol/L Ma

SAMPLE P2

Ca

Mg

Na

Cl

so4

CT

ZH 2S

Fe

PH

Eh

MEAS.

0.

0.

05

02

13.0

0.51

1.6

9.3

0.015

.0005

8.91

-0.317

MOD.

0.04

0.02

13.0

0.51

1.5

9.3

0.001

.0003

8.79

0.310

SAMPLE P1 1

Ca

Mg

Ma

Cl

so4

CT

IH 2S

Fe

PH

Eh

MEAS,

0.03

0.00

13.9

0.68

1 5

10.0

0.003

.0002

8.92

 0.322

MOD.

0.03

0.00

14.3

0.68

1.5

10.2

0.002

.0002

8.82

 0.317

FIGURE 50. Calculated mass transfer for the calcite-dolomite-cation exchange-CO2 dissolution model of ground-water evolution along the 

northern flow path of the Fox Hills Sandstone, Powder River Basin, Wyoming. Concentrations in millimoles per liter (mmol/L) and 

micromoles per liter (/xmol/L); Eh in volts.

+ 2 2CaMg(CO3 )2 +FeCO3 +2CH2O+SO4- 2 +2Na2EX+Ca 

=3CaCO3 +FeS+4HCO3-+2MgEX+4Na+

in combination with additional cation exchange (reaction 28 and 29). The addition of kaolinite and Namontmorillonite to these models results in only minor 

mass transfer changes, indicating that these aluminosilicates influence major element geochemistry primarily as cation exchange substrates. As illustrated previously in the discussion of CO2 sources, these proposed 

mass transfer models are not mathematically unique 

solutions, but instead represent plausible hypotheses to 

explain the geochemical evolution of ground-water composition in the Fox Hills Sandstone.

Comparison of measured or calculated concentrations 

to simulated concentrations, plotted in figure 52, shows 

good agreement in most cases. Iron presents the largest 

deviations, with analytical concentrations being much 

higher than modeled concentrations in several cases. As 

discussed previously, such deviations may indicate 

sample contamination by windblown dust.

CARBON ISOTOPIC EVOLUTION

Simulation of carbon isotopic compositions was accomplished using a FORTRAN computer code that incorporates the isotopic evolution equations of Wigley 

and others (1978). The freshwater carbonate <5 13C value 

of Keith and Weber (1964) (-4.8±5.1), and the organic 

6 13C value of Craig (1954) ( 28±4) were chosen to represent the isotopic composition of carbon sources in the 

aquifer. These estimated 613C source values were applied to the mass transfers calculated in figures 50 and 

51 to yield downgradient 613C values to be expected 

assuming that these reaction schemes actually control 

ground-water carbon composition. The coalification of 

lignite and sulfate reduction models yield nearly identical <5 13C values. Therefore, carbon isotopes do not 

allow for differentiation between these two carbon 

sources. Modeled and observed <5 13C values are plotted 

in figure 53, and show relatively good agreement, 

although two of the analytical values fall slightly above 

or below their corresponding modeled values. The scatter in observed <5 13C values may be caused by several influencing factors, the most significant of which is the
C66 REGIONAL AQUIFER-SYSTEM ANALYSIS

RAIN WATER b

MEAS. DISSOLVE:

0.24 mmol/L 00 

 

Ca ^ 0 02 0.11 mmol/L HALITE 2 

Na 0.01 o.57 mmol/L GYPSUM

Cl 0.002 0.03 mmol/L Ca(NO? ) 9

S04 0.01 3.51 mmol/L C02

"c 004 0.18 mmol/L UH 2O

 2.20 mmol/L DOLOMITE 

0.45 umol/L Fe(OH)3(am)

PRECIPITATE: 

1.31 mmol/L CALCITE

ADSORB:

0.09 mmol/L Ca 

0.38 mmol/L Mg

DESORB: 

0.96 mmol/L Na

SAMPLE P7

Ca

Mg

Na

Cl

so4

CT

N03

°2

Fe

pH

Eh

MEAS.

1.3

1.7

1.0

0.15

0.58

6.8

0.056

0.062

.0007

7.50

0.801

MOD.

1.4

1.8

1.1

0.11

0.58

6.8

0.056

0.062

.0004

7.41

0.850

1^

DISSOLVE:

0.38 mmol/L GYPSUM

0.69 mmol/L CH 20 

0.85 mmol/L DOLOMITE

0.28 mmol/L SIDERITE

PRECIPITATE: 

1.20 mmol/L CALCITE 

0.27 mmol/L FeS(am)

ADSORB: 

0.67 mmol/L Ca

1 96 mmol/L Mg

DESORB: 

5.28 mmol/L Na

DISSOLVE:

0.14 mmol/L HALITE 

0.19 mmol/L Na2 S04 

0.36 mmol/L CH 2 0 

0.19 mmol/L DOLOMITE 

0.18 mmol/L SIDERITE

PRECIPITATE: 

0.14 mmol/L CALCITE 

0.1 8 mmol/L FeS(am)

ADSORB: 

0.06 mmol/L Ca 

0.20 mmol/L Mg

DESORB: 

0.51 mmol/L Na

SAMPLE P6

Ca

Mg

Na

Cl

so4

C T

IH2S

Fe

pH

Eh

MEAS.

0.72

0.66

6.1

0.07

0.69

8.3

0.009

0.018

7.71

-0,219

MOD.

0.78

0.71

6.3

0.11

0.69

8.3

0.002

0.007

7.65

0.228

V

SAMPLE P3

Ca

Mg

Na

Cl

so4

C T

^"LJ Q

Fe

PH

Eh

MEAS.

0.04

0.00

14.4

0.65

1.6

10.3

0.006

.0014

8,77

-0.302

MOD.

0.03

0.00

14.1

0.65

1.5

10.1

0.001

.0003

8.81

-0.311

^

DISSOLVE: "

0.40 mmol/L HALITE 

0.94 mmol/L Na2S04 

0.13 mmol/L CH 2 0 

0.59 mmol/L DOLOMITE 

66. umol/L SIDERITE

PRECIPITATE: 

0.35 mmol/L CALCITE 

73. umol/L FeS, ,

ADSORB: 

0.99 mmol/L Ca 

1 .29 mmol/L Mg

DESORB: 

4.56 mmol/L Na

DISSOLVE: 

0.03 mmol/L HALITE 

0.11 mmol/L DOLOMITE 

0.09 umol/L FeS (a .

PRECIPITATE: 

0.11 mmol/L CALCITE 

0.17 umol/L SIDERITE

ADSORB: 

0.00 mmol/L Ca

0.11 mmol/L Mg

DESORB:

SAMPLE P2

Ca

Mg

Na

Cl

so4

C T

2H2S

Fe

pH

Eh

MEAS.

0.05

0.02

13 .0

0.51

1.6

9.3

0.015

.0005

8.91

MOD.

0.05

0.02

13.2

0.51

1.6

9.3

0.001

.0003

8.79

f-0.3l7J-0.309

SAMPLE P1 1

Ca

Mg

Na

Cl

so4

CT

IH 2S

Fe

pH

Eh

MEAS.

0.03

0.00

13.9

0.68

1. 5

10.0

0.003

.0002

8 92

-0.322

MOD.

0.03

0.00

14.4

0.68

1.5

10.2

0.002

.0002

8.85

 0.318

0.23 mmol/L Na

FIGURE 51. Calculated mass transfer for the calcite-dolomite-cation exchange-sulfate reduction model of ground-water evolution along the 

northern flow path of the Fox Hills Sandstone, Powder River Basin, Wyoming. Concentrations in millimoles per liter (mmol/L) and micromoles per liter (/imol/L); Eh in volts.

uncertainty of siderite <5 13C values and the lenticular 

nature of the flow system. Siderite has been shown to 

vary in <5 13C from highly negative (Hodgson, 1966) 

(<5 13C= 33.3) to moderately positive (Fritz and others, 

1971) (<5 13C=12.8) values, with large variations possible 

in samples collected within a meter of each other. The 

sulfate reduction model requires dissolution of significant amounts of siderite, and isotopic modeling considered siderite <5 13C values to be equivalent to 

freshwater carbonate <5 13C values (Weber and others, 

1964). Variations in siderite <5 13C values might cause 

significant variations in <5 13C values of dissolved carbonate. The lenticular nature of the aquifer impedes 

vertical mixing within the formation, thus samples 

obtained from different areas may have slightly different evolutionary histories and, therefore, different 513C 

values due to contact with carbonates of differing isotopic composition.

Calculation of adjusted carbon-14 ages yields virtually identical values for both models and suggests that 

the mechanism of CO2 production may not be an important factor in determining ground-water ages. Adjusted

ages of 19,800 to 33,400 years were computed for P2, 

P3, and Pll using the model of Wigley and others 

(1978) and measured initial carbon-14 activities as 

discussed in the section on Radioactive Isotopes. The 

isotopic compositions and mass transfers used in the 

model are listed in table 25 and figures 50 and 51, 

respectively. Flow rates calculated from these ages, 

presented in table 31, compare favorably with hydrologically derived values of 0.16-1.08 m/yr. Scatter in the 

ages may be the result of variable flow rates within the 

aquifer due to the lenticular nature of sand lenses and 

the presence of many interbedded shale beds, as well as 

failure to consider possible extreme <5 13C values for 

siderite.

CONCLUSIONS

Although the controlling reactions cannot be unequivocally determined, the major element geochemistry of 

confined ground waters in the Second Cat Creek sandstone and Lance-Fox Hills aquifers of the Northern
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS

EXPLANATION °'°2 ~1

RANGE OF MEASURED CONCENTRATION 

AT 95 PERCENT CONFIDENCE

C67

 MODELED CONCENTRATION 

CONCENTRATIONS IN MILLIMOLES PER LITER

O 0.01-

cc

0.00

D

o 1.0H

< 

o

0.0

O 11.on

cc

O 

O

(D 

cc 

O

< 

O

9.0-

7.0-

c/j LU

z

(D

2.0i

1.0-

0.0

LU

D

C/5

2.0 -i

0.0-

20-i

~ 10-

0

P2 P3 P11

0 10 20 30 40 50 60

DISTANCE FROM OUTCROP 

ALONG NORTHERN FLOW PATH, IN KILOMETERS

LUQ

I

O

0.5-

0.0' P2 P3 P11

0 10 20 30 40 50 60

DISTANCE FROM OUTCROP 

ALONG NORTHERN FLOW PATH, IN KILOMETERS

FIGURE 52. Comparison of measured and modeled concentrations of selected constituents in waters from the northern 

flow path of the Fox Hills Sandstone, Powder River Basin, Wyoming.

Great Plains appears to be controlled by the incongruent dissolution of dolomite driven by CO2 

dissolution and cation exchange. As a result, the waters 

evolve from a near neutral pH, low dissolved solids, 

calcium, and magnesium bicarbonate types to high pH, 

high dissolved solids, sodium bicarbonate types. The 

source of CO2 cannot be unequivocally determined, but

is assumed to be either the coalification of lignite or the 

oxidation of organic carbon by sulfate-reducing 

bacteria. Mass transfer and carbon isotopic evolution 

computations indicate that these reactions will affect 

dissolved carbonate <5 13C values similarly, with modeled 

isotopic compositions corresponding favorably to measured 613C values for both processes.
C68 REGIONAL AQUIFER-SYSTEM ANALYSIS

EXPLANATION

 MEASURED ISOTOPIC RATIO 

MODELED ISOTOPIC RATIO

DISTANCE FROM OUTCROP ALONG NORTHERN FLOW PATH, IN KILOMETERS

FIGURE 53. Comparison of measured 613C values and the range of 

modeled 613C values for waters from the northern flow path of the 

Fox Hills Sandstone, Powder River Basin, Wyoming. The range of 

modeled 613C values reflects the uncertainty in the actual 613C compositions of aquifer carbon sources.

TABI.K 31. Comparison of radiometric and hydrologic flow rato for the Lance-Fox 

Hills aquifer of the Powder River Basin, Wyoming

Sample

P2

t>'-t

PI 1

Hydrologic flow rate 

(meters per year)

0.16-1.08

.16-1.08

.16-1.08

Radiometric flow rate 

(meters per year)

1.07

.77

<2.14

Despite the profusion of aluminosilicate minerals in 

the aquifers studied, ground-water pH and dissolvedcarbonate chemistry appear to be controlled by 

equilibration with the often less abundant carbonate 

minerals. Clay minerals indirectly influence groundwater chemistry by providing substrates for cation exchange, thereby allowing for the high-pH, high- bicarbonate conditions attained in many aquifer systems. 

Similar conclusions were reached by Thorsten- son and 

others (1979) in their modeling study of the Fox 

Hills-Hell Creek aquifer in the Williston Basin, and by 

Foster (1950) in her laboratory study of carbonate- cation 

exchange equilibria.

The paucity of organic carbon poises oxidation potentials at values near the redox interface and effectively 

prevent the onset of sulfate reduction in the Second Cat 

Creek sandstone, suggesting that CO2 may be 

migrating into the aquifer from adjacent carbonaceous 

strata. In the Lance-Fox Hills aquifer, the presence of 

H2S and large quantities of carbonaceous material, as 

well as the ubiquitous nature of sulfate reducing bacteria, supports the hypothesis that CO2 is generated 

primarily within the aquifer, probably through the 

reduction of sulfate.

Significant quantities of methane were observed in 

dissolved gases extracted from the two deepest, and 

presumably oldest, samples from the Fox Hills Sandstone. Anomalously high <5 13C values indicate that 

methanogenic bacteria are active in these waters, 

becoming the dominant microbial population following 

the depletion of sulfate by more energetically efficient 

bacteria.

Dissolved gas samples also verify that measured Pco 

values may be significantly different than those calculated from field-measured pH due to exchange with atmospheric CO2 between the time of sample collection 

and measurement. Calculation of pH from measured 

Pco indicates that the in situ aquifer pH may differ 

from the field-measured value by up to 0.2 units for 

poorly buffered, low alkalinity waters, whereas the effects of CO2 exchange appear minimal for well buffered, 

high alkalinity samples.

Sulfate is presumed to be derived from the dissolution 

of gypsum or the oxidation of pyrite, either within the 

aquifer or in adjacent strata, whereas chloride enters 

the aquifer as minor amounts of leakage or diffusion 

through shale confining layers. As demonstrated in the 

Judith Basin study area, large increases in the concentrations of either of these constituents may be indicative of leakage into the aquifer through faults, fractures, 

or other highly permeable zones. In the Powder River 

Basin, recharge of oxidized waters and subsequent mixing with older, reduced waters results in high sulfate 

concentrations from migration of the redox interface 

and concurrent pyrite oxidation.

Aqueous iron concentrations are effectively controlled by amorphous ferric hydroxide, siderite, or amorphous ferrous sulfide, depending on Eh conditions. 

Manganese concentrations appear to be regulated by a 

solid solution between calcite and rhodochrosite, 

although saturation with respect to MnHPO4 in waters 

of the Lance-Fox Hills aquifer may indicate control of 

dissolved manganese by a phosphate phase. Apatite is 

also saturated in the Lance-Fox Hills aquifer, and probably controls dissolved phosphate concentrations. A 

solid solution between calcite and strontianite may influence strontium concentrations in the Lance-Fox 

Hills aquifer, although such solid solution behavior is 

unlikely in the Second Cat Creek sandstone.

The use of ground-water geochemistry can provide 

invaluable information regarding the hydrology of confined sandstone aquifers, especially when used in 

conjunction with available hydrologic data. Areas of 

significant recharge to the artesian systems are identified by the presence of relatively low dissolved solids 

concentration, calcium, and magnesium as the predominant cations, and high tritium and carbon-14 activities. 

Several samples of the waters from the outcrop did not 

exhibit all of these properties, serving as an indication
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C69

of minimal recharge to the confined system in these 

areas. Flow directions, determined by regional potentiometric surface maps, were geochemically characterized by increasing dissolved solids concentrations 

and log molar ([Ca]+[Mg])/[Na]2 ratios. It was determined, however, that variations in these operational 

parameters may become ambiguous in zones where 

waters of dissimilar chemistry converge and mix. 

Recharge temperatures for both the Second Cat Creek 

sandstone and Lance-Fox Hills aquifer, calculated from 

618O and <5D measurements and dissolved argon concentrations, are significantly lower than mean annual surface temperatures, suggesting that recharge to these 

systems is derived primarily from spring snowmelt, 

rather than warmer late spring and summer storms.

Areas of leakage or mixing between aquifers are geochemically characterized by abrupt changes in major 

element and isotopic chemistry, which are not characteristic of normally observed ground-water evolution. 

Such changes in a fractured, structurally deformed area 

of the Judith Basin were interpreted as evidence of upward leakage from deeper aquifers. Areal variations in 

major element chemistry similarly indicate the presence 

of a discharge zone along the Belle Fourche structural 

arch in the Powder River Basin.

Perhaps the most significant contribution of groundwater chemistry to the characterization of the 

hydrology of a confined aquifer system lies in the determination of ground-water ages and flow rates. The 

adjustment of carbon-14 ages using computed mass 

transfer and stable carbon isotopic data allows for the 

calculation of flow rates that are directly comparable to 

those derived from measured aquifer constants and 

potentiometric data. The recent success in the adjustment of carbon-14 ages utilizing mass transfer data obtained from detailed geochemical models, both in this 

study and in others using this approach, is indicative of 

the usefulness of this technique in determining flow 

rates in geochemically well-defined systems.

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SUPPLEMENTAL DATA
C74 REGIONAL AQUIFER-SYSTEM ANALYSIS

GEOCHEMICAL PRINCIPLES 

SOLUTION-MINERAL EQUILIBRIA

LAW OF MASS ACTION

The driving force of a chemical reaction can be related 

to the effective concentrations, or activities (a;) of the 

reactants and products, through the Law of Mass Action. In a system at equilibrium, for a particular set of 

pressure and temperature conditions, the rate of reaction 

in a forward direction is equal to the rate of reaction in a 

reverse direction. For the reaction:

bB+cC=dD+eE,

equilibrium conditions may be defined at a particular 

temperature and pressure by the thermodynamic equilibrium constant (K ):

K«,=

EQUILIBRIUM CONSTANTS

The K for a reaction of 25° Celsius (°C) and one bar

eq

total pressure may be calculated from the difference in 

the standard free energies of formation (AG^) between 

products and reactants:

.0 \ 

fC '

reaction f products f reactants,

f reaction

where:

R=gas constant (8.3147

Ke<] >

Joules

mole ° Kelvin

), and

T= standard temperature (298.15° Kelvin).

Because an equilibrium constant is dependent on temperature, the value calculated for 25 °C (298.15° Kelvin) 

must be adjusted if it is to be used at another temperature. In a system at constant pressure, the equilibrium 

constant is related to temperature by the standard enthalpy of reaction (AH^X) and the heat capacity of reaction at constant pressure (AC°RX):

X r

2.303RT0

r 2.303R

T2

Tr

2.303R

where T2 represents the temperature of interest and Tr 

represents the reference temperature (298.15°K). Unfortunately, heat capacity data are unavailable for a majority of aqueous species; in most instances, the above expression cannot be solved. A simplification of the above 

expression, the Van't Hoff equation, assumes that AC°RX 

equals zero:

KeqT^kgKaqT,- 1

2.303R

Except when AHRX is close to zero at 25 °C, the Van't 

Hoff equation allows the estimation of K^ values from 0 

to 50 °C, which are sufficiently accurate for most 

purposes.

Standard enthalpies of reaction are computed in a 

manner analogous to AGRX :

f products

AHRX =(dAH°D +eAH^-(bAH°B +cAH°c).

ACTIVITIES AND ACTIVITY COEFFICIENTS

It can be shown experimentally that an ionic compound is generally more soluble in an electrolyte solution 

than in pure water, provided the electrolyte does not consist of the same ions as the compound. This effect is the 

result of interactions between oppositely charged ions in 

solution, which cause an apparent shift in the chemical 

equilibria involved. For ,dilute solutions, the magnitude 

of this interaction is dependent on the ionic strength (I) 

of the solution, which provides a measure of the molal 

concentrations (m.) and charges (ZJ of the respective ions 

in solution:
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C75

To quantitatively describe the effects of ionic strength 

on equilibria, the concentrations of species in solution 

are adjusted by activity coefficients (7;) to yield the effective concentrations, or activities, of the species:

ai =7imi .

Activity coefficients for pure solids and liquids, and 

for dissolved species in infinitely dilute solutions, are 

equal to unity. As the ionic strength of the soluton 

increases, the dissolved ionic species lose some of their 

effectiveness, and the activity coefficient decreases. Activity coefficients may be calculated for individual ions 

in dilute solutions (I < 0.05) (Garrels and Christ, 1965) 

through the use of the Debye-Hiickel equation:

log 7;=

l+a.BT

or the Davies equation:

where A and B are constants dependent on the temperature of the solution, and a; is the experimentally determined ion size parameter (Kielland, 1937).

For gases at 25 °C and at pressures of less than 40 

atmospheres, the activity is essentially equal to the partial pressure of the gas:

COMPLEX EQUILIBRIA

For simple ions in solution, activities may be calculated from their respective activity coefficients and 

molalities. The activities of complex ions and ion pairs 

are, in turn, related to the activities of simple ions 

through mass action expressions; molalities of these 

complex ions and ion pairs may be calculated from their 

respective activity expressions. Conservation of mass 

and electro-neutrality constraints provide two additional 

expressions, allowing the set of simultaneous equations 

to be solved for the activities and molalities of all species 

in solution.

For instance, consider the dissolution of gypsum in 

pure water. The chemical reactions required to describe 

this system are:

H2O=H+ +OH- ,

CaSO4 -2H2O(c) =Ca+2 +SO4-2 +2H2O, 

CaSO4°=Ca+2+SO4- 2 , 

CaOH+=Ca+2 +OH- , 

HSO4 -=H++SO4 - 2 .

Five mass action equations can be formulated to describe these chemical reactions in terms of activities, and 

the molalities of the aqueous species can be related to 

these activities by seven activity coefficients, provided 

by the ionic strength and Debye-Hiickel equations. At 

this point, there are seven unknown molalities, seven 

unknown activities, seven unknown activity coefficients, 

and the unknown ionic strength. However, there are only 

five mass action equations, seven activity equations, 

seven Debye-Hiickel equations, and the ionic strength 

equation. To solve the problem, two additional equations 

are required to equal the number of unknowns. These additonal equations are provided by the mass balance and 

charge balance constraints. Because the solution is initially devoid of both calcium and sulfate, and because 

the independent addition of either calcium or sulfate during the dissolution of gypsum would result in the accumulation of a net electrical charge in the solution, the 

following mass balance:

m 0 +m +=m CaSO" CaOH + S

0 CaSO, HSO.

and charge balance:

2m ,.,+m ,+m + =2m _,+m _+m _

Ca+2 CaOH + H + SO4 ' HSO4 OH

expressions must be valid. Thus, the 22 independent 

equations may be solved iteratively to yield values for 

the 22 unknowns.

This example demonstrates that the set of simultaneous equations required to solve for the unknowns 

in even a simple natural solution rapidly becomes too 

large to calculate by hand. As a result, computer programs such as WATEQ2 (Ball and others, 1980) and 

PHREEQE (Parkhurst and others, 1980) have been 

developed to solve the necessary equations to describe 

complex equilibria in natural aqueous systems.

SATURATION INDICES

The activities of aqueous chemical species may be 

used to calculate the saturation state of a mineral in that 

solution. If a solution is undersaturated with respect to a
C76 REGIONAL AQUIFER-SYSTEM ANALYSIS

mineral, the mineral will tend to dissolve; if a solution is 

oversaturated with respect to a mineral, the mineral will 

tend to precipitate. Saturation with respect to a mineral 

indicates that the solution and the mineral are in 

equilibrium. It is important to realize, however, that 

equilibrium calculations indicate the direction in which a 

particular reaction will tend to proceed, but not the rate 

at which it will proceed.

The dissolution of calcite may be presented by the 

reaction:

CaCO3(c) +H+ =Ca+2+HCO-

(calcite)

so the corresponding mass action expression would be:

Ca 1 HCO:

equilibrium

Substitution of calculated activities into the above mass 

action expression will yield an ion activity product 

(IAP):

IAP=

calculated

Comparison of the calculated IAP value with the 

equilibrium K value will allow a determination of the sp

saturation state of a mineral. The saturation index (SI) 

relates measured compositions to the equilibrium state, 

and is defined as:

SI=log IAP

K

An SI equal to zero indicates saturation with respect to 

that mineral phase because the observed conditions are 

equivalent to equilibrium conditions. Similarly, examination of the above equations will reveal that if an SI value 

is greater than zero, the mineral is oversaturated, and 

that an SI value less than zero indicates 

undersaturation.

Knowledgeable interpretation of the saturation indices 

generated from complete and precise chemical analyses 

will, therefore, yield important information regarding 

the degree of equilibrium (or disequilibrium) attained 

between aquifer mineral phases and the aqueous solution, and will suggest which solid phases may control 

aqueous chemistry. However, caution must be exercised 

in the evaluation of plausible controlling phases for a 

particular system. As discussed previously, thermodynamic equilibrium calculations do not indicate the

rate at which a system will approach equilibrium. Thus, 

oversaturation with respect to a particular phase does 

not directly imply that the phase will precipitate from 

solution at a measurable rate.

PHASE DIAGRAMS

Phase diagrams were also utilized to determine which 

aquifer minerals are stable in a given aqueous solution. 

Aqueous dissolution reactions for pairs of mineral 

phases may be combined and balanced to yield an 

expression relating the stabilities of the two phases to 

the activities of their aqueous dissolution products. Solution of these equations in terms of two of three 

parameters ion activities, gas partial pressures, or oxidation potentials allows the stability boundary for 

that mineral pair to be plotted as a function of variation 

in the two parameters chosen. Activities of species not 

represented by either the ordinate or abscissa are fixed 

at values approximating solution composition. Superposition of aqueous compositions on such a diagram 

reveals the stable mineral or minerals in that water. For 

further discussion of such diagrams, or of solutionmineral equilibria in general, the reader is directed to 

Garrels and Christ (1965) and Stumm and Morgan 

(1981).

AQUEOUS SOLUBILITY OF GASES

The equilibrium activity of a particular gas in water 

(a jj is related by its Henry's Law solubility constant 

(Kgas) to the partial pressure of that gas (Pgas) in an adjacent gas phase by:

gas

Thus, knowing the atmospheric fractional abundance (N) 

of a gas and the total atmospheric pressure (Patm) at the 

location of interest, one can calculate the equilibrium activity of that gas in solution:

_ gas

gas vrp ^ A atm

Because Henry's Law solubility constants often vary 

with temperature, the aqueous concentration of a relatively inert gas, such as argon, krypton, or xenon, can 

reveal the temperature of ground-water recharge, assuming that the water has neither gained nor lost the gas of 

interest after contact with the unsaturated zone (Mazor,
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C77

1972). Additions of argon due to the radioactive decay of 

40K are generally insignificant for time intervals less 

than several million years.

OXIDATION-REDUCTION

Because the mobility of most elements exhibiting 

multiple oxidation states in natural systems is highly 

dependent on the oxidation potential (Eh) of the system, 

and because the oxidation of organic carbon provides an 

important driving force for many geochemical reactions, 

the recognition of redox zones and the estimation of Eh 

are very important in ground-water modeling studies.

The oxidation potential of a ground water is generally 

observed to decrease as the water flows from upland recharge areas to lowland discharge areas under confined 

conditions. Champ and others (1979) discussed such 

variations in Eh and the corresponding variations in concentrations of elements with variable oxidation states, 

including oxygen, nitrogen, iron, manganese, sulfur, and 

carbon. They concluded that a sequence of oxidationreduction reactions, based on thermodynamic principles 

and microbial catalysis, can account for the variations in 

oxidation potential. Their sequence of reactions defines 

three redox zones in ground-water flow systems that are 

described as: (1) oxygen-nitrate; (2) iron-manganese; and 

(3) sulfide. Ground water entering the recharge area of a 

confined aquifer system is oxidizing, containing both 

dissolved oxygen and nitrate, organic carbon present in 

the aquifer serves as the reductant, being catalytically 

oxidized to CO2 by micro-organisms that utilize these 

redox reactions as a source of energy for their metabolic 

requirements. The reductant, organic carbon with an oxidation sate of zero, is represented simplistically as 

CH2O in the following reactions. The pertinent 

geochemical reactions have been sequentially numbered 

for ease of reader reference.

Once the oxygen and nitrate have been microbially 

reduced by organic carbon:

02(g) +CH20=C02(g)+H20 (1)

4NO3-+5CH2O+4H+ =2N2(g)+5CO2(g) +7H2O (2) 

NO3-+2CH2O+2H+ =NH4 + +2CO2(g)+H2O (3)

the more reducing iron-manganese zone is encountered. 

Fe(III) and Mn(III)-Mn(IV) oxyhydroxides, stable in 

the oxygen-nitrate zone, are reduced to Fe(II) and 

Mn(II):

4Fe(OH)q(am) +CH20+8H+=4Fe+2 +CO,((r> +llH20, (4) '3(am| 2(g)

2MnO2+CH2O+4H+ =2Mn+2+CO2(g)+3H2O. (5)

After the reduction of iron and manganese, the ground 

water becomes sufficiently reducing to allow the onset of 

microbial sulfate reduction:

SO4- 2 +2CH2O+H+=HS-+2CO2(g) +2H2O, (6)

followed by acetate fermentation or CO2 reduction to 

form methane:

CH3COOH=CH4(g)+C02(g)

co2(g)+H2(g)=CH4(g)+2H2o.

' (7) 

(8)

The hydrogen required to reduce CO2 is derived from 

organic compounds present in the aquifer.

In some cases compounds other than organic carbon 

may act as reductants in the presence of oxidized species. Sulfide species originating from the dissolution of 

pyrite [FeS2(c)] can act to reduce both oxygen:

FeS2(c) +(7/2)02(g) +H20=Fe+2+2SO-2 +2H+ (9) 

and ferric iron: 

FeSw l + 14Fe+3 +8H9O=15Fe+2 +2SOr2 +16H+. (10)

£\Cf £ **

(pyrite)

Reduced gases such as methane may also serve as 

reductants on exposure to more oxidized species:

CH4(g)+202(g)=H2C03°+H20.

(11)

In this study, Eh was calculated using the appropriate 

oxidation-reduction reactions in conjunction with the 

Nernst-Peters equation:

Eh=EO+ 00591 

n

, (reduced)

where:

E°= standard electrode potential, for the combined 

half reactions, in volts,

n= number of moles of electrons transferred,

(oxidized) ,, ,. f ,,

- -=the mass action expression for the com-

(reduced)

bined half reactions.

For the oxygen-nitrate zone, the appropriate reaction involves the O2-H2O redox couple:
C78 REGIONAL AQUIFER-SYSTEM ANALYSIS

O2(gl +4H++4e-=2H2O ; E°=1.229 volts, (12) 

Eh=1.229-0.0592pH-0.0148 log (aH 0)2/(P0 ).

The Fe(II)-Fedll) couple (Doyle, 1968a; Whittemore 

and Langmuir, 1975) was used to calculate Eh when 

neither O2 nor H2S were present, as in many of the 

Judith Basin samples. Amorphous ferric hydroxide 

[Fe(OH)3(am)] and siderite [FeCO3(cl] are known to coexist 

in the Kootenai aquifers, and preliminary WATEQ2 

computations indicated that both of these phases were 

in equilibrium with sampled waters. Assuming that both 

Fe(OH)3(am) and siderite are in equilibrium with the solution, deciding on a pKsp range of 37 to 39 for the 

Fe(OH)3(am) dissolution reaction (Langmuir and Whittemore, 1971), and assuming the presence of a highly 

soluble amorphous siderite, the following reactions may 

be used to calculate Eh:

Fe(OH)3(am =Fe+3 +30H-; pK,^,^-37-39 , (13)

FeC03(c =Fe+2 +C03-2 ; pKFeC0^10.22, (14)

Fe+3 +e~=Fe+2 ; E°=0.771 volts. (15)

Adding reactions and substituting into the Nernst equation, the Eh expression becomes:

Eh=0.771+0.0592(pKFeCO -PKFe(OH), -3 log(a .

3(am) ° OH

1og(aco_ 2 )). (16)

WATEQ2 was used to calculate a _ and a _,, which

^ OH CO3 2

were in turn substituted into the above expression to 

evaluate the Eh of these waters. Saturation indices for 

siderite and Fe(OH)3(am), derived using the calculated Eh 

values and the WATEQ2 crystalline siderite stability 

constant, indicate equilibrium between these phases and 

all solutions for which this calculation was performed. 

The precision of the estimates, based on the variable 

stability of Fe(OH)3|am), was calculated to be ±60 millivolts of the intermediate (pKFe(OH) =38) value.

Finally, when measurable sulfide was present in the 

sampled waters, the SO~2 -H2S couple was used to calculate Eh:

SO4- 2 +10H+ +8e-=H2S°+4H2O; E°=0.301 volts , (16) 

Eh=0.301-0.0740pH-0.00740 log(aH s)(aH

Although recent evidence suggests that electroactive 

polysulfide (S~2 ) and thiosulfate (S2O~2 ) ions in equilibrium with dissolved sulfide control Eh in waters 

where these species are present (Boulegue and Michard, 

1979), the use of the more readily determined SO~2 -H2S

couple will result in minimal errors of approximately 50 

millivolts.

MINERAL DISSOLUTION AND WEATHERING

The CO2 produced during the microbial oxidation of 

organic matter provides an important driving mechanism for many important mineral dissolution and 

weathering reactions. Carbon dioxide dissolves, increases total inorganic carbon (CT) concentrations in the 

aquifer, and decreases pH through dissociation of the 

carbonic acid (H2CO^) formed:

C02(g) +H20=H2CO«,

H2CO3) =HCO3-+H+,

HCO-=COr2 +H+.

(17)

(18)

(19)

Although some mineral dissolution reactions can proceed without the consumption of acidity:

CaSO4 -2H2O(c) =Ca+2 +SO4- 2 +2H2O, (20)

(gypsum)

the acidity produced by the dissociation of carbonic acid 

provides a driving force for the weathering and dissolution of many other groups of minerals, including 

carbonates:

CaCO3(c) +H+=Ca+2 +HCO3-, (21)

(calcitel

CaMg(CO3 )2(c) +2H+=Ca+2 +Mg+2 +2HCO3- , (22)

(dolomite)

CaMg(C03 )2(c( +2H+=CaC03(c| +Mg+2 +HC03- , (23)

(dolomite) (calcite)

and aluminosilicates: 

K06 Mg025 Al25Si35010(OH)2(c) +3.15H20+l.lH^

(illite)

1.15Al2Si2O5 (OH)4(cf0.6K+ +0.25Mg+2+1.2H4SiO«, (24)

(kaolinite)

3Na0 33Al233Si367010 (OH)2(c) +H-

(Na-montmorillonite)

3.5Al2Si2O5 (OH)4(c| +Na++4H4SiO4 °,

(kaolinite)

3KAl2Si3O10(OH)2(c) +3H2O+2H+ =

(muscovite)

(25)
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C79

3Al2Si205(OH)4(c| +3K+

(kaolinite)

NaAlSi3O8(c) +6.1H2O+l.lH+ =

lalbite)

(26)

(allophanel

( +Na+ +2.6H4 SiO4 ° (27)

The coalification of lignite and other carbonaceous 

materials also results in the generation of CO2 , as well as 

in the release of considerable quantities of CH4 

(Teichmuller and Teichmuller, 1979). Experimental work 

by Foster (1950) has demonstrated the importance of 

CO2 derived from lignite, in combination with calcite 

dissolution and cation exchange, in producing the highbicarbonate waters of the Atlantic Coastal Plain.

ION EXCHANGE

Fine-grained sediments exhibit the capability of exchanging ions adsorbed on electrically charged particle 

surfaces. The net positive or negative charge on the particle is the result of ionic substitutions, imperfections, 

and broken bonds within the crystal lattice, or 

association-dissociation reactions that occur on the particle surfaces. To maintain electrical neutrality, ions of opposite charge, known as counter ions, are attracted to the 

particle and form adsorbed layers immediately surrounding the particle. These ion-exchange processes are most 

important for colloidal-sized particles, because of their 

relatively high charge-to-size ratio.

In natural systems, clay minerals form an important 

class of ion exchangers. The ion-exchange capacity of 

these layered aluminosilicates, which include the 

kaolinite, smectite, illite, chlorite, and vermiculite 

groups, results from imperfections and isomorphous 

substitutions in the crystal lattice, which generally 

create a net negative charge. In the kaolinite group, 

broken bonds at crystal plate corners and edges are the 

chief cause of the negative charge on the particles.

The oxide minerals, including silica and iron and manganese oxides, form a second important class of exchangers. The surfaces of these minerals contain incompletely coordinated metal and silicon cations at low 

pH, and an excess of hydroxyl or oxygen groups at high 

pH. The dissociation of water at these charged sites 

results in a charged surface layer surrounding the particle, the nature of which is dependent on solution pH. 

Positive surfaces prevail in low pH solutions; negative 

surfaces are formed in high pH solutions. Neutrality is 

attained at some intermediate pH, known as the zero 

point of charge (pHzpc), so the oxide has a tendency to adsorb major anions below the pHzpc, and major cations 

above.

As defined by van Olphen (1963), the cation exchange 

capacity (CEC) of a material is the excess of counter ions 

that can be reversibly exchanged for other cations in 

solution. Cation exchange capacity is usually expressed 

as the number of equivalents of cations that can be exchanged by 100 g dry weight of a material. The CEC's of 

several natural exchangers, as reported by Grim (1968) 

and Donald Langmuir, (Colorado School of Mines, oral 

commun., 1980), are presented in table 32. Clay exchangers generally have a greater affinity for divalent 

than monovalent cations due to the greater charge density of the former (Wiklander, 1964). Similarly, within 

groups IA and IIA of the periodic table, the affinity for 

adsorption by clays generally decreases with increasing 

hydrated ionic radii (Eisenman, 1962), as given by the 

Hoffmeister series:

Cs+ >K+>Na+>Li+ 

Ba+2 >Sr+2 >Ca+2 >Mg+2 .

The thermodynamics of exchange reactions can be described in terms of mass action expressions that relate 

the activities of cations in solution and on the exchanger 

to a thermodynamic exchange constant, Kex. Thus, for 

the following absorption-desorption reactions:

Ca+2+Na2EX=2Na++CaEX, 

Mg+2+Na2EX=2Na+2+MgEX,

(28)

(29)

where EX represents the negatively charged exchange 

site, the corresponding mass action expressions are 

given by:

KNaCa=

(a Na++ )2

and:

Na"1

where aN EX, * anc^ aMMEx represent the activities of gEx

these cations on the surface of the exchanger.

Because exchange reactions occur at a relatively fast 

rate, cation concentrations in ground water can be expected to be in equilibrium with the exchange surface. 

However, according to Freeze and Cherry (1979), many 

millions of pore volumes of water may have to pass 

through an aquifer before the ratios of absorbed cations 

completely adjust to the cationic composition of the input waters. This process may require time periods on the 

order of millions of years to complete, because of the 

relatively low concentration of cations in the input water
C80 REGIONAL AQUIFER-SYSTEM ANALYSIS

Material

CEC

(milliequivalents per 

100 grams at pH 7)

Kaolinite 

Glauconite 

Hike, Chloritc 

Smectites (mimtmnrillonite) 

Vcrmiculite 

Mn(IV) and Fe(III) oxyhydroxides

3-15

11-20

10-40

80-150

100-150

100-740

compared to the large surface area of the clay particles. 

Clays in strata deposited in marine or transitional environments are initially saturated with exchangeable 

sodium ions, the prevalent cation in seawater. These ions 

can be systematically desorbed by calcium and magnesium ions derived from mineral dissolution and present in infiltrating ground waters. The result is that, until exchanger compositions completely adjust to the cationic composition of infiltrating ground waters, a progressive decrease in aqueous calcium and magnesium 

concentrations, accompanied by a concurrent increase in 

sodium concentration, will occur as water flows through 

the aquifer.

ENVIRONMENTAL ISOTOPES

STABLE ISOTOPES

Many elements exhibit natural variations in their 

isotopic compositions because of the fractionation of isotopes during phase transitions and chemical and biological reactions. These isotopic fractionations are due to 

slight differences in the physical and chemical properties 

of different isotopes of an element, and are directly proportional to relative differences in their masses. As a 

result, isotopic fractionation of this type is most pronounced among elements with atomic masses less than 

or equal to that of calcium. The hydrologically significant elements considered in this report are hydrogen, 

carbon, oxygen, and sulfur.

The magnitude of isotopic fractionation associated 

with a particular process is indicated by the fractionation factor (a):

a=RA/RB

where RA and RB are the appropriate equilibrium isotopic ratios in phases A and B. Fractionation factors

may be calculated by means of statistical mechanics or 

measured experimentally; they generally approach unity 

with increasing temperature.

Isotopic composition is expressed in 5(Del) notation as 

per mil differences in the measured isotopic ratios of 

sample and standard. For example, the relationship for 

carbon isotopes is:

5 13C=

3C/12C)

sample

3C/12C), -1

standard

X1000

The carbon isotope ratios may be replaced by 2H/1 H 

(D/H), 18O/16O, or 34S/32S ratios to obtain the corresponding 5 values for these other elements. Typical standards 

used in evaluating isotopic composition include the 

SMOW (standard mean ocean water, for 518O and 5D) 

(Craig, 1961a), PDB-1 (CO2 gas produced from a belemnite of the Cretaceous Peedee Formation for 613C) (Craig, 

1957) and Canon Diablo troilite standards (troilite from 

the Canon Diablo meteorite, for 534S) (Jenson and Nakai, 

1963).

OXYGEN AND HYDROGEN

Natural variations in the 518O and 5D values of water occur from differences in the vapor pressures of waters 

containing different isotopes of oxygen and hydrogen. 

Water containing lighter isotopes will exhibit a significantly greater vapor pressure than water containing 

heavier isotopes. As a result, vapor formed from liquid 

water will be depleted (decreased 5) in 18O and D, and the 

remaining liquid will be enriched (increased 5) in these 

heavier isotopes. Because of this fractionation, 518O and 

5D vary characteristically as a function of latitude, elevation, condensation temperature, and distance inland, and 

can, therefore, be used as indicators of ground-water 

source and temperature of the recharge waters. Deviations from expected meteoric water compositions (Craig, 

1961b; Gat, 1980) in natural systems are indicative of 

evaporation and transpiration, mixing with nonmeteoric 

waters, or of isotopic exchange with the hydroxyl groups 

of silicate minerals at elevated temperatures.

CARBON

Stable carbon isotopes provide an extremely useful 

tool in the study of ground-water chemistry, aiding in 

the determination of controlling chemical reactions and 

mass transfer. Carbon isotopic data places an additional 

constraint on proposed mass transfer models, in that a 

geochemical model must correctly predict both the 

chemical and isotopic composition of a system.
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C81

In some instances, the relative importance of chemical 

control of ground-water chemistry by carbonates or 

aluminosilicates may be evaluated using 613C data. In 

temperate climates, plants use the Calvin photosynthetic cycle, which produces organic carbon with 6 13C 

values of -25±5 (Rightmire and Hanshaw, 1973). CO2 

produced in the unsaturated zone by the oxidation of 

this type of plant material should also have 613C values 

of approximately 25. Ground water in the unsaturated 

zone will exhibit 613C values representative of equilibration between dissolved bicarbonate, the predominant 

carbonate species between pH 6.4 and 10.3, and an 

essentially infinite gaseous reservoir of CO 2 . 

Equilibrium fractionation of 13C between HCO~ and

k^2<gi' aHco~-co.,, was reported by Deines and others 

(1974) to be 1.0092 at 10 °C, leading to a 9.2 per mil 13C 

enrichment in 6 13C _ over 6 13CCO , or a value

HCO 3 2igi

of 16 for bicarbonate in unsaturated zone waters. On 

reaching the water table, ground water is effectively 

isolated from the CO2 gas reservoir, and isotopic evolution proceeds under conditions that are closed with 

respect to exchange with that reservoir. Freshwater 

diagenetic carbonates, which are common in transitional 

sandstones, have 6 13C values of 4.8 ±5.1 (Keith and 

Weber, 1964), whereas CO2 derived from the oxidation of 

sedimentary organic carbon would have 613C values in 

the range of 28 ±4 (Craig, 1954). Congruent dissolution 

of calcite (reaction 21) or dolomite (reaction 22) driven by 

the oxidation of organic carbon (reactions 1-6) would 

liberate bicarbonate with a 6 13C value of 16, allowing 6 13C 

solution to remain approximately constant, as seen from 

mass balance considerations:

'final

m final

where m T is the molality of total inorganic carbon in 

solution.

Dissolution of aluminosilicates (reactions 24-27) also 

driven by the oxidation of organic carbon (reactions 1-6) 

would generate bicarbonate of the same isotopic composition as the organic carbon from which it was derived, 

causing the <5 13C of the evolving ground water to asymptotically approach a value of 28:

'con

m final

Using similar arguments, Wigley and others (1978) examined the isotopic evolution of a system controlled by 

dedolomitization, the incongruent dissolution of 

dolomite to form calcite (reaction 23). Again assuming an 

organic carbon CO2 source and a <5 13Ccarbonate value 

representative of either freshwater or marine 

(613C=-0.2±2.0) (Craig, 1954) carbonates, the authors 

calculated that solution 613C will asymptotically increase 

during closed (to a CO2 gas reservoir) system conditions, 

reaching a final value representative of equilibrium between input and output <5 13C values. Consideration of carbon mass balance and isotopic evolution should, 

therefore, allow for a distinction between the above 

types of reactions in a system where CO2 is being produced, but does not exist as a discrete gas phase.

Carbon isotopes may also be used to substantiate active methanogenesis (reactions 7 and 8) in aqueous 

sytems. Methane generated during these microbially 

mediated reactions is strongly depleted in 13C with 

respect to CO2 (Claypool and Kaplan, 1974), with 

aCOCH = 1.071 at 25°C (Bottinga, 1969). The activity of 

methanogenic bacteria would therefore be evident form 

the presence of methane and also from the 13C enrichment of dissolved carbonate. Migration of methane into 

an aquifer could not produce a similar increase in the 

6 13C of dissolved carbonate species, as Sackett and 

Chung (1979) have demonstrated that CO2 and CH4 do 

not attain isotopic equilibrium at low temperatures.

SULFUR

Sulfur isotopes can be used in geochemical studies to 

calculate the amount of sulfate reduction (reaction 6) 

that has occurred in an aquifer. Anaerobic bacteria, 

such as Desulfouibrio sp., reduce SO^ 2 to form H2S, 

which is depleted in 34S with respect to SO" 2 . The remaining sulfate is therefore enriched in 34S, and <534Ssulfate 

will increase accordingly. If both <5 34Ssulfide and <534Ssulfate 

values are known, calculated mass transfer in the sulfur 

system can be used to compute the expected final 

<534Ssulfate value, as demonstrated by Plummer and others 

(1983). Because <5 34Ssulfide determinations were unsuccessful in both study areas examined here, a discussion 

of this method is beyond the scope of this paper.

Sulfur isotopes have another use in hydrologic 

studies. Holser and Kaplan (1966) measured the 634S 

values in numerous samples of marine evaporate rocks 

of Phanerozoic age, and concluded that the 634S value of 

marine sulfate varied significantly during this period. 

Sulfate derived from marine sediments of one age will 

therefore differ from that derived from sediments of 

another age, so that differences in sulfur isotopic 

composition can be used to identify sulfate associated
C82 REGIONAL AQUIFER-SYSTEM ANALYSIS

with leakage between strata of different ages. However, 

considerable overlap exists between marine sulfates of 

different ages, so the age of the strata from which the 

sulfate was derived may not be unequivocally determined in many cases.

RADIOACTIVE ISOTOPES

TRITIUM

Tritium (3H) is a naturally occurring, radioactive isotope of hydrogen, which is produced in the upper atmosphere by the interaction of nitrogen with cosmic 

ray-generated neutrons (n):

13N+n=12C+3H (30)

The newly created tritium then combines with oxygen 

to form water, which is subsequently dispersed 

throughout the hydrosphere. As a result of the atmospheric testing of hydrogen bombs, naturally produced 

tritium has been far overshadowed during the past 25 

years by much higher levels of artificially produced 

tritium. These artificial tritium levels, which are between one and three orders of magnitude greater than 

pre-bomb levels, serve as a hydrologic tracer with which 

relatively young waters may be identified. Tritium concentrations on the order of 10 picocuries per liter (pCi/L) 

or less (2-4 tritium units or less; 1 TU=1 tritium atom 

per 1018 hydrogen atoms) are indicative of waters with 

ages of greater than 25 years, whereas greater concentrations indicate the presence of waters younger than 25 

years in age.

CARBON-14

The measurement and adjustment of carbon-14 ages 

of waters from confined aquifers remains the most suitable and widely used method of ground-water dating. 

Past attempts to use other radioactive isotopes, such as 

39Ar and 32Si (Oeschger and others, 1974) have met with 

limited success, although refined analytical techniques 

may allow the routine use of such dating schemes in 

future studies.

Carbon-14 is produced in the upper atmosphere by the 

interaction of cosmic ray-produced neutrons with stable 

isotopes of carbon, oxygen, and nitrogen. The most 

significant of these is the reaction between slow cosmic 

ray-produced neutrons and nitrogen:

14N+7i=14C-HH. (31)

The 14C atoms are incorporated into molecules of carbon 

dioxide and then rapidly mixed throughout the atmosphere, attaining constant concentrations representative of steady-state equilibrium.

Carbon-14 dating is based on the principle that, upon 

isolation from the atmospheric 14C reservoir, the concentration of this unstable isotope begins to decrease from 

radioactive decay by beta emission to 14N:

(32)

In ground-water systems, 14C concentrations may also 

be altered by mixing or interaction with other carbon 

sources, so that the accuracy of a date is dependent on 

one's ability to successfully isolate the effects of radioactive decay from those of other processes that change 14C 

content. In the absence of such processes, the equation 

for determining time (t, in years) since isolation from the 

atmosphere based on the 14C activity is:

t,.

v"^

0.301 log

where:

t, = 14C half life, 5,730 years

A= measured 14C activity, in percent modern

carbon

AND = 14C activity to be expected in the absence of 

radioactive decay, in percent modern 

carbon.

Estimation of AND is dependent on quantitatively determining the magnitude of carbonate sources and sinks in 

the aquifer system. Under favorable conditions, this can 

be accomplished through the use of geochemical mass 

transfer models.

Wigley and others (1978) divided carbon mass transfer 

into three stages for the purpose of 14C dating:

1. Evolution in the unsaturated zone, where infiltrating ground water is open to isotopic exchange with an effectively infinite, gaseous CO2 

reservoir.

2. Congruent dissolution of carbonates (reactions 21 

and 22) in the saturated zone, where ground 

water is assumed to be closed to exchange with a 

CO2 gas reservoir.

3. Incongruent dissolution of carbonates (reaction 

23) in the saturated zone, with closed system 

conditions continuing.

The effects of these three stages are dependent on isotopic fractionation between gas and aqueous phases, carbon mass transfer, and isotopic fractionation between
GEOCHEMISTRY OF GROUND-WATER IN TWO SANDSTONE AQUIFER SYSTEMS C83

aqueous and solid phases. Substitution of expressions 

representing the effects of each of the three evolutionary 

stages into the above decay equation yields:

t= 19035

100 613C2 -<513C. 

log^+log ( - ) +

A d^C

log(

/36 13C-613C.+

where:

d^C^ solution <5 13C at the end of stage 1, 

613C2 = solution 513C at the end of stage 2, 

<5 13C,= dissolving carbonate <5 13C,

|3= composite parameter representing reaction

progress, 

eps= additive fraction factor between solution

and precipitate, 

F= ratio of carbon input to output rate.

For a detailed derivation of this expression and examples of its use, the reader is directed to Wigley and 

others (1978). Application of the above expression to 

analytical data and computed mass transfer allows an 

estimation of ground-water age, which in turn can be 

used to derive flow rates and aquifer constants.

GEOCHEMICAL COMPUTER MODELS

WATEQ2

The computer program WATEQ2 (Ball and others, 

1979; 1980) was used to compute the equilibrium distribution of aqueous species and then calculate mineral 

saturation indices for the sampled ground waters. This 

PL/1 code is a revised and expanded version of 

WATEQ, the computerized aqueous chemical model of 

Truesdell and Jones (1974). The method of calculation 

involves solution of mass action and mass-balance equations through an iterative back-substitution on cation 

and anion mass-balance equations. Options are provided for calculation of activity coefficients using either 

the Debye-Hiickel or Davies equations.

WATEQ2 allows for consideration of ion association 

and solubility equilibria for several trace metals (Ag, 

As, Cd, Cu, Mn, Ni, Pb, and Zn), solubility equilibria for 

various metastable solids, and polysulfide speciation. 

Calculations of redox potential from various couples and 

from propagated standard deviations for many ion activity products are also provided. The program utilizes 

an extensive, low-temperature thermodynamic data 

base, which has been updated by using recent compilations of critically evaluated data (Ball, 1979; 1980).

Input data consist of laboratory determinations of 

total concentrations of major, minor, and trace 

elements, and field-measured temperature, alkalinity, 

and pH values. Eh may be determined in the field using 

a platinum electrode (Langmuir, 1971), or may be calculated from any of several appropriate redox couples 

(reactions 12-16). The significance of the calculated 

saturation indices is determined by the propagation of 

analytical errors, with saturation being considered any 

value of SI within the 95 percent confidence interval of 

zero. For many minerals, 95 percent confidence lies 

within several tenths of zero; for others, saturation may 

be implied at greater than one unit deviation from zero 

because of the poor precision of analytical results near 

the detection limit.

PHREEQE

Reaction simulation and mass transfer modeling of 

aquifer geochemistry was accomplished using PHREEQE (Parkhurst and others, 1980), a FORTRAN computer 

program designed to simulate chemical reactions. 

PHREEQE utilizes an ion-pairing aqueous-speciation 

model similar to that in WATEQ2 to simulate aqueoussolution chemistry. It can equilibrate an initial solution 

with one or more mineral or gas phases, calculating the 

final equilibrium solution composition, Eh and pH, as well 

as the associated mass transfer between solid, aqueous, 

and gaseous phases. The code also allows the user to add 

net stoichiometric reac- tions to a solution, to mix solutions of different composi- tion (constant volume), or to 

titrate one solution with another (additive volume).

PHREEQE was used to simulate changes in groundwater chemistry associated with equilibration of an initial rainwater solution with ion exchangers and various 

mineral phases. A computational problem arises in the 

attempt to mathematically equilibrate a given solution 

with Na/Ca and Na/Mg exchange, and with the mineral 

phases calcite and dolomite, simultaneously. Both ionexchange and mineral-equilibrium equations specify a 

a +2/a +2 ratio, indicating that the four equations

Ca Mg

simulating these reactions are not independent, and a
C84 REGIONAL AQUIFER-SYSTEM ANALYSIS

unique algebraic solution is not available. In natural 

systems, an additional variable is provided by the surface composition of the exchanger, resulting in an additional equation. PHREEQE is currently unable to consider the variable surface composition of an exchanger; 

therefore, simultaneous equilibration with Na/Ca and 

Na/Mg exchange in combination with calcite and 

dolomite equilibration was accomplished in a stepwise 

manner. Solutions were first equilibrated with calcite 

and dolomite, followed by equilibration with ion exchange. This two-step procedure was repeated until the 

resulting solution composition had stabilized, and the 

saturation indices of calcite and dolomite were within 

analytical uncertainty of zero. Comparisons of this stepwise method and simultaneous mineral-ion exchange 

equilibration for calcite-ion exchange and dolomite-ion

exchange simulations result in similar mass transfers 

and final solution compositions via the two methods.

ISOTOPIC EVOLUTION EQUATIONS

A. U.S. Geological Survey FORTRAN computer program was used to calculate carbon isotopic evolution 

associated with simulated mass transfers. This code 

utilizes a finite difference method to solve the isotopic 

evolution equations of Wigley and others (1978), 

yielding final 613C and AND values from initial 613C and 

13C measurements, simulated carbonate mass transfer 

and solution composition. Fractionation factors have 

been calculated from the temperature-dependent functions of Deines and others (1974).

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