Geochemistry of Groundwater: Symposium Proceedings (Developments in Water Science)

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SYMPOSIUM ON GEOCHEMISTRY OF GROUNDWATER 26th INTERNATIONAL GEOLOGICAL CONGRESS, PARIS, 1980

DEVELOPMENTS I N WATER SCIENCE, 16 OTHER TITLES IN THIS SERIES 1

G. BUGLIARELLO AND F. GUNTER

COMPUTER SYSTEMS A N D WATER RESOURCES

2

H.L. GOLTERMAN

PHYSlOLOGlCAL LIMNOLOGY

3

Y . Y . HAIMES, W.A. H A L L AND H.T. FREEDMAN

MULTIOBJECTIVE OPTIMIZATION I N WATER RESOURCES SYSTEMS: THE SURROGATE WORTH TRADE-OFF-METHOD

4

J.J. FRIED

GROUNDWATER PO L L U T l ON

5

N. RAJARATNAM

TURBULENT JETS

6

D. STEPHENSON

PIPELINE DESIGN FOR WATER ENGINEERS

7

v.

HALEK AND J. SVEC

GROUNDWATER HYDRAULICS

8

J. BALEK

HYDROLOGY A N D WATER RESOURCES I N TROPICAL AFRICA

9

T.A. McMAHON AND R.G. MElN

RESERVOIR CAPACITY A N D Y I E L D

10 G. KOVACS SEEPAGE H Y DR A U L l CS

1 1 W.H. GRAF AND C.H. MORTIMER (EDITORS) HYDRODYNAMICS OF LAKES: PROCEEDINGS OF A SYMPOSIUM

12-13 OCTOBER, 1978, LAUSANNE, SWITZERLAND

12 W. BACK AND D.A. STEPHENSON (EDITORS) CONTEMPORARY HYDROGEOLOGY: THE GEORGE BURKE M A X E Y MEMORIAL VOLUME

13 M.A. M A R I I ~ OAND J.N. LUTHIN SEEPAGE A N D GROUNDWATER

14 D. STEPHENSON STORMWATER HYDROLOGY A N D DRAINAGE

15 D. STEPHENSON PIPELINE DESIGN FOR WATER ENGINEERS (completely revised edition of Vol. 6 in the series)

SYMPOSIUM ON GEOCHEMISTRY 26th INTERNATIONAL GEOLOGICAL CONGRESS, PARIS, 1980

EDITED BY

WILLIAM BACK 432, National Center, U.S. Geological Survey, Reston, VA 22092 (U.S.A.) and

RENE

LETOLE

Dbpartement de Geologie Dynamique, UniversitbPierre et Marie Curie, 75230 Paris 05 (France)

Reprinted from Journalof Hydrology, Vol. 54, No. 113 (1981)

ELSEVIER SCIENTIFIC PUBLISHING COMPANY Amsterdam-Oxford-New York 1982

ELSEVIER SCIENTIFIC PUBLISHING COMPANY Molenwerf 1, P.O. Box 21 1 , 1000 A E Amsterdam, The Netherlands Distributors for the United States a n d Canada: ELSEVIER/NORTH-HOLLAND I N C 52, Vanderbilt Avenue New York, N.Y. 10017

Symposium oil (;eocliernisi;ry of Groundwater (1980 : Paris, im'rance ) Symposium ,n Geochemistry of Groundwater. (Developments in water science ; 16) "Reprinted from J:>urnal 3i hydrology, vol.

54,

no. 1/3 (1981)" Includes i n d e x . 1. Water chemistry--Congresses. 2. Hydrogeology-Congresses. 3 . Geochemistry--Congresses. I. Back, . 11. LCtolie, Renk. III. InternaWilliam, 1925tional G e o l o g i c a l C m g r e s s (26th : 1980 : Paris, France) IV. Journal of hydrology. V o l . 54. no. 1~/3. V. Series.

tiB855 .S95

l9?0 ISBN 0 - 4 4 1 i - 4 ~ 0 j b - j

551.119

81-15087 AfiCRZ

ISBN : 0-44442036-3 (VOl. 16)

ISBN:0-44441669-2 (Series)

0 Elsevier Scientific Publishing Company, 1982

A l l rights reserved. N o part of this publication may be reproduced, stored in a retrieval system or transmitted in any form or b y any means, electronic, mechanical, photocopying, recording or otherwise, without the prior written permission of the publisher, Elsevier Scientific Publishing Company, P.O. Box 330, 1000 A H Amsterdam, The Netherlands Printed in The Netherlands

Preface This compilation of articles is based on invited and unsolicited papers that were presented at the Symposium “Geochemistry of Groundwater and Aquifers” sponsored by the International Association of Hydrogeologists and the Geochemistry Section of the 26th International Geological Congress. The symposium was organized at the invitation and request of Dr. G. Castany, President, Hydrogeology Section, International Geological Congress, Dr. A. Burger, Institute of Hydrogeology, Neuchatel, and Dr. P.E. LaMoreaux, President, International Association of Hydrogeologists. Papers prepared for such international symposia always provide a significant, albeit incomplete, synoptic perspective of the state-of-the-art for those topics comprising the theme of the symposium, The papers often reflect the collective judgment of established scientists in identifying certain topics of nationalistic concern, deficiencies of knowledge and areas of desirable research. Included are papers that demonstrate a critical need still remains to provide an adequate description and explanation for the spatial and stratigraphic distribution of chemical constituents and their relation to the geological framework; some demonstrate the desirability and advantages of including isotopes in such geochemical studies of aquifers. Several papers demonstrate various levels of complexity of predictive models being developed to determine the concentration of chemical constituents at selected points within a specified time frame. Concepts and models are included here to determine dispersivity from regional geochemical processes; to emphasize hydrologic significance of chemical character of water that ranges from vadose flow to deep artesian units, and to elucidate the geochemical evolution of groundwater in large sedimentary basins, glacial aquifers and consolidated sandstone. A surprisingly large percentage of papers deal with brines of deep aquifers. This, no doubt, reflects the increasing awareness of the important role that groundwater plays in processes and activities such as generation of brines, production of geothermal energy, genesis of ore deposits, and deep-well waste injection programs. The wide range of problems of groundwater contamination and waste disposal, storage or containment are clearly exemplified by the last two papers, one of which discusses disposing of ash from power-generating plants - a problem that will expand as plants convert from high-priced oil to cheaper coal; and the last paper focuses on the increasing difficulty of selection of suitable waste-disposal sites as wastes continue to proliferate and available land continues to diminish.

viii

We want to express the appreciation of the editors and the authors t o all those scientists who so willingly reviewed and improved the manuscripts. We particularly acknowledge the editorial assistance of Laura Toran, U.S. Geological Survey. WILLIAM BACK 432, National Center U.S. Geological Survey Reston, VA 22092, U.S.A.

RENELETOLLE Dkpartement d e Gkologie Dynamique Universitk Pierre e t Marie Curie 75230 Paris 05, France

Contents Preface

....................................................

vii

I. Regional Relation o f Lithology and Chemical Character o f Water Factors of the chemical composition of seepage and groundwaters in the intertropical zone (West Africa) E.J. Roose (Paris, France) and F. Lelong (Orlians, France) . . . . . . . . . . . . . . . . 1 Geochemical and isotopic characteristics of spring and groundwater in the State of Sao Paulo, Brazil M. Szikszay, J.-M. Teissedre (Sao Paulo, Brazil), U. Barner (Tel Aviv, Israel) and E. Matsui (Piracicaba, Brazil) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 Repartition des elhments trace dans les eaux thermominirales (Minor-element distribution in mineral and thermal waters) J. Sarrot-Reynauld, J. Rochat et J. Dazy (Grenoble, France). . . . . . . . . . . . . . . 33

II. Geochemistry o f Brines and Deep Aquifers The origin and evolution of saline formation water, Lower Cretaceous carbonates, south-central Texas, U.S.A. L.S. Land (Austin, Texas, U.S.A.) and D.R. Prezbindowski (Tulsa, Okla, U.S.A.) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Dissolution of salt on the east flank of the Permian Basin in the southwestern U.S.A. K.S. Johnson (Norman, Okla., U.S.A.) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Patterns of groundwater salinity changes in a deep continental-oceanic transect off the southeastern Atlantic coast of the U.S.A. F.T. Manheim and C.K. Paul1 (Woods Hole, Mass., U.S.A.) . . . . . . . . . . . . . . . . Character of brines from the Belle Isle and Weeks Island salt mines, Louisiana, U.S.A. M.B. Kumar and J.D. Martinez (Baton Rouge, La., U.S.A.) . . . . . . . . . . . . . . . .

51 75 95 107

III. Istopes in Groundwater Sulfur and oxygen isotopes as tracers of the origin of sulfate in Lake Cr6teil (southeast of Paris, France) A. Chesterikoff, P. Lecolle, R. L6tolle and J.P. Carbonnel (Paris, France) . . . . . . 1 4 1 The Madrid Bain aquifer: preliminary isotopic reconnaissance F. Lopez Vera (Madrid, Spain), J.C. Lerman and A.B. Muller (Tucson, Ariz., U.S.A.) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1 5 1 Radiocarbon dating of groundwater of the aquifer confined in the Lower Triassic sandstones of the Lorraine region, France 167 B. Blavoux and Ph. Olive (Thonon-les-Bains,France) . . . . . . . . . . . . . . . . . . . . Uranium isotopes and 226Racontent in the deep groundwaters of the Tri-State region, U.S.A. J.B. Cowart (Tallahassee, Fla., U.S.A.) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 185

IV. Chemical Models o f Groundwater Systems Carbonate geochemistry of vadose water recharging limestone aquifers J. Thrailkill and T.L. Rob1 (Lexington, Ky., U.S.A.) . . . . . . . . . . . . . . . . . . . . 195 A geochemical method of determining dispersivity in regional groundwater systems W.W. Wood (Lubbock, Texas, U.S.A.) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 209 Flow-system controls of the chemical evolution of groundwater F.W. Schwartz, K. Muehlenbachs (Edmonton, Alta., Canada) and D.W. Chorley (Vancouver, B.C., Canada) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 225 Chemical evolution of groundwater in a drainage basin of Holocene age, east-central Alberta, Canada 245 E.I. Wallick (Edmonton, Alta., Canada) . . . . . . . . . . . . . . . . . . . . . . . . . . . .

X

The rate of flushing as a major factor in determining the chemistry of water in fossil aquifers in southern Israel A. Issar (Beer Sheva, Israel) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 285 Geochemical inputs for hydrological models of deep-lying sedimentary units: loss of mineral hydration water D.L. Graf and D.E. Anderson (Urbana, Ill., U.S.A.) . . . . . . . . . . . . . . . . . . . . . 297 H&t&ogen&te chimique et hydrologique des eaux souterraines d’un karst du HautJura neuchitelois, Suisse (Chemical and hydrological heterogeneity of groundwaters in the Upper “Jura Neuchitelois”, Switzerland) Y. Bouyer et B. Kubler (Neuchatel, Suisse) . . . . . . . . . . . . . . . . . . . . . . . . . . 315 V. Stressed Environments

Effect of leachage solutions from fly and bottom ash on groundwater quality D. A. Kopsick and E.E. Angino (Lawrence, Kans., U.S.A.) . . . . . . . . . . . . . . . . 341 Geological considerations in hazardous-waste disposal K. Cartwright, R.H. Gilkeson and T.M. Johnson (Champaign, Ill., U.S.A.) . . . . . . 357

1

FACTORS OF THE CHEMICAL COMPOSITION OF SEEPAGE AND GROUNDWATERS IN THE INTERTROPICAL ZONE (WEST AFRICA)

ERIC JEAN ROOSE and FRANCOIS LELONG

O.R.S.T.O.M., 75008 Paris (France) Laboratoire d e Giologie Appliquie, Domaine Uniuersitaire, 45046 Orlians C i d e x (France) (Accepted for publication February 20,1981)

ABSTRACT Roose, E.J. and Lelong, F., 1981. Factors of the chemical composition of seepage and groundwaters in the intertropical zone (West Africa). In: W. Back and R. LGtolle (Guest-Editors), Symposium on Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54: 1-22. In connection with a large research programme about the actual dynamics of ferrallitic and ferruginous soils of West Africa, 5000 samples of rainfall, throughfall, runoff, drainage and phreatic waters have been analysed during 4-11 years of field observations. Samples of eight stations, representative of different bioclimatic conditions (sub-Equatorial to pre-Sahelian), have been tested. The analysed parameters are: pH, resistivity, major cations and anions, total organic carbon and nitrogen, phosphorus, silica, aluminium and iron. The results show: (1)The slight influence of the bioclimatic differentiation on the mean chemical composition of the waters : all analysed waters are lightly mineralized (strong resistivity, total chemical charge generally lower than 100 mg/l), with an increasing mineralization from rainfall water to seepage water, at 2 m depth, but decreasing a t the water table level (except for Si and Na). (2) The marked variability of the amounts of dissolved chemical species compared to the seasons and the flow volumes. (3) The complexity of phenomena controlling the chemical composition of waters. In the soil layers, this composition would depend principally on biological and biochemical processes, in relation to the activity of organisms but at the level of phreatic waters the chemical composition would rather be controlled by physicochemical conditions (solution-mineral equilibria). INTRODUCTION

The chemical composition of groundwaters depends on many causes: rainwaters have a notable chemical charge, especially in marine salts (M. Schoeller, 1962) and in elements which originate from atmospheric dusts (silts, clays, pollen and bacteria) or from urban and industrial activity (Crozat, 1978; Dess&weDelepoule, 1978). Rain waters leach the greencover, carry away available constituents and are loaded with biophile elements, such as C, Mg, Ca and K (Duchaufour,

2

1977). The importance of this input (throughfall) is related t o the rainfall volume, the nature and density of greencover (Mathieu, 1972). In the soil layer many interactions occur: percolating waters rapidly acquire a chemical composition similar t o that of the groundwater; this fact has been confirmed in temperate regions for acid and calcareous soils (M. Schoeller, 1958; H. Schoeller, 1966). But the chemical charge of soil waters varies during the seasons and with the hydraulic potential : strongly bonded water in the microporous and matrix interstices is much more charged with products from mineral hydrolyses than free waters, which are renewed by rainfall waters and leach only a part of the chemical charge of the micropores (Vedy and Bruckert, 1979). Modern works emphasize the water-rock mineral interactions t o explain the chemical composition of underground water: Jacks (1973) shows that the major part of the ionic charge is originated from the rock-mineral dissolution; Tardy (1969) indicates that the mineralization of groundwater is related to the weathering type; BourriC. (1978) defines the respective proportion of major anions and cations in spring waters of the Vosges and the Massif Central, France, coming from atmospheric input and weathering reactions. In conclusion, the composition of groundwaters is often interpreted essentially in terms of solution-mineral equilibria. In connection with a large research programme about the actual dynamics of ferrallitic and ferruginous tropical soils of West Africa, runoff, seepage waters and phreatic groundwaters have been continuously sampled during several years, analysed and compared with rainfall waters and throughfall waters (Roose, 1980b). The chemical charge transported by soil and phreatic waters is in ionic and solid or colloidal form: this latter is very important and considerably influences the top soil profile differentiation. But in the present paper, the problem of the ionic charge of water will be treated only in view of precising the role of the soil layer towards the groundwater chemical composition.

METHODS

The experimental device in the field

Eight stations representative of natural soil-vegetation ecosystems of West Africa have been equipped with experimental devices -rain gauges, lysimeters, runoff and oblique drainage plots (Fig. 1)- allowing t o sample rainfall, rain water, throughfall, runoff, vertical and oblique drainage waters at each device. Phreatic waters have also been sampled in springs, near by the plots. The experimental device has been precisely described in recent works (Roose, 1978, 1980a). The eight stations (Fig. 2) are distributed from the southern, forested Ivory Coast, near Abidjan, t o the pre-Sahelian area near Ougadougou (Upper

3

Fig. 1. Experimental device at Divo, Ivory Coast. On a face of the trench, the gutters to sample the oblique drainage water can be seen.

Volta). These stations make it possible t o characterize different soiluegetation ecosystems, representative of a wide bioclimatic sequence, which extends from the zone of desaturated ferrallitic soils in the south, t o the zone of tropical fermginous soils in the north. Climate, vegetation, bedrock and geomorphology characteristics are given in Table I. During 4-11 successive years, the following samples were taken at each station and at each rainfall: (1) rainfall water and throughfall; (2) runoff water; (3) soil drainage waters at several depths (0.30-2 m), collected in lysimeters (vertical drainage) and in oblique drainage plots (oblique drainage); and (4)phreatic groundwaters, sampled in springs. On the whole, more than 5000 samples of different types of waters have been collected and stored prior t o be analysed in plastic bottles, which were previously washed with the water to be analysed, filled up t o the top, maintained in darkness and as soon as possible (one day t o six weeks) transported to the analytical laboratory of Adiopodoume (Ivory Coast).

Laboratory analysis The physicochemical parameters studied are: pH, temperature, resistivity, Ca, Mg, K, Na, SO4, C1, organic carbon, total nitrogen, NO,, NH4, total phosphorus, FeyA1 and S i 0 2 . The methods of analysis are those which are practiced in the O.R.S.T.O.M. Laboratory of Adiopodoume (Nalovic, 1968; GOUZY, 1973). The raw water is decanted, then filtered (folded fast filter paper Pratt Dumas@ No. 4 ) .

4

Fig. 2 . Localities of stations in West Africa (taken from Roose, 1978). Gonse : O.R.S.T.O.M.-C.T.F.T ;Gampela : C .T.F.T.-0.R.S .T.O .M. ;Saria : 0 .R.S.T.O.M.0.R.S.T.O.M.I.R.A.T.;Korhogo: O.R.S.T.O.M.;B0uake:I.R.A.T.-O.R.S.T.O.M.;Divo: I.F.C.C.; Azaguie : 0 .R.S.T.O.M.-I.R.F.A.; Anguededou : 0.R.S.T.O.M.-I .R.C.A.; Adiopodoume: O.R.S.T.O.M.; Agonkamey: O.R.S.T.O.M.; Ibadan: I.I.T.A. a = erosion site; = ERLO site; o = DV site.

.

TABLE I Ecological characteristics of the experimented stations

Climate rainfall (mm) PET* (mm) rainfall erosivity index (Wischmeiy, 1962) temperature ( C) Greencover

Landforms

grade of slope (%) length slope (m)

Adiopodoume

Anguedegou

*

Sub-Equatorial (two rainfall seasons)

2,150 1,250

2,100 1,300

1,7 50 1,314

1,550 1,280

1,200 1,300

1,260 26.2

1,000 26.2

885 26.2

825 26.0

512 26.1

4

Azaguie

wet dense forest evergreen

Divo

douake

-

semideciduous

-

Korhogo

Saria

transitional tropical 1,3 50

pure tropical

Gonse

1,660

830 1,885

860 1,905

676 27.0

450 28.0

430 28.1

wooded savannah Guinean Soudanian

bush savannah Soudano-Sahelian

strongly incised plateau

narrow hills

wide hills

wide hills

residual relief and long piedmont slopes

residual relief and very long piedmont slopes

1 3 1 11 .331?1?3'.1> I lr' 1 2 1 1'21 772 1 2 3 1?2 I ?

0

1

1

1 1 11

............................................ 1

1

1

1

1

I

1

1

I

1

. *

Y..*....*...,~....*......

75.

45,

IS.

0.0

30.

60.

135

105 90.

120

165

150

195 180

Volune ( l i t e r s )

Fig. 4 . Variation of calcium concentration (mg/l) in oblique drainage waters at the Adiopodoume station, as fraction of the volume of flow (liters).

at the level of the phreatic waters, except for SiOz and Na; these constituents seem to be less absorbed by plant roots than the others. The fact that maximum concentration values generally correspond to drainage waters is clearly shown in Fig. 6, giving the MPV-values in reference to the logarithm of the drainage volume (expressed in height of water). The plots of each station can be identified with abbreviations of localities. The drainage volume has been calculated from data of monthly rainfall, runoff and potential evapotranspiration, with reference to an available variable reservoir of water in the soil estimated a t the stations: maximum, 200 mm at Adiopodoume, minimum, 60 mm at Saria (Table 111). In Fig. 6 Adiopodoume and Anguedegou are plotted together, and Saria and Gonse in a similar manner.

9

6 0 0

50

N

0

.

+

40

v)

30 1

.1

. 20

.

-31 .21 1

.

.ll 10

1

11

1 1

1 11 1 1 1 1 1\ 11 11 11 1

+GL.1?

1

* 0

1 11

1121

rqr67457&117!>? 1 1 .! . ~ 5 2 2 1 ? . ! 7 ~ ~ 1 ~217 h I 121

.4?51

1

1

1

1

I3211

.12

0.

0

1

..........

1 1 1 1 1

1

1

1

1

Y

1

1

rz

h+.,..r,..

15.

0.0

.r.,.r.b.....o..b..~.-.o..~..~..*-

75.

45.

30.

60.

105 90.

165

135

120

150

195 180

Volume ( 1 1 ters)

Fig. 5. Variation of silica concentration (mg/l) in oblique drainage waters at the Adiopodoume station, as a function of the volume of flow (liters).

DISCUSSION

The first point of discussion is the influence of the bedrock nature on the chemical composition of waters. We can see in Table I1 and in Fig. 6 that waters in the chloritic schist or argillaceous sand areas do not differ from those in the granitic area. This is not surprizing, in spite of the well-known control of rock mineral on the groundwater composition (M. Schoeller, 1962; Tardy, 1969); in fact, in the whole bioclimatic sequence considered, the lateritization processes have destructed all the primary minerals, except quartz, and the weathering minerals are always the same: kaolinite, iron oxide and oxy-hydroxide, the stabilities of which in the surface are very great. The second point of discussion concerns the influence of flow volume. It

TABLE I1 Chemical characteristics of the different waters sampled at the eight experimental stations, West Africa - mean pH-values, electrical resistivities and concentrations of dissolved chemical species (MPV): major cations and anions, organic elements, nutrients and sesquioxides Station

Vegetation*

Rainfall

Through fall

Runoff

Drainage upper level

Spring lower level

pH:

Adiopodoume (R2) Anguededou Azaguie Divo Bouake Korhogo Saria (P7) Gonse

Resistivity

f. h. f.

f. s.f. S. S. S.

-

6.7

6.7 6.6 6.6 6.7 6.8 6.9 6.4 6.4

-

_.

-

6 .O

-

6.3 6.6 6.7 6.7 6.7 6.7 6.5 7 .o

6.8 6.5 6.8 7 .o 6.7 7 .O

-

-

-

6.6

5.3

-

6.1 6.9

(acm):

Adiopodoume (R2) Anguededou Azaguie Divo Bouake Korhogo Saria (P 7) Gonse

f.

h. f.

f. s.f. S.

S. S.

48,800 -

68,900 91,700 -

25,000 54,000 76,000 -20,000 -

(a)

28,700 32,300 34,000 18,100 27,700 46,900 76,540 31,800

27,200 21,800 23,000 9,500 32,600 32,000 24,700 27,700

17,900 19,600 22,500 6,200 25,600 23,200 -

25,900

38,100 44,600 21,800 -

SO4 : Adiopodoume (R2) Anguededou Azaguie

f. h. f.

1.2 -

2.5 -

3.1 8.8 3.5

5.1 12.6 5 .O

6.2 17.6 3.6

-

1.4

Divo Bouake Korhogo Saria (P7) Gonse

f.

s.f. S.

S.

1.0 1.6

1.7 1.7-1.9 (a)

S.

5.7 0.8 2.1 2.1 33

10.7 1.8 4 .O 3.4 3

2.8 3.5 1.2 3.2 1.1 1.3 0.8 0.7

4.4 5.2 2.8 10.4 1.6 2.5 2.3 0.4

6.6 8.4 3.7 6.4 0.8 2.5 0.8

28.2 14.7 12.5 18.0

1 4 .O 18.3 12.6 45.0

9.2 5.4

9.6 42.4

49.4 11.2 11.4 98.6 20.2

14.8

15.1

26.0

5.6 3.7 3.6 7 .O 4.5 2.1 2.3 4.9

3.7 6 .O 4.3 8.2 2.0 2.6 7.7 4.8

6.9 6.5 4.1 12.0 2.2 4.1

3.2 1.2 5.2

__

2.3 1.6

6

c1: Adiopodoume (R2) Anguededou Azaguie Divo Bouake Korhogo Saria (P 7) Gonse

f. h.

2.1 -

f.

f.

0.88 (b)

s.f. S. S. S.

0.3 0.3 -

4.9 1.8-3.4 0.8 0.6-2.2 -

(b)

(a)

3.9 -

-

0.3 1.9 -

HCO3 (estimated)

Adiopodoume (R2) Anguededou Azaguie Divo Bouake Korhogo Saria (P7) Gonse

f.

h. f. f. s.f. S. S.

4.3 -

18.0 -

-

-.

4.8 4.3

6.2 6.2-18 (a)

S.

5.5

-

__

Ca : Adiopodoume (R2) Anguededou Azaguie Divo Bouake Korhogo Saria (P7) Gonse

f. h.

1.8

f.

-

f. s.f.

2.4 (b) -

S.

1.9 2.14 -

S. S.

3.8 2.4-3.4

(b)

2.4 2.6-4.9

(a)

-

-

7.3

2.4 -

1.7 3.4 -

P P

TABLE I1 (continued) Station

Vegetation*

Rainfall

Through fall

Runoff

Drainage upper level

Spring lower level

Mg : Adiopodoume (R2) Anguededou Azaguie Divo Bouake Korhogo Saria (P 7) Gonse

f. h. f.

0.4 -

-

f.

0.44 (b) -

-

s.

0.1 0.31 -

0.4 0.39-1.2 (a) -

f. h. f.

0.3 -

3.9 -

f.

0.31 ( b ) 0.3 0.39 -

2.5-3.5 -

0.82 -

1.75 -

s.f. s. s.

2.2

-

0.8-1.2

(b)

2.7 2.3 0.9 1.9 1.2 0.6 0.5 0.7

1.8 3.3 1.6 5 .O 0.4 0.9 2.4 0.8

5.1 3.8 1.4 10.0 0.3 1.8

2.6 5.3 1.3 19.6 0.3 1.8 2.9 1.7

13.1 6.7 0.9 19.2 0.3 7.1

-

1.6

0.5 .-

-

0.4 0.6

-

K: Adiopodoume (R2) Anguededou Azaguie Divo Bouake Korhogo Saria (P 7) Gonse

s.f. s.

s. s.

(b)

1.3 1.24-7.1 (a) -

4.8 4.6 1.4 5.1 4.1 1.3 1.14 1.8

__

3.3

0.1 1.o 1.3 -

Na : Adiopodoume (R2) Anguededou Azaguie Divo Bouake Korhogo Saria (P 7) Gonse

f.

h. f. f.

s.f. s. s.

s.

1.0 (b) 0.15 0.12 -

0.7--1.2 ( b )

-

0.20 0 . 1 4 4 . 3 (a) -

1.91 1.2 1.o 0.67 0.66 0.35 0.38 0.2

2.64 1.85 1.8 3.74 0.8 0.44 4.7 0.6

3.12 2 .oo 2.4 4.9 1.o 0.90 1 .o

2.43 -

__ -

1.80 6 .O

-

Organic C :

Adiopodoume (R2) Anguededou Azaguie Divo Bouake Korhogo Saria (P 7) Gonse

f. h.

1.3 -

7.4 -

s.f.

-

-

S.

1.07 1.3 -

3.8 3.9-9.7 (a)

-

1.4 -

-

1.73 (b) -

-

0.9 0.63

1.1 0.61-0.88 (a) -

0.26

0.74 -

f. f.

S.

s.

1 .o

8.7 11.3 15 14.6 2.3 3.4 1.1 3 .O

8.1 11.4 9.7 6.7 1.4 3 .o 4.5

4.5 3.6 5.4 4.4 4.5 2 .o 1.1 2.8

3.3 4.3 2.3 6.1 1.9 1.2 1.9 2.4

3.4 5.9 2.2 5.9 1.9 2.1 2.3

1.6 -

0.77 1.13 0.9 3.5 0.74 0.25

1.61 1.70 0.9 3.7 0.93 1.o

0.17

0.30 0.52 0.3

0.42 0.60 0.2

11.2 11.5 9 12.9 8.5 5.2 2.5 3.2

-

-

1.2 0.7 -

Total N : Adiopodoume (R2) Anguededou Azaguie Divo Bouake Korhogo Saria (P 7) Gonse

f.

h. f.

f.

s.f. S. S.

S.

2.5

2.2-2.4

(b)

0.8 1.3

N-NO3 : Adiopodoume (R2) Anguededou Azaguie Divo Bouake Korhogo Saria f P7) Gonse

f. h. f.

f. s.f. S.

S.

0.15 (b) 0.48 -

0.13-0.21 (b) 0.21

1.17 0.51 0.4 2.4 0.46 0.20

0.24 -

0.27 -

0.73 0.58 0.6

-

-

-

S.

N-NH4 : Adiopodoume (R2) Anguededou Azaguie

f.

h. f.

__

-

0.03 0.5

TABLE I1 (continued) ~

Station

Vegetation*

Rainfall

Throughfall

Runoff

Drainage upper level

Divo Bouake Korhogo Saria (P7) Gonse

f. s.f. s. S. S.

0.27(b)

-

0.26

-

-

0.22-0.3(b)

-

0.10

0.7 0.25 0.13

Spring lower level

0.7 0.14 0.14

0.2 0.13 0.28

-

-

-

-

0.87

2.23 0.84 1.13 1.27 1.02 0.74 0.91 1

0.67 1.10

0.77 0.98 1.o 0.34 0.26 0.90

0.8

0.7

0.75 0.40 0.4 0.34 0.59 0.30 0.15 0.10

0.26 0.50 0.35 0.55 0.07 0.20 0.07 0.17

0.30 0.70 0.27 0.26 0.05 o .ia

-

0.09

-

PO4 : Adiopodoume (R2) Anguededou Azaguie Divo Bouake Korhogo Saria (P7) Gonse

f. h. f. f. sf. S. S. S.

0.32

-

0.99(b)

-

1.2-2.5 (b)

-

-

0.30 0.77

0.38 1.1-0.4(a)

-

-

o .a

0.97 0.38 0.57 0.13

-

1.25

-

-

-

0.92 0.15

-

F e z O 3: Adiopodoume ( R 2 ) Anguededou Azaguie Divo Bouake Korhogo Saria (P7) Gonse

f. h. f. f. s.f. s. S. S.

0.04

-

0.12

-

-

0.05(b)

0.08-0.17(b)

-

-

0.13 0.03

0.12 0.10-0.18 (a)

-

-

-

0.10

0.06

-

-

0.14 0.15

-

A1203 :

Adiopodoume (R2) Anguededou Azaguie Divo Bouake Korhogo Saria (P7) Gonse

f. h. f. f. s.f. S. S. S.

0.08

-

0.06 (b)

0.11

-

0.05-0.08 (b)

-

-

0.07 0.04

0.07 0.07

-

-

0.28 0.30 0.23 0.20 0.51 0.38 0.05 0.08

0.28 0.40 0.24 0.41 0.09 0.68 0.11 0.30

0.35 0.30 0.27 0.22 0.04 0.63 0.06

0.07

-

-

0.16 0.01 -

SiOz : Adiopodoume (R2) Anguededou Azaguie Divo Bouake Korhogo Saria (P7) Gonse

f.

0.7

1.3

h. f. f. s.f.

-

-

0.30 (b)

1.2-0.4 (b)

S.

0.7 0.68

1.4 0.78-0.9

S.

S.

-

-

-

(a)

3.5 2.2 4 .O 3.6 5.1 3.1 1.06 3.9

(a) = under gramineous plants or trees, respectively. (b) = sampling at Lamto: savannah or forest, respectively (Villecourt and Roose, 1978). (c) = mean value of 24 samples. * f. = forest; h. = hevea; s. = savannah; s.f. = wooded savannah.

7.2 3.8 5.4 11.5 2.4 6.8 16.3 4.4

8.3 5.2 5.6 15.5 2.8 6.3 -

6.5

8.4

-

19.2 35.6 -

16 + o A

No ppm

Troughfall Runoff water5 Drainage waters Sprmg waters

I

20

10

213 223

237 239

263

+ 0 log drainage

213

222

237 239

263

294

log dramage

so, w m 0

i

.

0

H

0

4

+ A

___

I 213 223 237 239 GO so

Bo

01 KO

263 A2

294 Ad An

log dmlnage

213 223 237 239

Go so

Bo

DI

KO

263

294

A2

Ad An

log drainage

Fig. 6. Mean concentrations of total dissolved solids in different types of waters, in relation with soil drainage intensity. The plotted points give the MPV concentration values as a function of the logarithm of the drainage height in the soil (see Table 11).

is known that, generally, the dissolved concentrations of surface and phreatic waters decrease for increasing specific flow (Meybeck, 1976; Bourrib, 1978). The empirical relation, established by Meybeck for water of the main rivers of the world is:

M

=

aQs-0.6

where a is a constant; M is the total dissolved concentration in mg/l; and Qs is the specific discharge flow in m3 s-l km-2. A striking feature, in Fig. 6, is the similarity of the concentration values in the different types of water for increasing drainage volumes, except for Si02 and Na of spring waters, which vary from 8ppm in southern Ivory Coast to 35 ppm in the pre-Sahelian area. This similarity is interesting. No doubt, dissolved amounts are perceptibly higher in rain and throughfall waters of the sub-Equatorial area than in those of the tropical and pre-Sahelian areas (see Table 11). However, the soil solutions are submitted to a much stronger evapotranspiration in the latter than in the former area. The theoretical concentration soil solutions of dissolved species coming from throughfall waters, after evapotranspiration, may be calculated using the concentration factor c. This factor is:

TABLE I11 Hydrologic balance of the different stations, West Africa AdiopodoumeAnguedegou

Azaguie

Divo

Bouake

Korhogo

Saria

Gonse

Annual rainfall (mm)

2,131

1,767

1,453

1,202

1,353

826

860

Real evapotranspiration (mm)

1,230

1,305

1,204

1,028

1,064

649

813

Runoff (mm) Deep drainage (mm)

22 879

35 427

15 235

4 170

41 248

41 136

22 25

Total flow (mm)

901

462

250

174

289

177

47

Concentration factor: (rainfall)/(drainage) ratio

2.42

4.14

6.18

7.08

5.46

6.07

34.4

18

c = RID where R is the mean annual rainfall height (mm); and D is the mean annual drainage height (mm). This ratio (Table 11) ranges from 5.46 and 6.07 in Korhogo and Saria, respectively, t o 2.42 in Adiopodoume. With the help of the concentration factor a balance of the chemical soilwater interactions may be established (Table IV): the negative values in column 3 indicate the amount of dissolved constituents which seem to be retained in the soil, positive values indicate the amount of dissolved constituents, which seem t o be added t o the initial concentration of throughfall water, when it percolates through the soil layer. In column 4 the percentages of each element in the soil, subtracted or added, have been calculated. The balance shows that the majority of dissolved elements is retained in the soil, probably because their biochemical precipitation or their consumption by roots or micro-organisms. This retention concerns primarily the nutrients (C, N, PO4); the absorption exceeds generally 50% of the theoretical concentration of the soil water [(theoretical concentration) = (throughfall concentration) x c , see Table IV, column (I)]. The major anions and cations are almost always retained, but in a smaller proportion. The sesquioxides show an irregular behaviour: they seem t o be sometimes absorbed, sometimes added. This irregularity is perhaps due to the presence of a part of these elements in colloidal form, the proportions of which may fluctuate in relation t o the conditions of sampling and conservation of the waters. Consequently, only a fraction of the chemical charge which comes into the soil, where it is concentrated by evapotranspiration, is carried down with drainage waters into the deep layers of soil and into phreatic waters. The retention of dissolved solids exists also in the deep layers of the soil (between 2 m depth and the water table), since the concentration of solution in spring waters is lower than in drainage waters, except for Si and Na, as above mentioned. In these very dilute solutions, chemical precipitation of salt is very unlikely. The chemical composition of water seems to be rather controlled by biochemical and biological processes: most of chemical elements released by decay of plant detritus and humification are recaptured by roots and micro-organisms in the aerated layers of the soils. The biological turnover is more active in the more dense forests than in bush savannah; so the flux of products is 5- t o 10-fold more intensive in the sub-Equatorial than in the pre-Sahelian areas, in relation to the much more important volume of drainage waters. But in terms of dissolved concentrations. the differences are not necessarily significant for soil waters which drain through the biologically active part of the soil. For the sesquioxides, the problem could be different. Throughfall and runoff waters, with near-neutral pH, are undersaturated in SiOz, with respect to quartz (concentration below mol/l, i.e. 6 mg/l; see Tardy, 1969), but soil drainage and spring waters are saturated or oversaturated (see Table 11).

19

TABLE IV Balance of the chemical interactions in the soil (2) dissolved amounts in drainage water (mg/l)

(3) balance = ( 2 ) - ( 1) (mg/l)

2.42 2.42 2.42 2.42 2.42 2.42 2.42 2.42 2.42 2.42 2.42 2.42 2.42 2.42 2.42

5.30 3.45 7.85 2.88 5.55 5.50 31.7 1.41 8.4 3.35 0.36 0.72 0.28 0.31 7.75

--3.90 -1.87 -1.60 -1.36 -0.50 -6.36 -11.86 -0.38 -9.51 -2.70 -0.29 -1.39 -0.01 t0.04 +4.6

+146

2.40 X 0.40 X 1.30 X 0.20 X 1.70 X 0.80 X 6.20 X 0.21 X 3.80 X 1.10 X 0.10 X 0.38 X 0.12 X 0.07 X 1.40 X

5.46 5.46 5.46 5.46 5.46 5.46 5.46 5.46 5.46 5.46 5.46 5.46 5.46 5.46 5.46

3.35 1.35 4.45 0.67 4.60 2.50 14.9 0.62 3.20 1.55 0.21 0.73 0.19 0.66 6.55

-9.75 -0.83 -2.65 -0.42 -4.68 -1.87 -18.95 -0.53 -17.55 -4.46 -0.34 -1.34 -0.47 0.28 -1.09

-74 -38 -37 -39 -50 -43 -56 -46 -85 -74 -62 --65 --70 74 -14

2.60 X 0.39 X 1.24 X 0.14 X 1.70 X 0.60 X 6.20 X

6.07 6.07 6.07 6.07 6.07 6.07 6.07

7.70 2.40 2.90 4.70 3.40 2.30 42.40

-8.08 +0.03 -4.63 +3.85 -6.92 -1.34 +4.77

-51 +1 -61 +453 -67 -37 4- 1 3

(1) (dissolved amounts in throughfall waters) X c (mg/l)

(4)

per cent leached (-) or absorbed (+) in soil ( 3 ) / ( 1 )x 100

Ad iopodoume : Ca Mg K Na

so4 c1 HC03 NO3 C N NH4 PO4 Fez03 A1203 Si02

3.80 X 2.20 X 3.90 X 1.75 X 2.50 X 4.90 X 18.0 x 0.74 X 7.40 X 2.50 X 0.27 X 0.87 X 0.12 X 0.11 X 1.30 X

-42 -35 -17 -32 -8 -54 -27 -21 -53 -45 -45 -66 -3

+15

Korhogo : Ca Mg K Na

so4

c1 HC03 NO3 C N NH4 PO4 Fe’203

A1’203

Si02

+

+

Saria : Ca Mg K Na

so4 c1 HC03

20

TABLE IV (continued) (1)

(dissolved amounts in throughfall waters) X c (mg/l)

NO, C N PO4 Fe203 A1203

Si02

.-

3.90 X 0.61 X 0.40 X 0.10 x

6.07 6.07 6.07 6.07 0.07 X 6.07 0.78 X 6.07

(2) dissolved amounts in drainage water (mg/l)

(3) balance = (2) - ( 1 ) (mg/l)

-

-

1.10 1.90 0.13 0.07 0.11 16.30

-22.6 -1.80 -2.30

-0.54 -0.31 +15.88

(4) per cent leached (-) or absorbed (+) in soil ( 3 ) / ( 1 x) 100 -95 -49 -9 5 -89 -74 +3,780

Little dissolution of quartz and a fortiori of the siliceous constituents (micas, phytolites, etc.) is possible, in the top layer of the soil. On the other hand, kaolinite which is the exclusive clay mineral in all the soils of the bioclimatic sequence, must be stable, since at pH 6-7 and for dissolved silica near mol/l, the solutions are saturated with respect t o this mineral, when dismg/l (Gardner, 1970). The addition solved A1 exceeds lo-’ mol/l, i.e. of A1 in the chemical charge of drainage waters might result from the colloidal mobility of this element. Finally, the mean dissolved concentrations in the waters are rather uniform from south t o north, in spite of very marked differences in bioclimatic conditions, perhaps because the decreasing biological activity from the Equatorial t o pre-Sahelian area would be balanced by the increasing importance of the evapotranspiration and of the concentration ratio. This bioclimatic regulation would be an important factor, together with the impact of human activity, which would explain the large variability of the mineralization values in surface waters vs. the specific flow (Meybeck, 1976).

CONCLUSION

The chemical analysis of runoff, throughfall, seepage waters (drainage) and spring waters from eight localities of West Africa, which are bioclimatically well differentiated, shows that the mineralization is increasing from rainfall and throughfall waters t o runoff and t o drainage waters, but decreases in spring waters, except for Si and Na. The instantaneous and seasonal variations of mineralization, in relation to rainfall intensities and dryness occurrences, are very marked. To estimate the chemical flux through the soil layers and the biogeochemical balances, it is necessary t o calculate the mean dissolved concentration balanced by flow volume (MPV: “moyenne pondbree par volume”), which assumes the accurate hydrologic budget of the soil t o be known.

21

The mean values obtained for the West Africa stations are at the same concentration levels, whatever the stations may be, in spite of the marked bioclimatic differences and of the increasing importance from south to north of the concentration ratio (related to evapotranspiration). This similarity would result from the regulation due t o biological activity: decreasing of dissolved concentrations in rainfall and throughfall waters and reducing of the plant residue decay in soil from the Equatorial to pre-Sahelian areas would be balanced by the increasing importance of the concentration ratio.

REFERENCES Bourrie, G., 1978. Acquisition de la composition chimique des eaux en climat temperg. Application aux granites des Vosges et de la Margeride. Univ. Strasbourg, Strasbourg, Strasbourg, Mhm. Sci. Geol. 52, 174 pp. Crozat, G., 1978. L’aerosol atmospherique en milieu naturel. Etude des differentes sources de potassium en Afrique de I’Ouest (C6te d’Ivoire). D. Sci. Thesis, University of Toulouse, Toulouse, 70 pp. D e s s h e Delepoule, A., 1978. Les eaux de la region Rouennaise. Contamination et interrelations des pluies, nappes et rivieres d’apres leurs caracteristiques chimiques et isotopiques. D. Spec. Sci. Thesis, University of Rouen, Rouen, 159 pp. Duchaufour, P., 1977. Pedologie, I. Pedogenese et classification. Masson, Paris, 477 pp. Gardner, L.L., 1970. A chemical model of the origin of gibbsite from kaolinite. Am. Mineral., 55: 1380-1389. Gouzy, M., 1973. Methodes d’analyses utilisees dans le Laboratoire d’Analyses du Centre ORSTOM d’Adiopodoum6. Rapp., O.R.S.T.O.M., Abidjan, 432 pp (mimeographed). Jacks, G. 1973. Chemistry of some ground water in silicate rocks. R. Inst. Technol., Stockholm, 73 pp. Mathieu, P., 1972. Apports chimiques par les precipitations atmospheriques en savane et sous fordt. Influence du milieu forestier sur la migration des ions et des transports solides (Bassin d’Amitioro, Cote d’Ivoire). D. Sci. Thesis, University of Nice, Nice, 441 pp. Meybeck, M., 1976. Total mineral dissolved transport by world major rivers. Hydrol. Sci. Bull., 21(2): 265-284. Nalovic, L., 1968. Les methodes d’analyse des sols et des eaux utilisees dans le Laboratoire du Centre ORSTOM d’Adiopodoumd. Rapp. O.R.S.T.O.M., Abidjan, 123 pp. (mimeographed). Roose, E.J., 1978. Pedogenke actuelle d’un sol ferrugineaux issu de granite pour une savane arboree du plateau Mossi (Haute-Volta) - Gonse - Campagnes 1968 a 1974. Rapp. O.R.S.T.O.M., Paris, 1 2 1 pp. (mimeographed). Roose, E.J., 1980a. Dynamique actuelle d’un sol ferrallitique tr& desature sous cultures et sous for& dense humide subdquatoraiale du Sud de la Cote d’Ivoire. Rapp. O.R.S.T.O.M., Paris, 204 pp. (mimeographed). Roose, E.J., 1980b. Dynamique actuelle de sols tropicaux d’Afrique de l’Ouest. Etude experimentale des transferts solides et en solution. Impact de l’utilisation des sols. D. Sci. Thesis, University of Orleans, Orleans, 587 pp. Schoeller, H., 1966. Comparaison de la composition chimique des solutions de sol a rendzine avec celle des nappes phreatiques sous-jacentes. C.R. Acad. Sci., Paris, 262: 337338. Schoeller, M., 1958. Variation de la composition chimique des solutions du sol avec la nature pedologique ‘et comparaison de la composition chimique de ces solutions avec celle des nappes phreatiques sous-jacentes. C.R. Acad. Sci., Paris, 246 : 2507-2508.

22 Schoeller, M., 1962, Les eaux souterraines. Masson, Paris, 462 pp. Tardy, Y., 1969. Geochimie des alterations. Etude des a r h e s et des eaux de quelques massifs cristallins d’Europe et d’Afrique. Mem. Serv. Carte G601. AlsaceLorraine, Strasbourg, 31, 199 pp. Vedy, J.C. and Bruckert, S., 1979. Les solutions du sol. Composition et signification pedogenetique. In: M. Bonneau and B. Souchier (Editors), Pedologie, Tome 2: constituants et proprietes du sol. Masson, Paris, pp. 161-186. Villecourt, P. et Roose, E.J., 1978. Charge en azote et en elements minbraux majeurs des eaux de pluie, de pluviolessivage et de drainage dans la savane de Lamto (C6te d’Ivoire). Rev. Ecol. Biol. Sol, 15(1): 1-20. Wischmeier, W.H., 1962. Rainfall erosion potential - Geographic and location differences of distribution. Agric. Eng., No. 43, pp. 212-215.

23

GEOCHEMICAL AND ISOTOPIC CHARACTERISTICS OF SPRING AND GROUNDWATER IN THE STATE OF SXO PAULO, BRAZIL

M. SZIKSZAY' , J.-M. TEISSEDRE2, U. BARNER3 and E. MATSU14 Instituto d e Geociencias, Universidade de ,960Paulo, ,960Paul0 (Brazil) 2Cornpanhia d e Tecnologia de Saneanento Ambiental, C.E.T.E.S.B., S6o Paul0 (Brazil) TAHAL Consulting Engineers, Tel Aviv (Israel) Centro de Energia Nuclear na Agricultura, Piracicaba (Brazil) (Accepted for publication April 16, 1981)

ABSTRACT Szikszay, M., Teissedre, J.-M., Barner, U. and Matsui, E., 1981. Geochemical and isotopic characteristics of spring and groundwater in the State of Sao Paulo, Brazil. In: W. Back and R. LGtolle (Guest-Editors), Symposium of Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54 : 23-32. A study of spring water shows that a correlation exists between the physical and chemical characteristics of the water and the lithology from where it issues. Water from crystalline rock can be classified as Ca-Mg-bicarbonate, with low conductivity and temperature; water from sediments and/or weathered crystalline rock as Ca-Mg-chloride-sulfate; and from volcanic rock, diabase and basalt as Na-bicarbonate water. Monthly samples of eight springs and of rain water in the region of Aguas da Prata were analyzed for the deuterium and ' $ 0isotopic contents expressed as 6D and 6l8O in order to determine the origin of these waters. The coincidence of the isotopic values of spring water with the regional meteoric line indicates a local source of recharge. Chemical anomalies of groundwater in the shallow Bauru and Basalto aquifers in the Parana Basin are probably caused by ascending water from the confined deep Botucatu-Piramboia aquifer through fracture and fault zones. GEOLOGY OF THE STATE O F SAO PAULO

About 25%of the State of Sio Paulo is covered by outcrops of the crystalline basement. The remaining area is covered by sedimentary rocks of the Parana Basin and the small Tertiary basins of Sao Paulo and Taubate, plus other restricted zones of coastal sediments and alluvial deposits. The crystalline basement is composed of igneous and metamorphic rocks such as granite, gneiss, phyllite, schist and quartzite. These Precambrian rocks extend in a strip encompassing the coast of Sao Paulo with a width of -100 km and a length of 480 km. The majority of the springs studied are in this strip, principally in the northeastern part of the State. Triassic-Jurassic sandstone of the Botucatu-Piramboia Formation crops out in an area of 17,000km2 and has a thickness of 300m. In the Early Cretaceous, several dozen basaltic lava flows occurred, accumulating to a

24

I I

I I

25

thickness of 1600 m which comprises the Serra Geral Formation. The Botucatu-Piramboia Formation together with the Serra Geral Formation represent the Sio Bento Group, which was covered by the Baum Group in the Late Cretaceous. The Bauru Group consists of sandstone, clayey sandstone or siltstone with, or without, carbonate cement. These sediments crop out in an area of l o s km2 and reach a maximum thickness of 300m. The Botucatu -Piramboia Formation, the Basalto and the Bauru Group constitute the principal regional aquifers which extend throughout the entire Parana Basin.

CLASSIFICATION OF CHEMICAL CHARACTER OF SPRING WATER

The studied springs discharge from almost all of the geological formations within the State of Sio Paulo, from the crystalline rocks t o the sediments of the Parana Basin (Fig. 1).Most of the springs, 43 in number, are from the magmatic and metamorphic rocks of the Crystalline Complex, nine from the sediments of the Paranfi Basin, and one spring is in Tertiary sediments of the

-1

Alluvlum

Tinguaite

Sandstone

Phonoli te

C

Spring and isotope relation

-y

Fault

w

Eq

Siltite - Shale

\^AAAA A A h

1

A ’

Diabase

Volcanic breccia Basement

Inferred fault

’/ Route / Rail way

3

River

Fig. 2. Geologic map of Aguas da PI:ata and 6 l8 O.-valuesfor sampled springs.

26

S5o Paulo Basin. Of these last ten springs, seven are from deep sources. The springs of Aguas da Prata, for which the isotopic analyses were made, issue from volcanic rocks and silicified sandstone (Fig. 2). The great majority of samples (Fig. 3) can be classified as Mg-bicarbonate with Ca being the cation of secondary importance; this water is associated with granite, granodiorite, gneiss and migmatite, and minor amounts of quartzite, schist and micaschist, representing Group I. In Group 11, water of sediments and weathered crystalline rock can be designated as predominantly Mg-chloride-sulphate (Fig. 3). Group I11 can be classified as Na-bicarbonate or mixed waters issuing predominantly from granitic rock, except for analyses 31 and 30 (Aguas da Prata) which are of water from silicified sandstone. Water of Group IV is designated as Na-bicarbonate and discharges from volcanic rock, diabase and phonolite, and from sandstone interbedded with basalt. The scattered numbers correspond to volcanic rock mixed with sandy sediment and basalt, classified as bicarbonate water. Samples 72, 73 and 74 (Aguas de S5o Pedro) are sodium chloride water from sandstone and shale and are the deepest springs of the Parana Basin.

100

t~

~

-Ca

0 -

-

~~

CI+N03

100

Fig. 3. Chemical character of spring water grouped according to four hydrogeologic units, Sao Paulo, Brazil.

27

Of those samples with Ca/Mgl, 69%are of water from gneiss and migmatites and the remaining samples from rocks such as quartzite, with dolomitic limestone and sandstone. Where Ca/Mg>2, 53% of the samples are of other types of metamorphic rock and the remaining 46% from basaltic and diabasic rock associated with sandstone.

ISOTOPIC STUDY OF SPRINGS I N AGUAS DA PRATA

Monthly samplings were made of eight springs from May 1978 to April 1979. During the same period rain water was also collected. The 6 l 8 0 determinations were made by the isotopic equilibrium methods between H 2 0 and CO, (Epstein and Mayeda, 1953), with small modifications. The 6D-values were determined according t o the method described by Friedman (1953) and Friedman and Hardcastle (1970). All values are referred t o the SMOW standards (Craig, 1961) and the analytical deviation are less than 0.20yoo for 6 l 8 0 and 2yoofor 6D. The correlation line of rain water of Aguas da Prata (6D = 8.96 l8 0 20) is different from the one obtained by Craig (1961) for data which were collected principally in the northern hemisphere (6D = 8 6 l 8 0 10). Alignment of the 6D and 6 l 8 0 of the spring-water samples with the regional meteoric line (Fig. 4) indicates that the spring water is of meteoric origin. Isotopic data on the springs (Table I) indicate that only small monthly variations of 6 l 8 0 occur in spite of the great isotopical variation in rainwater; thus, it is concluded that the volume of mixing within aquifers is

+

+

-18

-16

-14

-12

-10

-8

-6

-4

-2

1

I

I

I

i

I

I

I

I

0

- -20

oRain water .Spring water

QJ0 00

...;.: ....:?.*. -:. . .."Ip .' : .' d

- -40

. (

00

'

--60 --80

00

--loo - -120

--140

x 0 'O

TABLE I Values of 6 D Date of sampling

and 6 l80('/m ) in water samples from springs, Sio Paulo, Brazil

('//00)

6D May 2 5 , 1 9 7 8 -60 J u n . 2 4 , 1 9 7 8 -60 Sep. 2 , 1 9 7 8 -62 Sep. 3 0 , 1 9 7 8 - 6 1 Oct. 2 1 , 1 9 7 8 - 59 Nov. 1 8 , 1 9 7 8 -56 Dec. 1 6 , 1 9 7 8 -61 Jan. 2 6 , 1 9 7 9 -58 Feb. 1 7 , 1 9 7 9 -63 Mar. 1 8 , 1 9 7 9 -61 Apr. 6 , 1 9 7 9 -59

__

Mean values

F. Platina

F. Paiol

~-

60

F. PrataRadioativa

F. Villela

F. PrataAntiga

F. PrataNova

F. Vitoria

_

F. d o ijoi

_

~

6l80

6D

6l80

6D

6"O

6D

6l80

61)

6I8O

6D

6"O

6D

6l80

6D

6"O

-9.0 -9.0 -8.5 -9.0 -9.2 -9.1 -8.4 -8.9 -9.1 -8.7 -9.3

-59 -60 - 59 - 59 -57 -60 -56 -63 -54 - 57 -58

-9.0 -9.2 -8.8 -8.8 -8.9 -9.0 -8.6 -9.1 -8.7 -8.4 -9.0

-48 -47 -49 -48 -47 -52 -46 -44 -45 -43 -46

-7.2 -7.5 -7.6 -7.5 -7.6 -7.4 -7.3 -7.1 -7.1 -6.7 7.5

-45 -43 -45 -45 --43 -47 --45 -45 -45 -43 -48

-7.4 -7.1 -6.7 -7.1 -6.9 -7.3 -6.7 -7.2 -7.1 -7.5 -7.2

-57 -53 --56 -55 ---59 -56 -58 -57 -~54 -58 -54

-8.7 -8.6 -8.1 -8.6 -8.9 -8.6 -8.4 -8.5 -8.7 -9.4 -8.6

-53 -58 -57 -58 -56 -56 -58 -57 -58 -57 -56

-7.6 -8.7 -8.2 -8.7 -8.7 -8.6 -8.4 -8.8 -8.5 -8.8 -8.6

-48 -52 -52 -51 -55 -49 -54 -54 -54 -50 -53

-8.2 -7.9 -8.2 -7.9 -8.1 -8.1 -8.3 -8.0 -8.2 -7.7 -8.1

-51 -49 -49 -51 -52 -42 -47 -50 -46 -49 -47

-8.0 -7.9 -7.9 -7.7 -7.7 -7.7 -7.5 -7.6 -7.4 -7.4 -7.5

-58

-8.9

-47

-7.3

-45

-7.1

-56

-8.7

-57

-8.5

-52

-8.1

-48

-7.7

~

-8.9

F. = fonte, Portuguese for spring.

__

_ _ _

~

- _

'

__

__

29

relatively great. The three formations have characteristic values of 6 l 8 0 as = - 8 . 9 ~ 0 0 ) ; (b) sandstones follows: (a) volcanic breccia and tuffs = -7.1%0); and (c) diabases ( 6 l 8 0 = -7.7 to --8.7yoO).

HYDROGEOLOGY O F THE MAIN AQUIFERS

The Botucatu-Piramboia aquifer occurs within the areas of outcrops of sandstone around the margins of the Parana Basin, dipping toward the central axis of the sedimentary basin, from an altitude of 5 0 0 - 6 0 0 m down to 600-800 m below sea level in the subsurface near Rio Parana. The Botocatu aquifer seems to be continuous all along the subsurface of the basin. The Botucatu-Piramboia aquifer is confined by underlying impermeable argillaceous layers, and by thick basalt layers which overlie the sandstone. It is recharged directly by rainfall and rivers that traverse the sandstone outcrops. Drainage of the Botucatu-Piramboia aquifer is still a controversial question because it has no obvious outlet. It is possible, however, that small volumes of water discharge into the overlying basalts in the central parts of the sedimentary basin, and in zones of fractures and faults. Although the Basalto aquifer is heterogeneous in its lithology and hydrogeological parameters it can be considered as a regionally continuous aquifer because its great thickness and extention facilitates hydrologic interconnection between the different layers. Generally, it has a low permeability, but can have a relatively high permeability in fractures and fault zones. The basalt is covered in large part by the Bauru Formation, which forms the third and the most extensive upper aquifer of the Parana Basin. 3ecause of the argillaceous and silty material the Bauru forms only a moderately permeable aquifer. In some areas, especially in the southwestern part of the basin, the Bauru Formation passes into a more coarse composition. The Bauru is cut down t o the basaltic bedrock by the major rivers that traverse the basin and drain the aquifer.

HYDROGEOCHEMICAL CHARACTERISTICS O F THE WATER O F THE MAIN AQUIFERS

In the confined Botucatu-Piramboia aquifer the water has high total dissolved solids (TDS) contents with a relatively high concentration of sodium, chloride and sulfate (Table 11). The pH is slightly acidic in the replenishment areas and generally basic in the downgradient parts of the aquifer. The temperature of the water is proportional t o the geothermal gradient and increases up to 63°C at a depth of 1450 m (Barner and Teissedre, 1980). The groundwater of the Botucatu-Piramboia aquifer is typically Ca-bicarbonate. In the confined artesian and subartesian parts, the water changes into a Nabicarbonate or Nacarbonate type, depending on the pH-value, which varies from slightly to highly basic.

w

0

TABLE I1 Typical analyses of groundwater in the Parana Basin, Brazil Aquifer

Temp. (OC)

pH (in laboratory)

Conductivity (PS/cm)

B TDS

H C 0 3 * ' COs*' C1

SO4

F

Ca

Mg

Na

K

(/ail)

SiO, ~

Bauru Bauru i3asalto Botucatu Botucatu Botucatu (confined) Bauru*' Basalto*' *I *2

25.6 23.6 23.8 24.5 23.0 63.0 24.0 24.5

6.1 6.2 7.7 6.8 7.0 8.7 9.7 9.5

Expressed as CaC03. Water showing chemical anomalies.

190 200 2 50 25 127 850 360 710

159 162 212 24 92 615 254 535

83 94 158 10 48 20 45 32

0 0 0 0 0 193 73 48

0.5 3.0 2.5 1.0 5.5 125.0 3.5 16.5

0 0 0 0 2 104 58 188

0.34 0.17 0.21 0.10 0.10 11.6 0.53 8.20

21.6 14.0 38.4 2.0 12.0 4.7 0.8 5.6

5.0 2.3 8.5 1.0 6.0 1.8 0 0.5

5 .o 20.4 12.8 0.2 1.7 230.0 73.0 158.0

3.8 5.0 2.5 0.4 1.8 2.1 0.4 1.0

18.8 69.0 51.0 12.0 22.0 33.0 37.0 39.0

0 0

0 0 0 2,230 8 32

31

In the Bauru aquifer the water can be classified into two main groups: (1)Ca-bicarbonate; and (2) mainly Na-bicarbonate and subordinately Cabicarbonate. The water in the Basalto is generally similar to those of the Bauru aquifer with the only difference being the predominance of Mg in relation t o Ca. This classification demonstrates the chemical similarity of the water from the three aquifers; nevertheless, there is a clear evolution of the geochemical characteristics of the water in the confined aquifer resulting from changes in the sodium chloride and sulfate content. Because of these and associated geochemical differences, anomalies in the shallow aquifers can be used as an indicator of the presence of ascending deep water through fractures of faults zones.

CHEMICAL ANOMALIES

The chemical character of water from several boreholes in the Basalto and Bauru is similar to the water of the Botucatu-Piramboia aquifers. Plotting different physicochemical elements on maps shows the existence of zones with relatively high concentrations of salts in water of Bauru and Basalto aquifers. Zones with TDS above 250 mg/l in the Bauru and Basalto coincides with geochemical anomalies of other elements. The TDS that is high in relation to the “normal” water of those aquifers is due mainly t o increased concentration of the bicarbonate, sulfate, chloride and sodium. Beside these characteristics, higher concentration of other elements such as fluoride and boron has been noted. These elements are also present in the deep water of the Botucatu-Piramboia aquifer. In addition higher than normal regional temperatures and either neutral or highly basic pH were measured in the Bauru and Basalto aquifers in the anomaly zones. The observed geochemical and physical anomalies (Fig. 1)show an alignment along two principal axes oriented N-S and E-W. The E-W axis coincides with the line of the Rio Grande River along -150 km in the State. The N-S axis begins in the north at the curve of Rio Grande extending over a distance of -300 km until encountering an inflection of Rio Tiete in the south. Judging by the composition of water encountered in the shallow aquifers, which resemble the deep water of the Botucatu, it is possible that great fractures exist in the Basalto with a vertical extension of 400-500m. These fractures probably permit a vertical ascending circulation of water from the Botucatu into the Bauru aquifer. The ascending warm water may become enriched with constituents in the Basalto and cause secondary mineralization. Indications of tectonic disturbance exist at the Rio Tieti3 alignment within the anomaly zone. It was confirmed through drilling that a difference of 200 m exists in thickness of the basalt between the two margins of the river,

32

and well logs show a strong local stratigraphic unconformity where the Bauru Formation is in direct contact with Permian sediments. The map of water levels in the Botucatu aquifer also shows a strong inclination in the gradient toward the anomaly zone along the Rio Tiet6, strengthening the hypothesis of drainage of the lower confined aquifer into fracture and fault zones in the shallow aquifers.

ACKNOWLEDGEMENTS

The authors wish t o express their gratitude for the grant received from the National Commission of Nuclear Energy (C.N.E.N), Foundation for the Development of Research in the State of SaG Paulo (F.A.P.E.S.P.) and the National Council of Scientific and Technical Development.

REFERENCES Barner, U. and Teissedre, J.-M., 1980. Geochemistry and high temperatures of groundwater in the Parana Basin deep sandstone aquifer. Commun. 26th Int. Geol. Congr. Paris. Craig, H., 1961. Standard for reporting concentrations of deuterium and oxygen-18 in natural water. Science, 133: 1833-1934. Epstein, S. and Mayeda, T., 1953. Variation of " 0 content of waters from natural sources. Geochim. Cosmochim. Acta, 4: 213-224. Friedman, I., 1953. Deuterium content of natural water and other substances. Geochim. Cosmochim. Acta, 4 : 89-103. Friedman, I. and Hardcastle, K., 1970. A new technique for pumping hydrogen gas. Geochim. Cosmochim. Acta, 34: 125-126.

33

REPARTITION DES ELEMENTS TRACES DANS LES EAUX THEBMOMINERALES

JEAN SARROT-REYNAULD, JACQUES ROCHAT e t JEAN DAZY U.E.R. de Giologie e t de Pharmacie, Uniuersiti d e Grenoble, Grenoble (France) (Accept6 pour publication le 11 mai, 1981)

ABSTRACT Sarrot-Reynauld, J., Rochat, J. and Dazy, J., 1981. Ripartition des elements traces dans les eaux thermominerales. (Minor-element distribution in mineral and thermal waters.) In: W. Back and R. Letolle (Guest-Editors), Symposium of Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54: 33-50. This paper is an approach to define the origin and mineralization of thermal waters, based on new data on spring waters from southern France (Massif Central and Alps) and from Iran (eastern Azerbaijan). The behaviour of some minor elements, such as B, F-, Br-, I-, Sr2+, Rb', Li', and the significance of characteristic ratios between minor and major water constituents enable us to distinguish between mineral and thermal waters associated with granitic, metamorphic or volcanic rocks and those originating from sedimentary evaporites. Trace-element distribution seems to be controlled by magmatic influences, the chemical composition of the wallrock, the presence of large distended fractures as well as by the effect of temperature on groundwaters. Ratios such as Br-/Iand Sr2+/Rb+enable us to determine the eroded-rock type. In some mineral waters the occurrence of chloride being relatively independent of other halogens seems to indicate a different origin for chlorine compared to other elements. RESUME A partir d'analyses nouvelles des eaux des sources thermominerales du Sud-Est de la France (Massif Central et Alpes) et d'Azerbaidjan oriental iranien, on definit les rapports entre elements traces (B, F-, Br-, I-, Sr2+, Li', Rb') e t majeurs qui peuvent caracteriser les eaux provenant d'un circuit thermal traversant diverses varietes de roches plutoniques, metamorphiques ou sedimentaires riches en evaporites. L'influence du magmatisme et de la presence de grandes fractures de distension qui se manifestent par des anomalies geothermiques locales ou regionales et des emanations fumerolliennes, ainsi que la composition des roches lessivees par les eaux thermales semblent au moins aussi importantes dans la repartition des elements traces que les temperatures atteintes par les eaux en profondeur. Certains rapports tels que Br-/I- ou Sr2+/Rb+permettent de determiner le type de roche lessivee mais la presence de chlorure de sodium dans certaines eaux, sans relation constante avec les autres halogenures, parait confirmer son origine independante de celle des autres dements.

34 INTRODUCTION

Les etudes isotopiques ont montre que la plus grande part des eaux thermominerales proviennent d’eaux meteoriques infiltrees 8 grande profondeur, et qui se sont mineralisees lors de leur infiltration puis de leur remontee vers la surface. Elles acqui6rent alors des temperatures d’autant plus elevees qu’il peut exister des anomalies geothermiques dans les zones qu’elles parcourent et leur mineralisation resulte de lessivages superficiels ou profonds et d’apports fumerolliens. Afin de caracteriser chaque famille d’eau thermominerale et mieux connaitre l’origine de la mineralisation, et par 18 le circuit thermal emprunte par les eaux qui parviennent aux emergences, nous avons determine, avec les elements majeurs, les anions B, F-, Br-, I-, et les cations Li’, Rb’ et Sr2+ dans les eaux de sources du Sud-Est de la France (Massif Central et Alpes) et d’Azerbaidjan iranien, situees dans des conditions geologiques comparables ou pouvant servir de references.

METHODES DE DOSAGE DES ELEMENTS TRACES

Le bore total a ete dose par colonmetrie l’aide du sulphophenyl azo-2dihydroxy-I,8 naphtalhe disulfonate de sodium. Les resultats sont exprimes - .~ en moles de bore par litre, le seuil de detection &ant voisin de 3 ~ 1 0 M Les iodures, bromures et fluorures ont ete doses 8 l’aide d’electrodes specifiques 8 membranes cristallines et nous avons utilise des solutions tampons (Vesely et al., 1978) pour minimiser les interferences qui sont un inconvenient de la methode. Pour les iodures et les bromures, les tampons sont representes par un ajusteur de force ionique (NaNO,, 2 M ) mais dans le cas des eaux sulfurees, le soufre a ete oxyde prealablement. Le seuil de detection est de M pour les iodures et de 3 ~ 1 O M - ~ pour les bromures. Pour les fluorures, nous avons employe un tampon pH (acide acetique, acetate de Na, NaC1, acide cyclohexane diamine tetracetique) qui permet egalement d’ajuster la force ionique et de complexer les cations t r i d e n t s . Le seuil de M. detection est alors de Les cations Li’, Rbf, Sr2+ont ete doses par absorption atomique et Na+ e t K+ par spectrophotometrie d’emission. Les seuils de detection sont de M pour Li+, de 2 ~ l O M - ~pour Rb’ et Sr2+,et de 3*10-’ pour Na+ et K+.

SITUATION GEOLOGIQUE DES REGIONS ETUDIEES

Azerbaidjan oriental iranien (Fig. 1) Le secteur de 1’Azerbaidjan oriental iranien situe ? 1’Est i de Tabriz correspond a une discontinuite majeure entre les chaines du Zagros et de 1’Alborz

35

Fig. 1. Schema g6ologique de I’Azerbaidjan oriental iranien. Fig. 1. Schematic geological map of eastern Iranian Azerbaijan.

et celles des Taurides, des Pontides et du Petit Caucase. C’est une zone de distension qui se prolonge vers l’Ouest en Anatolie orientale turque. La mise en place des roches volcaniques commencee au Cretace a ete particulierement intense au Paleogene, avec des epanchements d’andesites e t de basaltes et l’intrusion de diorites et granodiorites. Le volcanisme neoghne s’est manifeste tout d’abord par des projections acides puis par l’epanchement de basaltes miocenes et pliocenes. Le volcanisme quaternaire est 9 l’origine de la formation des grands massifs du Sahand (3700 m) et du Sabalan (4800 m). Celui-ci est constitue de dacites, trachytes, trachyandesites, de coulees d’andesites et de basaltes e t de tufs. L’epaisseur des volcanites tertiaires et quaternaires y est superieure 9 3500m, mais la nature de son soubasement n’est pas connue. Seuls les sediments du Cretace au Pleistocbne, ou s’intercalent des evaporites oligocenes, affleurent 9 sa peripherie. Les principales sources thermominerales se situent sur le trace ou 9 l’intersection de fractures recentes satellites d’accidents majeurs, soit B l’intx5rieur du massif volcanique quaternaire du Sabalan, soit 9 sa peripherie, soit dans les massifs volcaniques paleogenes, soit enfin simplement dans les sediments paleogenes et neogenes (Salehi, 1978). Leurs eaux chaudes, le plus souvent carbogazeuses, sont bicarbonatees ou chlorobicarbonatees sodiques et calciques (Sareine, Bouchli, Khalkhal, Abress, Bostanabad) ou chlorosulfatees et bicarbonatees calciques (Ilando, Pyremar). Les eaux des sources de

36

37

Sardabeh, Qotur Sui, Chabil, Mayur et Sarab sont sulfureuses et sulfatees calciques et sodiques. Les sources thermominerales d'Azerbaidjan oriental sont donc en liaison etroite avec de grands accidents de socle parfois encore actifs comme la grande faille de Tabriz et avec un volcanisme recent tres important et qui se manifeste encore par la presence d'anomalies geothermiques.

Sud-Est de la France (Fig. 2 ) Les sources thermominerales du Massif Central et des Alpes occidentales sont localisees pour la plupart B l'intersection de fractures de direction meridienne avec l'une ou l'autre des deux grandes familles d'accidents de direction N50"E et N140"E qui correspondent dans le Massif Central franqais aux failles varisques et armoricaines, et dans les Alpes B des decrochements senestres ou B des failles dextres. Ces accidents anciens comme les accidents N-S qu'ils recoupent ont rejoue plusieurs fois et particulierement lors des derniers mouvements de l'orogenhe alpine.

Massif Central Dans le Massif Central, les sources de Royat, Vichy, Le Mont-Dore, SaintNectaire, La Bourboule, Chaudes-Aigues, Vals, sont liees & de grands accidents du socle granitique et metamorphique, mais accompagnent tres souvent des edifices volcaniques importants: chaine des Puys, massifs du Mont Dore et du Cantal, volcans du Vivarais. Ces edifices, mis en place du Miocene au Quaternaire, comportent des basanites et phonolites, des domes trachytiques ou rhyolitiques, des projections volcaniques et de vastes coulees de basaltes. Leur diametre atteint parfois des dizaines de kilometres. Souvent en liaison avec le volcanisme quaternaire, les eaux des sources du Massif Central peuvent n'avoir traverse que les roches du socle cristallin tandis qu'& peu de distance elles traversent des appareils volcaniques comme B Vals-les-Bains (Jeantin, 1979). Elles sont essentiellement bicarbonatees sodiques ou chlorobicarbonatees calciques et sodiques. La majorite est carbogazeuse, le gaz carbonique etant d'origine profonde (Maisonneuve et Risler, 1979). Nous avons utilise, outre les analyses nouvelles realisees sur les sources de Vichy, Saint-Nectaire et La Bourboule (Terrier, 1980), celles publiee recemment (Henou, 1973; Thuizat, 1973; Jacob, 1975; Cailleaux, 1976; Cailleaux et al., 1976; Michard et al., 1976, 1978; Ouzounian, 1978; Degranges et al., 1978), malgre l'absence frequente de donnees concernant les halogenures, le rubidium ou le strontium. Comme en Iran, les venues thermominerales sont liees & la presence de grands accidents de socle qui ont permis la remontee de fluides mineralisateurs ou tout au moins de gaz profonds tel le dioxyde de carbone, et & l'exception peut-etre du bassin de Vichy tres proche du fosse oligocene de la Limagne, les eaux de ces sources ne sont pas en contact avec des evaporites.

38

Alpes franqaises La plupart des eaux thermominerales des Alpes franqaises ont lessive des evaporites. Les sources des zones internes des Alpes: Brides, Salins, La Lethere en Tarentaise (Sarrot-Reynauld et Simeon, 1979; Simeon, 1980)’ Monetier-les-Bains en Brianqonnais (Poulain, 1977; Sarrot-Reynauld et al., 1977) emergent directement du Trias gypseux A l’intersection de grands accidents submeridiens e t transverses A 1’Arc des Alpes occidentales, les sources de 1’Echaillon en Maurienne et de Plan de Phazy apparaissant au contact du socle cristallin et du Trias. Si celui-ci contient des gypses et anhydrites, il n’a pas encore ete possible d’y trouver du chlorure de sodium pourtant present en fortes quantites dans les sources de Salins et de l’Echaillon. Dans les zones externes des Alpes, les eaux des sources d’Allevard et d’Uriage (Dazy et al., 1980) emergent des assises du Lias, mais ont lessive les evaporites du Keuper comme celles de Saint-Gervais, La Motte-les-Bains ou Digne. Les eaux des sources d’Aix-les-Bains situees A la peripherie des chaines subalpines proviennent (Bosch et al., 1976; Dazy et Grillot, 1981) du melange d’eaux meteoriques rechauffees en profondeur mais peu mineralisees avec des eaux moins profondes ayant lessive les evaporites de l’Oligocene, A la faveur d’un systeme complexe de fractures satellites des grands decrochements. Les sources thermominerales des Alpes franqaises sont sulfatees calciques ou chlorosulfatees sodicocalciques. La plupart est sulfureuse mais celles de Salins, Oriol-Cornillon, La Lechere et Plan de Phazy sont carbogazeuses. Leur thermalisme n’est pas lie au volcanisme mais au rejeu recent d’accidents de socle.

COMPARAISON DES RESULTATS D’ANALYSES

Le Tableau I regroupe les resultats de moyennes analytiques. Certains caracterisent l’ensemble des sources d’une meme localite. Divers traitements statistiques se sont reveles decevants. Pour rendre la comparaison des resultats plus aisee e t definir des types d’eaux, nous avons trace pour les anions et pour les cations, separement, des diagrammes semilogarithmiques dont les colonnes sont disposees de telle faqon que les segments horizontaux traduisent les rapports molaires: B/F- = 1,

Cl-/F- = 100,

Br-/I- = 10,

F-/Br- = 1,

K/Li+ = 10,

Na+/Li+ = 10,

K+/Rb+ = 100, et

Rb+/Sr2+ = 1, Ca2+/Mg2+ = 1

Cl-/Br- = 100,

c1-/so:-

= 1,

N ~ + / K += 10, Ca2+/Sr2+ = 100,

TABLEAU I Teneurs en ions majeurs ou en traces des eaux thermominerales du Sud Est de la France et d'Azerbaidjan oriental iranien Localisation

IBI F~ 1 0 - 6 10-3

C110-3

Br10-6

I10-7

-

-

HCO; 10-3

Li+ Na' M 1 0 - ~ 103

Kf

0,18 1,08 0,08 0,12 1,77 0,19 0,71

0,28 1,95

23,5 78,6 8,O 11,2 39,35 15.6 39,6 21,7 2,95 73,8

93 74,8 8,9 11 68 14,3 40 72 43 82

0,95 3,18 0,22 0,72 4,05 12 2,6 2,65 0,42 2,77

17,7 18,7 16,6 9,3 11,45 15 22 27 6,95 12 1,92

6,8 14 1,9 5,14 8,28 16 4 7 4,54 3,95 4.34

0,533 0,54 0,12 0,37 0,09 0,35 1.35 0,3 0,07 0,09 0,Ol

40 206 67 78,3 9-7 70 56 140 15,2 45,3 1,1

2 2,5 0,46 1,48 0,36 0.66 1,06 2 0,14 0,93 0,08

22 39 15 58 11,7 23 28 20 3,5 3,5 2.3

11,5 2,4 4,7 8,2 0,6 0.2 0 18,8 2,4 0 5,4 0

1,5 0,71 0,14 0,08 0,002 0,05 0,004 0,13 0,2 0,03 0,08 0,003

83 39 7,2 39 1

6,2 3.12 0,92 1,6 0.02 0.36 0.69 2,05 1,05 0,8 0,43 0.15

210 100 28 14 2.3 3.5 10 24.5 27 10 8 2,3

SO:-

10-3

Rh'

10-~

SrZt 10-~

CaZ+

Mgz'

Temperature emergente ("C)

144 48 27,5 15 511 180 620 367 19 96

5,05 1,34 0.34 0,5 4,2 2,l 6,5 3 0,16 6,45

2,08

3,5 1,36 5,O 2,32 0,2 0,43

23 13 51 80 36 43 34 53 16 66

1.900 2.200 1.300 800 896 1.206 1.230 2.150 400 1.023 138

16 20 13 8,25 12,3 16,5 8,l 11,8 4,57 8,12 2,47

4,6 6 2 2 2 4,45 1.34 5,46 2,38 3,21 1.04

36 33 50 40 35 27 42 27 16 40 46

510 140 45 4.200 S 65 40 355 85 34 55 100

1.6 3,32 7,8 6.5 0,22 1.15 2,22 2,55 2.82 0,77 1,45 6.75

1,31 0,65 0.74 4,5 0,001 0,25 0,98 1.4 0,95 0,41 1.85

55 88 43 43 40 57 45 30 38 40 50 36

Massif Central. Neyrac Vals Camuse St. Laurent Les Bains Chaudes Aigues St:Nectaire-Parc Le Mont-Dore--Chanteurs Royat--EugCnie La BourbouleChoussy La Bourboule Fenestre VichrLys

-

352 -

0,04 0,17 0,55 0.18

0,549 2,19 0,56 2,15 40,5 6,23 27,2 49,6 2,3 9,77

505,5 105 690 426 71,4 130

0,03

218 199 144 168 74 68.5 161 138 25 69,4 17,6

0,18 0,09 0,14

0,18 0,05

36 210 3 -4 79 9,4 65,6 29,7 115,5 9,5 37,l 1,16

313 223 80 120 20 46 23 186 110 40 3 3

0,146 0,094 0,024 0,198 0,069 0,215

78.3 43,4 6,42 48 0.22 3.04

0,026 0,05 0,34 0,015

0,63

-

6.7 33,7 15,5 25 55 7,75 87,5

2,5 3,33 1,66 3,33 4,17 16,5 5

0,14 0,56 0,32 0,22 0,86 0,44 1,3

1,08

0,88 0,08

0,75

6 17,3

-

8,3 80

-

96

1,66 0,04 0.28

Alpes franqaises: Brides Salins La,Lechere L'Echaillon-St. Jean M o n e t i e r L a Rotonde Plan de Phazy-La Rotonde Saint Gervais Uriage Allevard Digne Aix-les-BainsAlun

0,21 0,07 0.06 0,29 0,13 0,04

137,5 690 42,5 300 187,5 837,5 1.150 812.5 77,5 312,5 25

15 23 28 40 2,36 7,87 11,8 28,3 4,72 19,7 2

Azerbaidjan oriental. Bouchli Pyremar Sareine Bostan Abad Sarah Khalkhal Mayur Abress Ilando Qotur Sui Chabil Sardabeh

0,07

0,032 0,067 0,052 0,02 0,Ol

0,28

23 11,7 2,7 0,14 0.39

350 262 17,5 412 25 25 26 200 60 4.5 42,5 0.37

23,6 23,6 14,3 5,5 4,7 13.3 56 11,8 4,7 0,s

18,l 2

3,54 2,6 1 3,25 0,86 1,04 3,3 4,68 3.12 4.8 4,68 3,l

10 2,4 44 11,6 4,s 7.6 1

0.41

w (0

40

41 RAPPORTS ENTRE LES DIVERS IONS

Sud-Est de la France (Figs. 3 et 4) Massif Central. Dans le Massif Central (Fig. 3), des profils pratiquement identiques traduisent l'homogeneite des teneurs en cations des eaux de Royat, Saint-Nectaire et du Mont-Dore. Pour ces sources, le rapport B/F- est superieur a 10 mais il est voisin de 2 pour les sources de La Bourboule et, inferieur a 1a Vichy. Le rapport Cl-/Fest compris entre 250 et 1350 sauf pour les eaux de Vichy (Cl-/F- = 15). Le rapport Cl-/Br- est voisin de 1000 tandis que le rapport Br-/I- est proche de 100. Le profil C1-Br-I quasiment rectiligne (Cl-/Br- = 10 x Br-/I-) que l'on retrouve pour les eaux de la source Choussy a La Bourboule, est tres caracteristique. Le rapport SOi-/I- generalement voisin de 3000 n'est que de 150 pour les sources Choussy et Fenestre e t le rapport Cl-/SO;- est compris entre 8 pour les eaux de la source Fenestre et 50 pour celles de la source du Parc B Saint-Nectaire. Malgre la proximite du socle cristallin et de gisements de mineraux fluores, les fluorures ne sont pas tres abondants sauf dans les eaux du bassin de Vichy. Le bore parait le plus souvent lie au COz qui est d'origine profonde. Cette relation, classique dans un certain nombre de gisements geothermiques (Cormy et al., 1970), confirmerait l'origine fumerollienne du bore (Cailleaux, 1976). La repartition des cations majeurs ou en traces est identique dans les eaux de Royat et du Mont-Dore ou celles de Vichy et de Chaudes-Aigues. Li relativement abondant dans les eaux issues du socle cristallin et volcanique du Massif Central comme dans les roches qui le constituent (Letolle, 1967), est bien correle avec K ( r = 0,96) et un peu moins bien ( r = 0,86) avec Na. Le rapport Na+/Li+ propose comme g6othermomGtre (Michard, 1979) varie de 38 a 140, indiquant des temperatures en profondeur comprises entre 140' et 240°C alors que le geothermom&treNa+/K+,sensible aux echanges de bases, varie de 10 a 40 indiquant des temperatures allant de 120' B 260°C. L'utilisation des rapports K+/Rb+ et Li+/Rb+ comme geothermometres n'a pu etre testbe valablement faute de donnees assez nombreuses. Le rapport Sr*+/Rb+,voisin de 1 pour les sources de Vichy et de Chaudes-Aigues, atteint 20 pour les sources du Mont-Dore. Alpes frunpises. Dans les eaux des sources des Alpes (Fig. 4) ayant lessive des evaporites, le rapport B/F- est beaucoup plus faible que dans les eaux carbogazeuses du socle cristallin ou volcanique du Massif Central. I1 est m6me inferieur 2 1 pour les sources de Saint-Gervais, Allevard, Digne ou Aixles-Bains qui sont pour la plupart sulfureuses. Le rapport Cl-/F- vane de 20 pour les eaux d'Aix-les-Bains et La Lechere, 100 pour celles de Monetier et de Digne et ti 2000 pour celles de Salins et Plan de Phazy. Fig. 3. Rgpartition des ions dans les eaux des sources du socle cristallin et volcanique du Massif Central franqais. Fig. 3. Ion assessment in spring waters from the crystalline and volcanic basement of the Central Massif, France.

42

43

Le rapport Cl-/Br- qui est de 650 dans l'eau de mer varie de 30 A SaintGervais, a 75 a La Lechere et atteint 300 dans les eaux des sources de Salins, Brides, 1'Echaillon et Aix-les-Bains. Les eaux de Saint-Gervais, Monetier et La Lechere sont donc enrichies en bromures, fait qui semble caracteriser les eaux issues des formations evaporitiques dans lesquelles le rapport Br-/I- est de l'ordre de 1000. Le rapport SOi-/I- est compris entre 1000 et 20.000 et le rapport Cl-/SO$- ne depasse pas 1 0 a Salins et peut etre inferieur a 1ti La Lechkre. Les eaux sulfatees issues des evaporites se caracterisent donc par des rapports Br-/I- eleves, les teneurs en chlorures fluctuant de faGon independante a la fois des sulfates et des autres halogenures. L'identite des diagrammes des teneurs en cations des sources des Alpes traduit bien par ailleurs l'homogeneite des roches traversees malgre les centaines de kilombtres qui separent certaines emergences ou l'age different des evaporites: le profil des eaux d'Aix-les-Bains issues des evaporites oligocenes est absolument analogue a celui des eaux de Monetier-les-Bains issues du Trias gypseux. Le rapport Na+/Li+ varie de 40 a 500 et le rapport Na+/Kf de 13 a 110, c'est-a-dire beaucoup plus que dans les eaux du Massif Central. Les temperatures en profondeur calculees par ces deux geothermomktres vont de 50 a 230°C mais sont tres fortement discordantes pour les sources sulfureuses de Saint-Gervais (230" et lOO"C), Uriage (55" et 84"C), Allevard (95" et 59"C), Digne (50" et 110°C) e t Aix-les-Bains (80" et 220°C). En fait, Li ne parait pas en relation directe avec Na et K. Le rapport K+/Rb+compris entre 250 et 2500 parait mieux correle ( r = 0,84) avec les temperatures des eaux en profondeur calculees par les geothermometres geologique et Na-K-Ca qui donnent en general des resultats concordants. Le rapport Li+/Rb+varie de 50 a 500 dans les eaux issues des evaporites alors qu'il varie de 70 a 230 dans celles issues du socle du Massif Central. Le rapport Sr2+/Rb+est toujours egal ou superieur A 50, fait que l'on ne rencontre que dans les eaux issues des evaporites et qui semble les caracteriser. Le rapport Ca2+/Sr2+est voisin de 100 comme dans les eaux du Massif Central mais le rapport Ca2+/Mg2+peut atteindre 10 ce qui est rare dans les memes eaux. En resume, l'ion sodium, fortement lie au chlorure, semble se surimposer A un profil dont l'homogeneite est par ailleurs remarquable, sans que l'on puisse determiner si sa presence dans certaines eaux est due a des conditions geologiques (yeux de sel), thermiques ou a des apports fumerolliens. Azerbaidjun oriental iranien (Fig. 5) Dans les eaux des sources de Bouchli, Pyremar, Abress, Sareine et Ilando, la liaison est tr&smarquee entre les fortes teneurs en C 0 2 d'origine profonde (Sarrot-Reynauld et al., 1978) et les teneurs importantes en ions bore. Le rapport B/F- varie de 2 6. Les eaux sulfureuses de Chabil, Sardabeh, Sarab, Fig. 4 . Rdpartition des ions dans les eaux issues des dvaporites des Alpes franqaises. Fig. 4 . Ion assessment in spring waters originating from evaporites of the French Alps.

44

45

Qotur Sui, Mayur et Khalkhal presentent de fortes teneurs en fluorures (0,12 < B/F- < 0,65) que l’on est tente de relier a des apports fumerolliens. Les analyses isotopiques montrent qu’une bonne part des composes du soufre y est elle-meme d’origine volcanique (Salehi, 1978). Dans les emergences oh ils sont les plus abondants, les ions bore ou fluor semblent donc chacun en relation avec les gaz associes. Les sources tres chaudes de Bouchli et Pyremar, situees a des distances considerables de part et d’autre du Massif du Sabalan, presentent des profils absolument analogues. Leur localisation sur un meme accident tectonique recoupant des formations de composition et d’age differents y faciliterait des apports fumerolliens identiques. Par ailleurs, le rapport Cl-/F- va de 500 a 700 pour les sources caybogazeuses mais n’est plus que de l’ordre de 1 0 pour les sources sulfureuses. Le rapport Cl-/Br- qui ne parait plus lie 2i la nature des degagements gazeux aux emergences varie de 500 pour les sources de Sareine a 1 pour celle de Mayur. Les rapports Br-/I-, sauf pour la source de Bostanabad, proche d’evaporites oligocenes, sont plus voisins de ceux observes dans les eaux du Massif Central que de ceux des eaux issues des formations evaporitiques des Alpes. Certaines teneurs elevees en iodure (Mayur) resulteraient de la presence d’H2S qui les presewerait de l’oxydation lors de melanges eventuels avec des eaux superficielles riches en oxyg&ne. Cette oxydation peut en effet se traduire par la precipitation partielle d’iode avant l’emergence ou au contraire un entrainement par des vapeurs chaudes. Les variations des teneurs en iodure vis-a-vis des autres halogenures pourraient donc etre dues au caractere tres oxydable des iodures et aussi a ce que leur solubilite croit avec la temperature mais decroit des iodures aux bromures et aux chlorures, les fluorures ayant un comportement independant de celui des autres halogenures. Les rapports variables entre halogenures dependent donc de la temperature des eaux en profondeur, de la nature des roches traversees et des pH e t Eh. Les rapports SOi-/I- vont de 30.000 B Qotur Sui, 1000 a Bouchli, A 650 a Mayur. Les teneurs assez fortes en sulfate des eaux des sources de Bouchli, Pyremar et Sareine laissent supposer que celles-ci ont lessive en profondeur des evaporites mais bien que leur temperature elevee entraine une diminution de solubilite des sulfates, deja nette vers 100°C, les concentrations atteintes ne sont pas considerablement plus elevees que celles des eaux thermales du Massif Central. Le rapport Cl-/SOi- inferieur B 1 pour la plupart des sources sulfureuses est compris entre 3 et 22 pour les autres sources mais les chlorures paraissent se surimposer aux sulfates dans certaines eaux, aucune liaison systematique n’existant entre ces ions. L’analogie entre les profils des teneurs en cations permet de definir un groupe de sources riches en Li pour lequel K+/Li+ est inferieur a 10 et Na+/Li+inferieur B 100 (Bouchli, Pyremar, Sareine, Ilando), un groupe pour Fig. 5. Rbpartition des ions dans les eaux des sources de la rhgion du Sabalan (Iran). Fig. 5 . Ion assessment in spring waters of the Sabalan area (Iran).

46

lequel K+/Li+ est superieur A 10 et Na+/Li+ superieur A 100 (Sardabeh, Mayur, Qotur Sui), un groupe pour lequel K+/Li+ est inferieur a 1 0 mais Na+/Li+ superieur a 100 (Chabil, Sarab, Khalkhal) et enfin un groupe de sources pauvres en Li mais riches en Sr (Bostanabad, Abress) pour lesquelles K+/Li+ est superieur A 1 0 et Na+/Li+inferieur a 100 (Rochat et al., 1979). Les ions Na', K+ et Li+ semblent evoluer independamment les uns des autres malgre des correlations fortes entre eux pour l'ensemble de la region (Na+Li', r = 0,85; Na+-K+, r = 0,94; K+-Li+, r = 0'96). Les temperatures des eaux profondes fournies par le g6othermomGtre Na+/Li+, comprises entre 40 et 2OO0C, different de celles obtenues A partir du geothermom6tre Na+/K+ (de 100 A 45OoC) et de celles calculees par les geothermombtres Na-K-Ca ou geologique (de 120 a 320°C). Les variations du rapport Li+/Rb+ sont par contre trGs bien correlees ( r = 0,96) avec ces dernieres et un peu moins bien avec les temperatures aux emergences ( r = 0'85). Ceci peut signifier que l'apport d'eau froide superficielle se fait sensiblement dans les memes proportions pour toutes les sources. Les rapports K+/Rb+, Sr2+/Rb+et Ca2+/Sr2+varient dans des proportions beaucoup plus grandes dans les eaux thermominerales que dans les roches eruptives du Sabalan en raison de la solubilite tres differente de ces ions. CARACTERES DES EAUX DES DIVERSES FORMATIONS - FACTEURS INFLUANT SUR LEUR COMPOSITION

Les groupes de profils des teneurs en anions des sources coyncident rarement avec les groupes de profils de leurs teneurs en cations. Cette independance est dejA connue pour les elements majeurs. La repartition des anions peut etre modifiee par des apports gazeux, fumerolliens ou superficiels et les variations de pH et de Eh differentes qui modifient les equilibres chimiques. Les variations des rapports entre cations paraissent surtout liees aux temperatures des eaux en profondeur et A la nature des mineraux des roches traversees. Independance des chlorures vis-d-vis des autres anions Les eaux d'Iran, comme celles des Alpes, ont des teneurs en chlorure qui ne depassent qu'exceptionnellement celles des eaux du Massif Central ou les chlorures ne proviennent pas d'evaporites (Schoeller et Schoeller, 1979) mais d'apports fumerolliens ou du lessivage du socle granitique et des laves (Jacob, 1975), les fortes teneurs en chlorures resultant veritablement de la decomposition des mineraux et non pas d'une simple concentration des chlorures des precipitations par evaporation. Dans les regions etudiees, il existe toutefois pour chaque type de formation des eaux a profils tr6s proches qui ne diffGrent que par la presence ou l'absence de NaC1. Celui-ci est donc relativement independant des autres elements.

47

Par ailleurs, si la concentration en bromure et iodure augmente souvent en valeur absolue avec celle des chlorures, il n’existe jamais de rapport constant entre les chlorures et les autres halogenures. Le rapport entre Br- et I- varie, lui, de faqon importante sous l’effet de la temperature et de reactions d’oxydation mais aussi en fonction de la nature des roches dont ils proviennent, en particulier dans le cas des evaporites. Les chlorures se montrent egalement independants des fluorures et des sulfates.

Role des apports fumerolliens L’influence des apports fumerolliens sur la composition des eaux est tr&s marquee par la liaison entre teneur elevee en bore et en C 0 2 et celle existant entre fluomre et soufre. Les apports gazeux lies aux grandes failles et aux montees volcaniques plus marques en Azerbaidjan que dans le Massif Central le sont peu sauf peut-6tre A Salins et Plan de Phazy dans le domaine alpin, o t ~la presence du bore pourrait etre liee A la nature m6me des evaporites.

Role de la temperature La temperature des eaux en profondeur et aux emergences determine la solubilite des sulfates e t halogenures et celle des composes du bore, mais ces elements ne peuvent &re utilises comme geothermom&tresfaute de blocage des teneurs acquises en profondeur lors de la remontee vers la surface (Michard, 1979). La teneur (generalement assez basse) en cations liberes par la decomposition des mineraux, en particulier des feldspaths, sous l’action de la temperature des eaux en profondeur varie relativement peu lors de cette remontee. Les tentatives de correlation des variations des differents rapports evoques avec les variations de temperature aux emergences ou calculees par les geothermom6tres classiques ne se sont averees concluantes que pour une ou deux regions au plus. Ces correlations n’ont donc pas de valeur generale, en raison certainement des differances dans la nature des roches traversees par les eaux e t des conditions d’equilibre resultant d’eventuels melanges avec les eaux superficielles. I1 conviendrait d’ailleurs avant tout essai de correlation de tenir compte du sodium present sous forme de NaCl (Schoeller e t Schoeller, 1979). Le rapport Li+/Rb+ est le plus sensible aux variations de temperature. Le rapport Sr2+/Rb+,temoin du rapport entre alcalinoterreux et alcalins, parait l’etre peu et doit traduire surtout la nature des roches traversees par les eaux.

Role de la composition minkralogique des roches traverskes e t des mklanges Les differences de nature des mineraux avec lesquels les eaux sont en con-

48

tact expliquent que malgre de grandes analogies de situation geologique on ne retrouve pas les memes rapports ioniques pour les eaux issues des roches cristallines et volcaniques d’Iran ou du Massif Central franqais. Malgre quelques cas particuliers, les eaux ayant lessive les evaporites paraissent caracterisees par des rapports Sr2+/Rbf superieurs B 50 e t Br-/Isuperieurs ou egaux a 1000, le rapport B/F- etant inferieur B 3 e t le rapport SO:-/- voisin de 10.000 etant plus constant que le rapport Cl-/I-. Ces caractbres, lies B la nature meme des sediments evaporitiques, permettent de differencier les eaux qui en sont issues de celles provenant du socle cristallin ou volcanique qui, en Iran comme dans le Massif Central, montrent des rapports Sr2+/Rb+compris entre 1 e t 20 mais des rapports Br-/I- voisins de 100 pour les eaux du Massif Central e t variant de 5 B 160 pour la region du Sabalan. La nature des mineraux avec lesquels les eaux sont en contact intervient d’ailleurs pour des elements tels que F, B ou Li qui peuvent 6tre abondants dans des contextes petrographiques varies. Si en effet les fluomres semblent souvent d’origine fumerollienne ou en relation avec des gisements de fluorine (L’Echaillon, Vichy) la presence de fluomres dans les eaux d’Aix-les-Bains peut etre due 51 la seule presence de mineraux du type fluoapatite dans les sediments albiens. Le bore peut egalement provenir de la mise en solution en milieu acide des borates bien representes dans les formations marines: carbonates, schistes ou evaporites, mais il faut constater qu’il est le plus souvent associe au COz d’origine profonde. Certaines teneurs elevees en Li (Saint-Gervais, Brides, La Lechere) peuvent etre dues a des conditions particulieres de sedimentation des evaporites ou de lessivage de mineraux tels que les micas du socle cristallin, mais il est possible qu’il existe aussi une relation entre les fortes teneurs en Li et les fortes teneurs en CO, comme on le constate surtout en Azerbaidjan oriental. Si la nature des formations traversees par les eaux lors du circuit thermal determine pour une grande part leur composition, il convient d’insister i3 nouveau sur l’importance des equilibres chimiques conditionnes par la temperature, le pH, le Eh e t les forces ioniques des solutions sur cette composition qui peut etre profondement modifiee lors de melanges avec les eaux semisuperficielles ou superficielles. Le fait que de tels melanges ne soient pas uniquement binaires rend particulikrement complexe l’etude des phenombnes hydrothermaux.

CONCLUSION

Le grand nombre de parametres intervenant dans la composition des eaux thermominerales: apports fumerolliens, temperatures en profondeur, nature des roches traversees, equilibres chimiques, pratiquement tous lies B des conditions locales, explique la variete des diagrammes representatifs de cette

49

composition dans les diverses regions etudiees, au sein d’une meme region ou d’une meme station thermale. La prise en compte des teneurs en halogenures e t elements traces dans les eaux complete toutefois utilement l’etude de leurs elements majeurs. La mise en evidence de l’independance entre les chlomres et les autres halogenures en est un bon exemple. S’il semble difficile de definir un profil unique de composition des eaux issues de chaque type de formation, et bien qu’un certain nombre de caracteres soient communs aux eaux issues de certaines formations, il est delicat de definir des geothermometres chimiques utilisables sans tenir compte du type des formations geologiques traversees par les eaux thermales. Une connaissance approfondie des conditions geologiques locales s’avere essentielle comme pour tous les problemes hydrogeologiques et l’etude de la repartition des teneurs en Sr2+, Rb’, Li’, F-, Br-, I-, B devrait permettre de mieux connaitre l’origine des elements acquis en profondeur par les eaux thermominerales, comme les analyses isotopiques ont permis de preciser l’origine de l’eau elle-meme.

BIBLIOGRAPHIE Blavoux, B., Dazy, J. et Sarrot-Reynauld, J., 1980. Information about the origin of thermomineral waters and gas by means of environmental isotopes in eastern Iranian Azerbaijan and the southeast of France. 2eme Congr. Geol. Int., Paris, Symp. Geothermie. Bosch, B., Dazy, J., Lepiller, M., Marc&, A., Olive, Ph., Poulain, P.A. et Sarrot-Reynauld, J., 1976. Donnees nouvelles sur quelques sources thermomindrales des Alpes franqaises. Proc. Int. Congr. Thermal Waters, Athens, 2: 32-40. Cailleaux, P., 1976. Les elements traces dans les eaux thermominerales du Massif Central. These 3eme cycle, Universite Paris VI, Paris. Cailleaux, P., Fouillac, C., Michard, G. et Ouzounian, G., 1976. Etude gkochimique des sources thermales de Chaudes-Aigues. C.R. Acad. Sci., Paris, Ser. D, 292: 1237-1240. Cormy, G., Demians d’Archimbaud, J. et Surcin, J., 1970. Prospection gkothermique aux Antilles Franqaises, Guadeloupe et Martinique. Geothermics, Spec. Iss., 2(Part I): 557 -597. Dazy, J. et Grillot, J.C., 1981. Relations entre thermomineralisme, geochimie des reservoirs et structures tectoniques perialpines. Exemple de la region d’Aix-les-Bains (Savoie, France). Rev. Geol. Dyn. Geogr. Phys. (A paraitre). Dazy, J., Olive, Ph. et Rochat, J., 1980. Nouvelles donnees geochimiques et isotopiques sur les eaux thermales d’Uriage-les-Bains. C.R. 105eme Congr. Natl. Soc, Sav., Caen. Degranges, P., Pepin, D., Risler, J.J., Baubron, J.C. et Bosch, B., 1978. Etude chimique et isotopique de l’eau minerale et des gaz thermaux de Royat (Puy-de-Dbme). Cah. Artoriol. Royat, 6 : 14-25. Didon, J. et Gemain, Y.M., 1976. Le Sabalan, volcan plioquaternaire de 1’Azerbaidjan oriental (Iran). These 3eme cycle, Universite Grenoble I, Grenoble. Henou, B., 1973. Les sources minerales et thermales du Cantal: leur cadre geologique. These 36me cycle, Universite de Clermont, Clermont. Jacob, P., 1975. Contribution a l’etude geochimique des eaux froides et thermominerales du massif du Mont-Dore. These 3eme cycle, Universite Paris VI, Paris.

50 Jeantin, M., 1979. Les sources minerales carbogazeuses de la Haute-Ardeche. Etude geologique et hydrogeologique. These 3eme cycle, Universite de Grenoble, Grenoble. Letolle, R., 1967. Les laves du Mont-Dore. Recherches sur la geochimie et la genese des magmas. These Doct. Sci., Universite de Paris, Paris. Maisonneuve, J. et Rider, J.J., 1979. La ceinture perialpine carbogazeuse de 1’Europe occidentale. Bull. Bur. Rech. Geol. Min. (Fr.), 2eme Ser., Sect. 3, No. 2, pp. 109-121. Michard, G., 1979. Geothermometres chimiques. Bur. Rech. Geol. Min. (Fr.), 2eme Ser., Sect. 3, No. 2, pp. 183-189. Michard, G., Stettler, A., Feuillac, C., Ouzounian, G. et Mandeville, D., 1976. Subsuperficial changes in chemical composition of thermomineral waters of the Vichy Basin. Geochem. J . , 1 0 : 155-161. Michard, G., Eward, M., Fouillac, C. et Lambret, B., 1978. Acquisition des ions alcalins terreux par les eaux thermominerales carbogazeuses. Earth Planet. Sci. Lett., 34: 170-1 74. Ouzounian, G., 1978. Etude des elements en traces dans les eaux thermominerales du Sud de la France. These 36me cycle, Universite Paris VII, Paris. Poulain, P.A., 1977. Les eaux minerales et thermominkrales dans le dkpartement des Hautes-Alpes. These, Universitd de Grenoble I, Grenoble. Rochat, J., Salehi, E. et Sarrot-Reynauld, J., 1979. Geochimie des eaux thermominerales. Repartition et origine des halogenures et des elements traces dans les eaux thermales d’Azerbaidjan oriental (Iran). C.R. 1046me Congr. Natl. SOC.Sav., Bordeaux, Fasc. 111, pp. 123-134. Salehi, E., 1978. Sources thermominerales et anomalies gkothermiques en Azerbaidjan oriental. These Doct. Sci., Universitk Grenoble I, Grenoble. Salehi, E. et Sarrot-Reynauld, J., 1977. Determination des taux de melange des eaux thermominerales et superficielles par utilisation des geothermometres chimiques en Azerbaidjan oriental (Iran). C.R. l02eme Congr. Natl. SOC.Sav., Limoges, Fasc. 11, pp. 191-202. Sarrot-Reynauld, J. et Simeon, Y., 1979. Tectonique et anomalies geothermiques. Donnees nouvelles sur les sources thermominerales de Tarentaise: Brides, Salins et La Lechere (Savoie). C.R. 104eme Congr. Natl. SOC.Sav., Bordeaux, Fasc. 111, pp. 111122. Sarrot-Reynauld, J., Poulain, P.A. et Marc&, A., 1977. Tectonique et anomalies g6othermiques. Les sources thermominerales des bordures orientales et meridionales du massif du Pelvoux. Geol. Alp., 53: 75-82. Sarrot-Reynauld, J., Salehi, E. et Blavoux, B., 1978. Caractkristiques isotopiques et origine des eaux thermominerales d’Azerbaidjan oriental (Iran). C.R. 103eme Congr. Natl. SOC.Sav., Nancy, Fasc. IV, pp. 141-152. Schoeller, H. et Schoeller, M., 1979. Une etude des eaux thermominerales du Massif Central (France). Bull. Bur. Rech. Ge‘ol. Min. (Fr.), Sect. 3, No. 2, pp. 121-156. Sirneon, Y., 1980. Etude hydrogeologique des sources thennominerales de Tarentaise: Brides-les-Bains, Salins-les-Thermes, La Lechere. These 3eme cycle, Universite de Grenoble, Grenoble. Terrier, Cl., 1980. Le Mont-Dore. Ses sources thermales. Son environnement climatologique. These Pharmacie, Universite Grenoble, Grenoble. Thuizat, R., 1973. Les sources thermominerales de Chatel-Guyon et leur environnement geologique (Massif Central franqais). These 3eme cycle, Universite de Clermont, Clermont. Vesely, J., Weiss, D. et Stulik, K., 1978. Analysis with ion selective electrodes. Ellis Horwood Ltd., New-York, N.Y., 245 pp.

51

THE ORIGIN AND EVOLUTION OF SALINE FORMATION WATER, LOWER CRETACEOUS CARBONATES, SOUTH-CENTRAL TEXAS, U.S.A. LYNTON S. LAND and DENNIS R. PREZBINDOWSKI Department of Geological Sciences, University of Texas a t Austin, Austin, TX 78712 (U.S.A.) Amoco Production Company, Tulsa, OK 74102 (U.S.A.) (Accepted for publication April 16, 1981)

ABSTRACT Land, L.S. and Prezbindowski, D.R., 1981. The origin and evolution of saline formation water, Lower Cretaceous carbonates, south-central Texas, U.S.A. In: W. Back and R. L6tolle (Guest-Editors), Symposium on Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54: 51-74. Systematic chemical variation exists in formation water collected from a dip section through Lower Cretaceous rocks of south-central Texas. Chemical variation can be explained by an interactive water-rock diagenetic model. The cyclic Lower Cretaceous shelf carbonates of the Edwards Group dip into the Gulf of Mexico Coast “geosyncline”, and can be considered, to a first approximation, as part of a complex aquifer contained by Paleozoic basement beneath, and by relatively impermeable Upper Cretaceous clay and chalk above. The hydrodynamic character of this carbonate system is strongly controlled by major fault systems. Major fault systems serve as pathways for vertical movement of basinal brines into the Lower Cretaceous section. Formation water movement in this sytem has strong upfault and updip components. The “parent” Na-Ca-C1 brine originates deep in the Gulf of Mexico basin, at temperatures between 200 and 25OoC, by the reaction: halite f detrital plagioclase f quartz f water

+

albite f brine

Other dissolved components originate by reaction of the fluid with the sedimentary phases, K-feldspar, calcite, dolomite, anhydrite, celestite, barite and fluorite. Significant quantities of Pb, Zn and Fe have been mobilized as well. As the brine moves updip out of the overpressured deep Gulf of Mexico basin, and encounters limestones of the Stuart City Reef Trend (the buried platform margin), small amounts of galena precipitate in late fractures. Continuing to rise upfault and updip, the brine becomes progressively diluted. On encountering significant quantities of dolomite in the backreef facies, the Ca-rich brine causes dedolomitization. Although thermochemical consideration suggests that small amounts of several authigenic phases should precipitate, most have yet to be found. Minor amounts of several kinds of calcite spar are present, however. As the brine evolves by dilution and by cooling, no systematic changes in any cation/Cl ratio occur, except for regular updip gain in Mg as a result of progressive dedolomitization. The formation water, highly diluted by meteoric water, eventually discharges along faults as hot mineral water.

52

INTRODUCTION

The origin of saline formation water in sedimentary basins has been problematical since it was first recognized that basinal fluids typically contain dissolved solids in concentrations considerably in excess of seawater. Vast differences in major-ion ratios quickly dispelled early assumptions that basinal fluids were connate and represented buried seawater (Chave, 1960). Since then, different mechanisms have been advocated to account for the composition of subsurface water, and indeed, different mechanisms probably operate in basins with different lithologies and different burial histories. In some cases saline formation water may evolve in near isochemical rock-water systems during burial, as increasing temperature and pressure induce reactions which transfer components from the solid t o the dissolved state. At the other end of the spectrum, fluid bearing no resemblance to the interstitial burial water may be imported from another part of the basin, or even from outside the basin, for example, by meteoric recharge, and modified by rock-water interaction. The primary goal of our study was to document the burial diagenetic history of the Lower Cretaceous Edwards Group in south-central Texas with emphasis on the most deeply buried rocks. Our basic questions were: How has burial affected the rocks? Does rock chemistry or mineralogy record early or late (burial) diagenetic events? What percentage of cementation (or any other diagenetic reaction) is due t o burial? In attempting t o answer these questions we systematically sampled formation water from 43 producing oil and gas wells from eleven fields in five Texas counties. Thirteen additional samples were obtained from updip water wells. The present paper reports our water analyses, and a model which, we believe, best accounts for our observations. We do not pretend to fully understand this gigantic aquifer, much of which is inaccessible, and preface our remarks with the recommendation that more data are badly needed. Study area

In Early Cretaceous time a carbonate platform existed over most of what is now Texas, fringed by a “reef” complex (bank edge) which separated extensive areas of shallow-water carbonate sedimentation from the ancestral Gulf of Mexico. Between 300 and 1000 m of Lower Cretaceous shallow-water limestone, dolomite and evaporites accumulated over Paleozoic and Lower Mesozoic rocks, and was subsequently covered by clay and chalk deposited by Late Cretaceous transgressions. Burial of the platform occurred as thick clastic deltaic complexes prograded over it throughout much of the Tertiary. Today, the bank edge (the Stuart City “reef”) is buried to depths of between -3.7 and 4.7km in the study area. Fig. 1 is a map of the study area showing the major zones of faulting and the oil and gas fields from which samples have been obtained. Because of the distribution of producing oil and

53

Fig. 1, Map of the study area showing the position of the Stuart City Reef Trend and the producing fields which were sampled. Lower Cretaceous peritidal sediments are exposed in outcrop today northwest of the Balcones Fault zone, whereas the Stuart City Trend (the bank edge) is buried today to depths of between 3700 m (southwest) and 4700 m (northeast) in the study area.

gas fields, collection of samples from a complete dip section and at all depths in the section is not possible. As a result, our samples are concentrated at five distinct depth intervals, i.e. -4200m (the Stuart City shelf edge), 3200m (the Karnes Trough), 2100m (the Atascosa Trough), 780m (the Luling Fault zone), and shallower depths (water wells). All areas of production are associated with fault trends, and all samples come from the uppermost zones of the thick “aquifer”. Sampling and analysis Formation water was obtained from producing wells as close to the wellhead as possible, either from the first separator or from the well-head itself. Only the most prolific water-producing wells in each field were sampled. In the case of gas wells, considering gas production, and temperature and pressure changes during production, less than 5% of the water produced is calculated t o be distillate.

54

At each well, three samples were taken whenever possible. A 500-ml sample, acidified with HNO,, was taken for chemical analysis and 125-ml samples were taken for 6l8O and 6D analyses (untreated) and for 6 13C analysis (added t o 5 ml of saturated SrC1, and 1 0 ml of 5 N NaOH). In addition, as much information as possible was obtained on production rates of oil/gas and water, well manipulation (fracturing, acidizing, etc.), well treatment, scaling problems, and previous water analyses. Well scale was obtained wherever possible. Although innumerable possibilities exist for errors in sampling of this sort, and it is very difficult, if not impossible, to determine in situ formation conditions, the data presented and discussed (Table I) represent an internally consistent data set. In the course of our sampling we typically obtained erratic results from gas wells producing less than -1.59 m3 (10 bbl) of water per day, or wells from which gas aspirated during collection. Water from these wells usually had enriched 6D-values and reduced total dissolved solids (TDS) content. In the case of three samples where several wells connected to a single separator, or wells in which multiple completions were indicated, one or more components (especially K) had aberrant concentrations relative to nearby samples. We present these data in Table 11, but have not included them in our discusion. Wells which were not producing water more-or-less continuously, or which had been extensively manipulated (acidized or treated) were not sampled. No water-flooding is used in this area, and we have no reason t o suspect that the analyses reported in Table I represent anything other than the ion ratios of in situ water. Because of the care taken during sampling, we believe the ionic strengths are representative as well. Analytical results

Fig. 2A-D presents the results of the cation analyses plotted vs. chloride, using a different symbol for each depth interval sampled. Sulfate is very dilute in these waters, less than 300mg/l, and was not routinely analyzed. Alkalinity is also very low, and determination of bicarbonate is complicated by the presence of organic alkalinity (Carothers and Kharaka, 1980) and loss of CO, as the water is produced. We made no attempt t o determine in situ carbonate saturation states. Regarding the major cations, and chloride, with the exception of Mg (discussed below), the relative ionic proportions of water deeper than -2 km are approximately constant. We observe no statistically significant differences in Ca/C1 or Na/C1 ratios among any samples from this area. The Na/C1 ratio is neither that of halite nor that of seawater (Fig. 2). Sr and K do not correlate as well with chloride, in part due to more severe analytical difficulties. There is a suggestion, especially in the 3200and 2100-m depth intervals, that some shallower water contains slightly less Sr and K relative t o chloride than does deeper water and that a correlation with depth may exist.

General Crude, Kenedy Field, Karnes C o u n t y :

A. Wernli No. 1 J. Polson No. 1 E. Rolf No. 1 C. Strawn No. 1

13,076-13,253 13,033-13,288 13,065-13,261 13,384-13,452

157 162 160 157

1.213 1.052 1.061 1.213

161 163 174 168

1.075 1.071 1.049 1.052

118

4,800 340 890 5,600

24.5 6.77 8.53 28.6

830 360 370 800

4,390 868 1,090 4,980

173.0 43.5 51.5 189.0

5.45

229 156

74.0 19.3 21.5 77.5

370 >800 240 366

30.5 26.0 17.6 22.2

1,400 2,220 610 510

7.39 9.08 6.93 3.98

340 330 201 220

1,610 1,240 615 1.110

55.2 37.5 57.2

3,370 2,800 3,470

15.6 26.2 17.7

630 520 660

63.0 56.5 52.0 64.0

4,200 3,400 2,900 4,400

20.1 17.6 14.9 20.5

2,800 1,400 2,900 4,400 1,400

10.8 8.9 13.9 18.9 10.3

2,050

-

4.95 5.45

f12.8 f3.4 +11.4 +9.7

-17 -27 -21 -17

64.4 59.1 41.0 44.7

6.25 5.45 5.75 6.7

+13.9 f9.5 +13.6 +14.7

-22 -20 -21 -15

2,740 1,800 3,210

123.6 111.3 132.9

4.7 6.2 4.55

+14.4 +29.6 +18.0

-26 -18 -6

580 470 620

1,670 1,580 1,380 1,710

2,380 2,240 1,880 1,310

145.0 125.0 114.3 146.0

6.2

-9 -7 -14 -21

881

5.1 5.9

f10.4 +8.5 +10.7 +10.3

830 760 1,220 1,330 900

1,650 2,170 2,950 3,250 2,400

86.6 71.5 110.6 140.0 83.2

5.85 5.8 5.8 6.05 6.6

+6.7 f8.5 4-12.3 +12.9 +13.0

+10 -10

-

556

1,120

47.0

4.75

+13.1

-

701 -

858

Shell Oil, Pawnee Field, Bee County:

A. Gordon No. 2 A. Olson No. 1 J.A. Leppard No. 1 S.E. Turner No. 2

13,624-14,002 13,446-13,922 13,804--13,898 13,716-13,971

241 -

177

M.G.F. Oil Co., Monteola Field, Bee County: Ruhmann No. 1 Ruckman B No. 1 Schulz No. 1

13,440--13,666 13,575--13,645 13,424-13,587

176 170 171

1.141 1.131 1.153

-

114 132 137 133

1.165 1.147 1.133 1.165

296

123 125 121 129 118

1.094 1.083 1.127 1.158 1.096

246 386 180

35.9 32.0 51.0 63.5 38.0

1.059

-

20.1

-

Siiell Oil, Person Field, Karnes C o u n t y : Y. Cisneros No. 1 A. Dugie No. 2 Kruse No. 1 C. Kainer No. 2

10,935-10,960 10,904--10,918 10.906--10,920 10,953--10,980

-

48 260

-

-

-

Gulf Oil, Fashing Field, Atascosa C o u n t y : Emma Tartt No. 1 W.T. Hurt No. 3R W.T. Hurt No. 2U W.T. Hurt No. 1U E.S. Koehler No. 1 U

10,817--10,835 10,524-10,822 10,600-10,664 10,712-10,772 10,628--10,712

80

-

-15 -15 -14

-

-

Exxon, San Miguel Field, McMullen C o u n t y : L.M. Gubbels No. 1

10,149-10,182

112

5.63

-

-

cn cn

TABLE I (continued) Production depth (ft.*)

Well

Density

HCO2

Na

K

Ca

ME

Sr

C1

pH

6l80

6D

Br

89 99 116 89

1.134 1.149 1.139 1.132

334 739 >BOO >BOO

53.0 56.0 53.8 51.5

2,300 2,950 2,475 1,885

15.8 17.5 16.3 15.7

2,150 2,350 2,250 2,350

768 925 890 1,750

119.0 128.0 117.0 114.0

5.9 4.95 5.3 5.6

+9.5 $12.5 f9.2 +10.1

-18 -18 -13 -15

732

-

-

63.3 56.5 49.0

3,190 2,340 1,175

17.5 16.6 15.0

2,270 2,210 2,060

2,150 2,300 2,350

134.0 125.0 83.8

5.25 5.1 5.2

+10.8 +10.6 +10.4

-14 -18 -20

-

83

1.153 1.142 1.132

81 74

1.077 1.090

-

32.5 33.0

605 875

6.33 7.50

1,360 1.526

390 355

63.8 68.7

5.55 5.75

f6.4 f6.5

-16 -22

441

-

1.059 1.053

-

25.1 23.1

281 253

3.53 2.26

890 647

490 320

47.6 42.6

5.85 6.6

+5.2 +5.2

-24 -20

-

1.019 1.020 1.019 1.018 1.019

730 650 514 316

260 260 250 250 260

1.77 1.67 1.71 1.58 1.63

580 550 550 540 540

53 53 51 49 69

14.4 13.9 14.2 13.5 14.8

6.5 6.7 6.7 6.5

-2.5 -3.0 -2.3 -1.7 -3.6

-27 -26 -22 -17 -23

Termerature ("C)

-

Exxon, Jourdanton Field, Atascosa County: S.P.J.S.T. Lodge No. 1 S.P.J.S.T. Lodge No. 2 J. Sandeen No.'s 1 and 2 A.N. Moursand No. 2

7,386-7,392

-

7,35+7,396 7,380-7,384

-

Exxon, Imogene Field, Atascosa County: Duren and Richter No. 5 Duren and Richter No. 2 Coward and Couch No. 1

7,568-7,573 7,599-7,606 7,455-7,451

-

-

-

Exxon, Charlotte Field, A tascosa County: J.T. Uppright No. 1 E.J. F'ruitt No. 36

6,996-7,000 6,44*6,572

-

Exxon, Horn Field, Atascosa County: E.J. F'ruitt No. 8 E.J. Pruitt No. 7

6,921-6,929 6,9174,927

Gulf Oil, Darst Creek Field, Guadalupe County: Dix and McKean No. 31 Thomas Dix No. 1 C. Knobloch No. 1 Anderson No. 1 Thomas Dix No. 22

* 1 ft. = 0.3048 m.

2,548-2,562 2,633-2,635 2,613-2,616 2,657-2,659 -

52 49

50

-

6.99 6.63 6.71 6.50 7.17

-

85 -

-

TABLE I1 Chemical analyses of oil-field water with one or more components having aberrant concentrations, south-central Texas Well

Production depth (ft.*)

Temperature ("C)

Density (g/cm3)

HC03 (mg/l)

Na (g/l)

K (mg/l)

Ca @/I)

Mg (mg/l)

-

-

-

540

3.32

-

Sr (mg/l)

C1 k/l)

pH

6"O

6D

(o/oo)

(%o)

Br (PPm)

General Crude, Kenedy Field, Karnes C o u n t y : McDowell No. 1

15,655-15,814

194

1.013

350

-

166

-

-

5.55

f0.9

+2

-

860

37.2

6.55

-

f11.8 4-14.2

-15 -21

-

Shell Oil, Pawnee Field, Bee C o u n t y :

E.P. Benham No. 1 S.E. Turner No. 1

-

1.044

-

291

18.7

-

148

1.098

450

31.4

850

14.1

800

4,000

81.2

5.7

4-13.1

-26

-

-

1.076

-

27.2

1,940

8.6

445

1,330

64.0

5.0

4-18.0

-25

-

134 130 132

1.079 1.155 1.121

180

-

22.2 56.5 47.0

250 3,400 3,700

17.2 21.8 13.9

520 1,460 960

810 2,230 1,970

66.5 133.0 106.0

6.1 4.2 5.7

-4 -8 f12

265

flO.O 4-6.4

1.063

256

22.4

1,050

7.23

1,030

480

52.7

5.9

f3.2

1.022

620

-

-

-

-

-

-

6.6

-2.9

13,690-13,778

Shell Oil, Buchel Field, Bee County: Hay No. 1 -

-

-

-

-

174

-

-

M.G.F. Oil C o . , Monteola Field, Bee County: Boone No. 1

13,440-13,696

Shell Oil, Person Field, Karnes County: T. Yanta No.'s 1 and 2 C. Wishert No. 1 L. Urbanczyk No. 1

10,890-10,916 10,930-10,960 10,925-1 0,938

4-73

-

Exxon, Jourdanton Field, Atascosa County: O.H. Pfeil No. 2

7,399-7,403

78

-

344

G u l f Oil, Darst Creek Field, Guadalupe C o u n t y . Dix and McKean No. I 1

* 1ft. = 0.3048 m.

2,533-2,548

49

-25

-

58

The relation between Mg and chloride is most informative. Because the correlation between Ca and chloride is excellent (Fig. 2B), we present the data as a scatter plot of Mg against Ca (Fig. 3). It must be stressed again that our “depth intervals” are imposed by sampling limitations, and samples are unavailable at intermediate depths in areas where no oil or gas production occurs. Clearly, Fig. 3 shows that shallower waters are progressively and systematically enriched in Mg relative t o Ca (or chloride). Each “depth interval” has a characteristic Mg/Ca (or Mg/C1) ratio, and that ratio varies systematically in the dip section. Fig. 4 presents a scatter plot of 6D vs. depth. All water is depleted in deuterium relative to SMOW, and there is a tendency for samples t o be

.. 1 -

... Y

.

L

0 0

2

1

4

3

5

6

mCI-

.

0.8-

.

-

0.6

..

...: ’ *.

mCa2’0.4-

0.2.

u

. .. . . .I

.

I

0 0

1

2

3

4

5

6

59

.

..

.

."...

0

0

-

2

L

*

:=

.. .. .

1

.

3

4

5

6

mCI0.08

.

0.06.

?3r2+

0.04.

0.02.

...-

....

.

.

.

.

*'

AA

O ! 0

1

2

3

4

5

-

6

r"CI-

Fig. 2. Scatter plot of molality of: (A) sodium; (B) calcium; (C) potassium: and (D) strontium, respectively, vs. molality of chloride. Circles depict samples from -4200 m (Stuart City Trend), squares depict samples from -3200 m (Karnes Trough), and triangles depict samples from -2100 m (Atascosa Trough). Note that the deepest samples are the most saline, but relatively low salinity samples can be found a t any depth. The Na/Cl ratio of the brine is constant, and less than that of seawater. Note also that the Ca/Cl ratio is essentially constant.

somewhat enriched with increasing depth. Within each of the three deeper zones, there is a suggestion that shallower wells are slightly depleted relative to deeper wells, but no parallel relationship with chlorinity exists. The 6 l80 of the water is buffered by the Edwards limestone, as originally explained by

60

mCa2+

0.8

-

0.6

-

0.4

-

. .

. Fig. 3. Scatter plot of molality of calcium vs. molality of magnesium. Symbols as in Fig. 2. Note that unlike previous data, each depth interval has a consistent and unique Mg/Ca (or Mg/C1) ratio. The lines are best fits to the data poidts, and using the slopes of the lines and the data of Rosenberg and Holland (1964),an equilibrium temperature for calcitedolomite has been calculated. In all cases the calculated temperatures are much hotter than the reservoirs today or at any time in their past. For the three depth intervals, reservoir temperatures today are -165OC (circles), 126OC (squares) and 89OC (triangles). Therefore all water is undersaturated with dolomite.

-301:

::*

. . .. r

H

8m

c

-10 .

0

. . .. .

.

?

1

2

DEPTH

3

4

5

(km)

Fig. 4. Scatter plot of 6 D vs. depth. Note that all samples are depleted relative to SMOW, but most are slightly enriched relative to local meteoric water today (6D = -21Tm). Note also that shallower water in each depth interval (fault trend) is in cases depleted relative to deeper samples in the same fault trend.

61

Clayton et al. (1966), and the 6’*0 of the water in almost all cases is in equilibrium with the limestone at the reservoir temperature (Prezbindowski, 1981). N o useful information relative t o the origin of the water can be obtained from oxygen-isotope data except to be sure that the water has been in contact with the rocks long enough t o have equilibrated with them. Discussion In trying to account for the data presented above, we considered three possible “models ” : Model 1 . Meteoric water is recharging the outcrop of the aquifer, moving downdip, and evolving in chemical composition as rock-water interactions take place a t increasing depths and temperatures. Model 2. Connate water (alternatively, Cretaceous seawater, a Cretaceous sabkha brine, or Cretaceous meteoric water) has evolved t o the present formation water compositon by closed-system burial diagenesis, or diagenesis in a system permitting loss of material by compaction, but not addition. Model 3. Brine is formed in the deep Gulf of Mexico by reaction of deep basinal water with Jurassic evaporites, is injected into the aquifer at depth, and moves updip. We have accepted model 3 as best accounting for the observed facts, and present our discussion accordingly. Fig. 3 presents, we believe, a compelling argument for this model. Because this is a carbonate aquifer, the Mg/Ca ratio of the water is certainly related t o the calcite-dolomite system. Clay constitutes only a few percent of the rock section. Dolomite is abundant in outcrop of the Edwards Formation, and decreases in abundance into the subsurface, being limited in distribution in cores deeper than 1km. Water shallower than -1 km approaches equilibrium with calcite plus dolomite using the Kdolomite-valuesof Langmuir (1971) or Hsu (1963) (Pearson and Rettman, 1976). A different situation exists with respect to the deeper water, however, using the relation: log(mCaz+/mMgz+) = -1000T-’

+ 2.98

where T is in kelvins, derived from experimental data in 2 M divalent cationchloride (MC1, ) solutions at elevated temperatures (Rosenberg and Holland, 1964). We assume, as did Rosenberg and Holland (1964), that Y ~ ~ z +Y M ~ Z + . Since the ionic strength of Rosenberg and Holland’s experimental solutions (6) is similar t o or greater than ionic strengths observed in the Edwards Formation (1.3-53, Rosenberg and Holland’s molality ratios should be comparable to our molality ratios irrespective of the values of the activity coefficients. Using the 1 M data of Rosenberg et al. (1967) (ionic strength = 3) would additionally favor the following argument, namely that all the water deeper than -1 km is grossly undersaturated with respect to dolomite, and the

62

water becomes increasingly undersaturated with increasing depth. Consider the three possible models in light of these data: Model 1. It is impossible for water to move downdip, lose Mg relative to Ca, becoming progressively more undersaturated with dolomite. This model can therefore be ruled out, unless the Mg is being controlled by reactions other than those involving calcite-dolomite. We are unaware of any other Mg-bearing phase of volumetric significance in this carbonate aquifer. Model 2. If water, no matter what its original salinity, was in near equilibrium with calcite dolomite or was oversaturated with dolomite by virtue of a high Mg/Ca ratio at shallow depth (like Cretaceous seawater or a sabkha brine), it could never evolve to its present composition with increasing temperature (burial) by dolomitization because the water is everywhere undersaturated with dolomite today. It is very unlikely that the very low Mg/Ca ratio of the brine could be primary because neither Cretaceous seawater nor any reasonable Cretaceous sabkha brine are likely to have had such a composition. Model 3. Because most of the dolomite in these rocks is associated with the shallower (bankward) parts of the carbonate platform (Rogers, 1967), a Ca-rich water moving updip would encounter progressively more and more dolomitic country rock as it moved updip, and could thus dedolomitize the rock and increase the Mg/Ca ratio of the water. The amount of Mg gained by the water would be limited by the amount of dolomite originally present in the rocks and by the channelization of the flow. Because dolomite is absent in the platform margin, the water could remain undersaturated with dolomite until significant quantities of dolomite were encountered updip. Fig. 5 is a photomicrograph of a sample from the Jourdanton field in the Atascosa Trough. Blocky spar calcite lines vuggy pores in the dolomitic limestone, and the poikilotopic spar encircles many of the dolomite crystals, some of which are extremely corroded. There is a distinct transition zone between the calcite lining the pores and the porous dolomite groundmass making up the bulk of the rock. This transition zone is characterized by an increase in corroded and replaced dolomite crystals as one moves toward the pore space. The average 6 l80of the poikilotopic spar calcite (dedolomite) is -5.3YO0 (PDB) as compared t o 6 l 8 0 of -2.1YoO for the dolomite and 6 ' * 0 of -3.8 Yoo for nearby bulk limestone. These data are consistent with a subsurface origin for the spar, whereas the isotopic composition of the bulk limestone and dolomite is due to early, pre-burial meteoric diagenesis (Prezbindowski, 1981). Both petrography and isotopic data indicate that dedolomitization is the latest diagenetic process t o significantly alter these rocks. We suggest that dedolomitization is being caused by the present-day formation water. A second line of argument concerns the TDS content of the water, which averages 3 m in chloride. Why does the concentration of dissolved solids in the water increase with increasing depth [a common phenomenon in sedimentary basins (Rittenhouse et al., 1969)]? Edwards water is of the Na-Ca-C1 type

+

63

Fig. 5. Photomicrograph of dedolomite (poikilotopic spar calcite) from the Pheil No. 3 well, Atascosa County, 2253 m, white light. Width of field is 940pm. Note that many of the dolomite rhombs are badly corroded where they have been unreplaced by the spar.

(again, very typical of sedimentary basins), dominated by NaCl, which accounts for -70% of the TDS. The only reasonable source for the dissolved Na and chloride in solutions this concentrated is halite. Although other sources of chloride have been proposed, accounting for the vast amount of chloride in the large volume of water in this aquifer from any other source except halite is very unlikely. Processes such as ultrafiltration (clay-membrane filtration, e.g., Hanshaw and Coplen, 1973; Kharaka and Smalley, 1976), suffer overwhelming objections in cases such as this (model 2). The rocks we are studying are carbonates, generally containing less than 1%clay, and nowhere more than 10% clay. Such clay as is present in this section is concentrated in the basal part of the section and increases in abundance updip. The closest adjacent shales are 1km above these deposits, separated from them by impermeable chalk. In addition, the Edwards and all overlying units are normally pressured at the present time. N o evidence of earlier overpressuring exists and the rocks may have never been overpressured since they and the overlying clastics have never been more deeply buried than they are today. If membrane filtration is important then the membranes are far removed from the water and a driving mechanism (overpressure) has apparently always been absent. If membrane filtration is involved, why are essentially constant ionic proportions always observed in the water (except for Mg)? One would expect systematic variation in, say, Ca/Na with increasing salinity if filtration were operative, as certain species (Na') passed through the shale membrane

64

more easily than other species (Ca”) (White, 1965). Yet the Ca/Na ratio of the water is constant (Fig. 2A and B). Although 6D does seem to increase slightly with increasing depth, no correlation with salinity exists and the observed variation is far less than predicted experimentally (Coplen and Hanshaw, 1973). In addition, all water is depleted in deuterium relative t o SMOW [or the Cretaceous ocean (Lowenstam, 1961)]. If the brine represents the residue after filtration of seawater, meteoric water, or a sabkha brine, it should be enriched in deuterium, not depleted. For these reasons, membrane filtration by shales is believed to be of no consequence in this geological system. Could meteoric water, moving downdip, evolve into the Edwards brine (model I)? Although meteoric water plays an important role in diluting upward-moving brine, it cannot evolve into a brine by interaction with the available country rock. In other areas where carbonate coastal plains of similar lithology are known to be recharged by meteoric water, the tremendous increase in TDS as we observe in the Edwards is not fqund. For example, in the Florida aquifer, water at 500 m depth contains only -850 mg/l TDS and the dominant anion is sulfate (Back and Hanshaw, 1970), compared t o 4000 mg/l in the C1-dominated Edwards. In addition, the junction between saline and potable water (the “Bad Water” line just downdip from the Balcones Fault zone corresponding to 1000 mg/l TDS, Fig. 1)is a relatively distinct feature, not a gradation over many kilometers. Because there is no halite associated with the country rocks, chloride supplied by saline water moving updip, not meteoric water moving downdip, is the only possible explanation for the rapid basinward increase in chloride. Isochemical evolution during burial to yield the water we observe in the rocks today is also unlikely. It has been adequately demonstrated that during and soon after deposition, the Cretaceous sediment deposited on the Comanche shelf was massively altered by meteoric diagenesis (Mueller, 1975; various papers in Bebout and Loucks, 1977; Prezbindowski, 1981). The interstitial water in which the sediment was initially deposited (either seawater or, rarely, a sabkha brine) had been replaced by many pore volumes of meteoric water prior to the beginning of burial in the Late Cretaceous. If the TDS content of the average Edwards water today was produced by dissolution of synsedimentary halite by connate meteoric water or even connate marine water (assuming the synsedimentary evaporites somehow escaped Cretaceous meteoric diagenesis), then an average Edwards water today containing -2 g Na+/l would need to have dissolved -2.3 cm3 of halite per liter of water. Assuming an average post-diagenetic porosity of -lo%, then -0.6% of the original rock section would need to have been halite. Bedded halite has never been described or postulated from Edwards carbonates. Only very minor amounts of gypsum are still present, which, together with solution-collapse features, are most commonly associated with restricted areas of the platform and exposed today far updip and in outcrop. N o evaporites whatsoever are associated with the buried platform margin (the Stuart City

65

Reef Trend), and anhydrite or former evaporites (solution-collapse features) are rare in rocks which are buried today t o depths in excess of 2500m, and are localized in the fault troughs. Thus the origin of the Edwards brine from dissolution of synsedimentary evaporites is unlikely. There simply were insufficient volumes of evaporites in the rock section, and what little evaporites were deposited were mostly removed in Cretaceous time prior to burial (Rose, 1972). Carpenter (1978) proposed that a brine from Mississippi, similar t o that described here, was a modified sabkha brine, or a true connate fluid buried along with its cogenetic evaporites. Aside from the arguments already presented, that very little gypsum or anhydrite and no bedded halite are associated with Edwards deposits, other arguments oppose such an origin for the Edwards brine. Volumetrically, such a connate brine would have been insignificant, even if evaporites had been extensive (Garret, 1970; Kinsman, 1971). A connate sabkha brine should have been enriched in deuterium relative t o SMOW, yet all the Edwards brine we analyzed is depleted. Carpenter's model proposes modifications of a connate sabkha brine by mineral-water reactions (principally dolomitization) t o yield the observed formation water. We have already pointed out that the Mg/Ca ratio of the Edwards brine disputes the role of dolomitization in the formation of the brine as the water cannot continue to lose Mg beyond saturation with respect to dolomite as temperature increases with burial. Carpenter uses the Cl/Br of the water as another line of evidence, arguing that the Cl/Br of formation water resembles seawater-derived brine from which considerable halite has precipitated. Conventional wisdom has it that a brine derived from the dissolution of halite will have the Cl/Br of the halite. Such a solution would contain less Br relative to C1 than seawater or a sabkha brine, because of the value of the distribution coefficient : which is -0.035 at 25OC (Holser, 1979b). However, conventional wisdom is apparently flawed in this case. If the partitioning of bromide between halite and brine were reversible and represented a true homogeneous equilibrium, then a solution formed by dissolving halite initially precipitated from seawater should have a Cl/Br exactly that of seawater, not that of the salt. At equilibrium, if the distribution coefficient were a true reversible equilibrium constant, then the solution formed by dissolving halite would be considerably enriched in bromide relative to the salt. In order to investigate the behavior of Br during halite dissolution, we reacted Jurassic halite from a Texas salt dome with distilled water on a shaker table at 25°C. The Cl/Br ratio for the solution after only 18hr. of reaction (13,500) was lower than that of either the halite or the solution immediately after reaction (both 17,000). After nearly eight months the Cl/Br of the solution had fallen to an apparent steady state (SOOO), but did not approach

66

the value to be expected of an homogeneous reversible equilibrium with = 0.035 (600). Similar behaviour of halite on dissolution has been noted by Kuhn (1968) and a similar mechanism has been suggested to account for very low Br concentrations in halite (Holser et al., 1972). Holser (1979b) summarizes some of the kinetic problems that have plagued Cl/Br studies, and which may explain our observations. Therefore the molar Cl/Br ratio of the Edwards brine (160-250, Table I) might be due to simple dissolution of Jurassic halite since: (1)in situ basinal Jurassic halite should contain more bromide than the displaced and recrystallized basin-margin material used in our experiment; and (2) the distribution coefficient of Br in halite increases with increasing temperature (Holser, 1979b). We find arguments based on Cl/Br ratios unsatisfactory at this stage of our knowledge, and recommend that our experimental observations be tested on a more rigorous basis, and at higher temperatures such as exist during burial. A more satisfactory explanation for the TDS content of Edwards brine is that a highly saline fluid originates deep in the Gulf of Mexico by reaction with Jurassic evaporites, moves up-fault and up-dip, and is diluted and undergoes changes in composition by rock-water interaction as it moves.

D

Origin of '>parent" Edwards brine Many explanations have been proposed for the origin of Na-Ca-C1 formation water. In addition to mechanisms which we have rejected (e.g., membrane filtration), Carpenter (1978) proposed that such a brine results from the generalized reaction: halite

+ carnallite + water + limestone

-+

Na-Ca-C1

brine

+ dolomite

His model is clearly inapplicable in this case unless the water is produced at a temperature of at least 320°C (see Fig. 3), the temperature in equilibrium with calcite, dolomite and the Mg/Ca of the deepest brine. At a reasonable average Gulf of Mexico coast geothermal gradient of -30"C/km, this would require -10 km of burial, whereas the Cretaceous section we are studying has never been more deeply buried than it is today (-5 km). In addition, carnallite is a rare evaporite mineral, and Carpenter's stoichiometry requires -10% of the evaporite section t o be carnallite, a figure hardly in accord with the abundance of carnallite in cored evaporites (Holser, 1979a). N o subs.urface dolomitization of these rocks has been detected, and in fact the opposite reaction is taking place (Fig. 5 ) . As a general explanation for Na-Ca-C1 brines, Carpenter's (1978) hypothesis appears t o us to be unsuitable. We propose another reaction, namely: detrital plagioclase

+ halite iwater

+

Na-Ca-C1

brine

+ albite

for the following reasons. First, albitization is a known process in the Gulf of Mexico coast subsurface (Garbarini and Carpenter, 1978; Boles, 1979;

67

Land and Milliken, 1981). Pettijohn (1963) pointed out that albite is the typical feldspar of deeply buried greywacke sequences, and could not have been detrital because pure albite-bearing source rocks are uncommon. But Pettijohn offered no explanation for this materially significant anomaly. Albitization is a plausible explanation. Second, interest in geothermal energy has stimulated research into various aqueous “geothermometers”. One which has been used with success, based ultimately on feldspar equilibria, is the Ca/Na geothermometer. Using the data of Helgeson (1972, fig. 9), or the empirical formula of Fournier and Trusdell (1973), and the ion ratios of Edwards brine from Fig. 2., a temperature of -200°C is calculated. This is a reasonable temperature for updip Jurassic salt in the Gulf of Mexico today (6000m x 30°C/km). If we consider a rather simple case where the components of a brine are generated by successive addition of common sedimentary minerals, halite plus water yields a brine of a fixed composition (mNa+,mcl+, pH, etc.) at constant temperature and pressure. Adding quartz and albite saturates the brine with silica and alumina. Adding the metastable phase intermediate plagioclase (the typical detrital component of sandstones) causes albitization and consequent replacement of Na in the brine with Ca. Predicting the Ca/Na ratio of the resultant brine accurately is difficult, and depends on the availability and reactivity of plagioclase and whether or not the brine remains in contact with halite quartz plagioclase. If a halite-saturated (or near-halite-saturated) brine flowed into a clastic section lacking halite as a phase, Na removed from the brine by albitization would cause undersaturation with respect to halite but Na could not be replaced by continued halite dissolution. At -200°C a brine in equilibrium with albite and K-feldspar should have a Na/K activity ratio (or molality ratio) of -16 (Helgeson, 1972, figs. 6 and 7), less than the value of 26 observed for Edwards brine. Edwards brine is therefore undersaturated with respect to K-feldspars. K-feldspars have been observed to be almost completely removed from Frio sandstones in Brazoria County, Texas, as shallow as 4250m (Land and Milliken, 1981), and we suggest one possible reason for the potassium composition of Edwards brine is that most of the K-feldspars may have been removed from the Jurassic and Lower Cretaceous sandstone aquifers before the Edwards rocks were as deeply buried as they are today. Thus as the Edwards brine forms, insufficient K-feldspars are present in the discharge conduits, and the brine remains undersaturated. We presume a similar situation to prevent the brine from equilibrating with dolomite. We have already pointed out that a brine temperature as high as 320°C is unlikely, and thus the most likely explanation for the Mg/Ca of the brine is that it forms in a dolomite-starved system. The two most saline brines (Wernli No. I and Strawn No. 1 ) contain only 96 and 84ppm SiOz, respectively. If albitization had occurred at -2OO0C, then the water should contain -250ppm Si02 (Crerar and Anderson, 1971).

+

+

68

But three facts suggest that Edwards brines today should contain less dissolved silica than when they initially formed: (1)the most saline brines we sampled are only -3 saturated with halite, and thus have probably been diluted since they formed; (2) cooling to 160°C, the present bottom-hole temperature of the wells, may have caused precipitation of about half the original silica in the deep aquifer; and (3) the solubility of silica in nearhalite-saturated brines may be lowered significantly because of the reduced activity of water. Clearly the Si02 values obtained today do not preclude quartz-saturation a t the time the brine originally formed. Therefore it is possible that a paucity of K-feldspar and dolomite in the pre-Edwards aquifers downdip from our study area is the reason for the undersaturation of the brine with respect t o these phases. It should be noted that K-feldspar, quartz and dolomite are extremely rare, or absent entirely from the deep Edwards carbonates themselves. We suggest that the initial major-ionic composition of the brine is controlled by a tendency toward equilibrium with the common sedimentary phases halite, quartz, plagioclase and K-feldspar. The other components of the brine are probably controlled by reaction with other common sedimentary phases. Since 2mca2+ mNa+approaches mc,-, Ca in the brine cannot be derived from dissolution of anhydrite. Dissolution of anhydrite would result in Ca being balanced by sulfate or bicarbonate (in the case of sulfate reduction). Since the concentrations of both sulfate and bicarbonate are very low, the cause of the high Ca concentration in the brine cannot be anhydrite dissolution. The sulfate concentration in the brine is controlled by the equilibration of anhydrite with a Ca-rich brine, and the bicarbonate by equilibration with calcite and C 0 2 . The high Sr content, which exceeds Mg in the deepest brines, is probably controlled by equilibration with celestite in the evaporite section. Attempts to calculate actual brine compositions using existing algorithms (WATEQ - Truesdell and Jones, 1973; SOLMNEQ - Kharaka and Barnes, 1973; EQ3/6 - Wolrey, 1979) have not been very successful at present writing, owing t o lack of reliable thermochemical parameters in aqueous solutions near halite saturation and in excess of 150OC. Even so, values not too different from the Edwards brine ion ratios are obtained when halite, quartz, albite, anorthite, calcite, anhydrite, celestite and fluorite are “reacted” t o saturation, and K-feldspar and dolomite held below saturation. Three other lines of evidence favor our hypothesis that Edwards brine evolves deep in the Gulf of Mexico basin and moves up-fault and up-dip to its present position. (1)The underlying Sligo carbonate section is known to be overpressured. Pressure gradients in the Sligo Formation range from -1.88.104 t o 1.99*104 Pa/m (0.83 to 0.88 psi/ft.), whereas a gradient of 1.15*104Pa/m (0.51 psi/ft.) was measured in the (Edwards) Wernli No. 1 well (Prezbindowski, 1981). The Strawn No. 1 well, which produced the most saline brine we sampled (Table l),flowed at -318 m3 (ZOO0bbl) of water per day. Irrespective of

+

69

. 60C

PPm

4 00

Zn

200

. 0

0

11 0

.

150

200

g CI’/ liter

Fig. 6. Scatter plot of Zn (ppm) vs. C1 (g/l) for five samples from the Stuart City Reef Trend. The most saline well, the Strawn No. 1 , was sampled twice over an interval of 13 months, and both points have been plotted. No other samples had Zn contents greater than 1 ppm.

the measured bottom-hole pressure of the well, a 4084-m column of water having a density of 1.213g/cm3 exerts approximately the same pressure as a 4954-m column of fresh water having the same temperature change with depth. Considerably less than 870m of relief exists between any possible area of updip recharge and the land surface above the Stuart City Trend, and therefore the formation water must have a strong updip component of movement. (2) We find minor amounts of galena associated with late fractures in the Strawn No. 1 well, the well which produced the most saline brine we collected. Dissolved Pb and Zn have been detected in concentrations greater than l p p m only in wells from the Stuart City Reef Trend and not from shallower zones. The most saline wells, the Strawn No. 1 and Wernli No. 1, contain the highest concentrations of heavy metals (Fig. 6), and the Strawn well additionally contains fluorite as scale. Although the data are admittedly sparse, the distribution of both the metallic phases and high dissolved metal values are consistent with an out-of-the-basin fluid movement. The source for the metals is unknown. Since the metals are associated with the bank margin, not shallow shelf deposits, concentration of the metals by algal mats (Davis, 1977) is not indicated. ( 3 ) Hydrocarbon typing suggests that a significant fraction of the oil being produced from Edwards fields was sourced from Jurassic rocks (Prezbindowski, 1981). Again, out-of-the-basin transport is indicated.

70 CONCLUSIONS

We present our conclusions with regard to the three “models” we previously suggested as possible origins for the Edwards brine: (1)Arguments against brine formation by basinward recharging meteoric water (model I), or modification of connate Cretaceous meteoric water or Cretaceous seawater (model 2): (a) No source of NaCl is known from Edwards shelf carbonates. Very minor amounts of gypsum are known from outcrops of the inland shelf, or from the fault troughs (anhydrite). No evaporites are known to exist or have been postulated to have existed associated with the bank edge where the most saline brine is now found. (b) The “Bad Water” line is a sharp feature, not a gradual downdip increase in TDS. The “Bad Water” brine has preserved textures in the rocks which have been destroyed updip by the active meteoric regime today (Longman and Mench, 1978), indicating that the active meteoric regime has never extended further downdip than it does today. Additionally, the “Bad Water” contains abundant H2 S, and occasionally hydrocarbons, facts inconsistent with a dominant basinward recharge, but consistent with sulfate reduction by upward-moving hydrocarbons on encountering the meteoric regime. (c) It is impossible for water to evolve to the Mg/Ca ratio of the brine by dolomitization, and all brine deeper than -1 km is undersaturated with dolomite and is, in places, dedolomitizing. It is unlikely that any possible connate solution had such a low Mg/Ca ratio as observed in any brine below -1 km. (d) Large volumes of water would be necessary to concentrate chloride from a dilute water by any mechanism, including membrane filtration. Such large volumes of water would have caused extensive isotopic reequilibration of the rocks, yet the rocks largely retain their early diagenetic chemistry (Prezbindowski, 1981). Therefore it is impossible to involve large volumes of isotopically depleted water in any postulated origin for the Edwards brine. (2) Arguments against brine formation by evolution of a connate Cretaceous brine (model 2): (a) No evaporites are known from the bank-edge part of the section where the brine is most saline today [see (l), (a) above]. (b) The Mg/Ca ratio of the brine is too low t o have evolved to its present composition by dolomitization [see (l), (c) above]. (c) The volume of brine buried with the Cretaceous sediment, even if it had escaped the extensive Cretaceous meteoric diagenesis evidenced in the rocks, is insufficient to supply the volume of brine in the rocks today. (d) The amount of carnallite required relative to halite to yield the Ca content of the Edwards brine is much higher than observed in evaporite sections world-wide. Subsurface dolomitization, required t o replace Mg

71

Fig. 7. Schematic north-south cross-section through San Antonio and the Muil Field (see Fig. 1).Jurassic and Lower Cretaceous sediments (the “aquifer”) are shown by the limestone pattern since most are carbonates. They are overlain by Upper Cretaceous clay and chalk. The northern-most outcrops are associated with the Balcones Fault zone near San Antonio. The shallow part of the section is underlain by Paleozoic “basement”, or by Triassic rocks further south. Jurassic salt is shown by solid hatchures. Darkened areas at A , B and C depict the three producing zones, the Stuart City Reef Trend (the carbonate platform margin), the Karnes Trough and the Atascosa Trough, respectively. Note the faulting associated with them. The temperature axis is speculative below “ 5 km.

derived from carnallite by Ca, is unknown in the study area and in fact the reverse reaction is occurring. (3) Arguments in favor of basinal origin of the Edwards brine (model 3 ) : (a) Evaporites are abundant in the deep Gulf of Mexico, and albitization is a documented reaction of major significance. (b) The distribution of Mg in the brine is best explained by dedolomitization as the brine moves updip. Dedolomite has been found, and its isotopic chemistry is consistent with its formation in the present aqueous regime. (c) Galena and dissolved heavy metals are associated only with the most saline brines at the bank edge, indicating a basinward origin, and a model for their origin is consistent with at least some current theories (Banaszak, 1979). (d) Wells in the Stuart City bank edge flow, and the hydrodynamic drive is updip. (e) Hydrocarbons produced from the Edwards Formation have a strong Jurassic component indicating a downdip source. Fig. 7 presents a cross-section of the study area. We interpret the flow paths of the Edwards brine today t o be up-fault and up-dip. We presume the samples from -4.2 km to tap a slightly different source of brine than the shallower samples because of small but systematic differences in 6D (Fig. 6) and Cl/Br ratio (Table I). It would be tempting to account for the distribution of hydrocarbons in this area by hydrodynamic trapping, but we are unconvinced that sufficient water movement can be taking place t o accomplish migration. We face the same dilemma as faced by studies of sandstone diagenesis (e.g., Land and

72

Dutton, 1979), namely that some evidence (in this case the hydrocarbon distribution) favors extensive water movement, whereas other evidence (lack of extensive oxygen-isotopic modification of the country rock) does not. The volume of flow in this gigantic aquifer is, we believe, a major outstanding problem.

ACKNOWLEDGEMENTS

This research was supported by the National Science Foundation, NSF EAR-76 17774, and the Geology Foundation of the University of Texas. We sincerely thank the many oil and gas companies, and especially their production crews and supervisors, for their cooperation and patience in permitting and helping us to obtain water samples. Alden Carpenter provided the heavy metals analyses on samples collected for him at his request. Rick Schatzinger first detected the galena in the Strawn No. 1 core. Karl Hoops performed the chemical analyses and spent considerable time perfecting the bromide analytical method. We thank him for his careful, reliable analytical work. Charles Kreitler kindly pointed out editorial and logical flaws in an early manuscript.

REFERENCES Back, W. and Hanshaw, B.B., 1970. Comparison of chemical hydrogeology of carbonate peninsulas of Florida and Yucatan. J. Hydrol., 10: 330-368. Banaszak, K.J., 1979. A coherent basinal-brine model of the genesis of Mississippi Valley Pb-Zn ores based in part on absent phases. Am. Inst. Min. Eng., Prepr. No. 79-94, 9 PP. Bebout, D.G. and Loucks, R.G. (Editors), 1977. Cretaceous carbonates of Texas and Mexico: applications to subsurface exploration. Bur. Econ. Geol., Univ. Texas, Rep. Invest. No. 89, 332 pp. Boles, J.R., 1979. Active albitization of plagioclase in Gulf Coast Tertiary sandstones. Geol. SOC.Am., Abstr. Progr., p. 391. Carothers, W.W. and Kharaka, Y.K., 1978. Aliphatic acid anions in oil-field waters - implications for origin of natural gas. Am. Assoc. Pet. Geol. Bull., 62: 2441-2453. Carpenter, A.B., 1978. Origin and chemical evolution of brines in sedimentary basins. In: K.S. Johnson and J. Russell (Editors), 13th Annu. Forum on Geology of Industrial Minerals. Okla. Geol. Surv. Circ., 79: 60-77. Chave, K.E., 1960. Evidence on history of sea water from chemistry of deeper subsurface waters of ancient basins. Am. Assoc. Pet. Geol. Bull., 44: 357-370. Clayton, R.N., Friedman, I., Graf, D.L., Mayeda, T.K., Meents, W.F. and Shimp, N.F., 1966. The origin of saline formation waters, I. Isotopic composition. J. Geophys. Res., 71: 3869-3882. Coplen, T.B. and Hanshaw, B.B., 1973. Ultrafiltration by a compacted clay membrane, I. Oxygen and hydrogen isotopic fractionation. Geochim. Cosmochim. Acta, 37 : 22952310. Crerar, D.A. and Anderson, G.M., 1971. Solubility and solvation reactions of quartz in dilute hydrothermal solutions. Chem. Geol., 8: 107-122. Davis, J.H., 1977. Genesis of the southeast Missouri lead deposits. Econ. Geol., 72: 443450.

73 Fournier, R.O. and Truesdell, A.H., 1973. An empirical Na-K-Ca geothermometer for natural waters. Geochim. Cosmochim. Acta, 37: 1255-1275. Garbarini, J.M. and Carpenter, A.B., 1978. Albitization of plagioclase by oil-field waters. Geol. SOC.Am., Abstr. Progr., p. 406. Garret, D.E., 1970. The chemistry and origin of potash deposits. In: J.L. Rau and L.F. Dellwig (Editors), 3rd Symp. on Salt, N. Ohio Geol. SOC.,Cleveland, Ohio, 1: 211222. Hanshaw, B.B. and Coplen, T.B., 1973. Ultrafiltration by a compacted clay membrane, 11. Sodium ion exclusion at various ionic strengths. Geochim. Cosmochim. Acta, 37: 2311-2327. Helgeson, H.C., 1972. Chemical interaction of feldspars and squeous solutions. In: W.S. Mackenzie and J. Zussman (Editors), The Feldspars. Proc. N.A.T.O. Adv. Study Inst., Manchester University Press, Manchester, pp. 184-217. Holser, W.T., 1979a. Mineralogy of evaporites. In: R.G. Bur% (Editor), Marine Minerals. Mineral. SOC.Am., Short Course Notes, 6: 211-294. Holser, W.T., 1979b. Trace elements and isotopes in evaporites. In: R.G. Burns (Editor), Marine Minerals. Mineral. SOC.Am., Short Course Notes, 6: 295-346. Holser, W.T., Wardlaw, N.C. and Watson, D.W., 1972. Bromide in salt rocks: Extraordinarily low content in the Lower Elk Point salt, Canada. In: G. Richter-Bernburg (Editor), Geology of Saline Deposits. Proc. Hannover Symp., UNESCO, Paris, pp. 69-75. Hsu, K.J., 1963. Solubility of dolomite and composition of Florida groundwater. J. Hydrol., 1: 288-310. Kharaka, Y.K. and Barnes, I., 1973. SOLMNEQ: Solution-Mineral Equilibrium Computations. U.S. Geol. Surv., Comput. Contrib. PB 215 899,Progr. No. G204, 81 pp. Kharaka, Y.K. and Smalley, W.C., 1976. Flow of water and solutes through compacted clays. Am. Assoc. Pet. Geol. Bull., 60: 973-980. Kinsman, D.J., 1971. Discussion of: Subsurface brines and the formation of Mississippi Valley-type ore deposits. Trans. Inst. Min. Metall., 80: B61-B63. Kuhn, R., 1968. Geochemistry of the German potash deposits. Geol. SOC.Am., Spec. Pap., 88: 467-504. Land, L.S. and Dutton, S.P., 1979. Reply to Discussion, Cementation of sandstone. J. Sediment. Petrol., 49: 1359-1361. Land, L.S. and Milliken, K.L., 1981. Feldspar diagenesis - Frio Formation, Brazoria County, Texas Gulf Coast. Geology, 9: 314-318. Langmuir, D., 1971. The geochemistry of some carbonate ground waters in central Pennsylvania. Geochim. Cosmochim. Acta, 35: 1023-1045. Longman, M.W. and Mench, P.A., 1978. Diagenesis of Cretaceous limestones in the Edwards aquifer system of south-central Texas: A scanning electron microscope study. Sediment. Geol., 21: 241-276. /0l6 ratios, and strontium and magnesium conLowenstam, H.A., 1961. Mineralogy, 01* tents of recent and fossil brachiopods and their bearing on the history of the oceans. J. Geol., 69: 241-260. Mueller, H.W., 1975. Centrifugal progradation of carbonate banks: a model for deposition and early diagenesis Ft. Terrett Formation, Edwards Group, Lower Cretaceous, central Texas. Ph.D. Dissertation, University of Texas, Austin, Texas, 300 pp. Pearson, Jr., F.J. and Rettman, P.L., 1976. Geochemical and isotopic analyses of waters associated with the Edwards limestone aquifer, central Texas. U.S. Geol. Surv., Rep. Edwards Undergr. Water Dist., 35 pp. Pettijohn, F.J., 1963. Data of geochemistry, chemical composition of sandstones - excluding carbonate and volcanic sands. U.S. Geol. Surv., Prof. Pap. 440-S, 19 pp. Prezbindowski, D.R., 1981. Burial diagenesis, Edwards Formation, Lower Cretaceous, south-central Texas. Ph.D. Dissertation, University of Texas, Austin, Texas, 235 pp. Rittenhouse, G., Fulton, R.B., Grabowski, R.J. and Bernard, J.L., 1969. Minor elements in oil field waters. Chem. Geol., 4: 184-209.

74 Rogers, J.K., 1967. Comparison of some Gulf Coast Mesozoic carbonate shelves. Gulf Coast Assoc. Geol. SOC.Trans., 1 7 : 49-60. Rose, P.R., 1972. Edwards Group, surface and subsurface, central Texas. Bur. Econ. Geol., Univ. Texas, Rep. Invest. No. 74, 1 9 8 pp. Rosenberg, P.E. and Holland, H.D., 1964. Calcite-dolomite-magnesite stability relations in solutions a t elevated temperatures. Science, 145: 700-701. Rosenberg, P.E., Burt, D.M. and Holland, H.D., 1967. Calcite-dolomite-magnesite stability relations in solutions: the effect of ionic strength. Geochim. Cosmochim. Acta, 31: 391-396. Truesdell, A.H. and Jones, B.F., 1973. WATEQ, a computer program for calculating chemical equilibria of natural waters. U.S. Geol. Surv., Comput. Contrib. PB 220 464, Progr. No. C 737, 7 3 pp. White, D.E., 1965. Saline waters of sedimentary rocks. Am. Assoc. Pet. Geol. Mem., 4: 342-3 66. Wolery, T.J., 1979. Calculation of chemical equilibrium between aqueous solution and minerals. the EQ3/6 software package. Univ. Calif. Lawrence Livermore Lab. UCRL52658,39 pp.

75

DISSOLUTION OF SALT ON THE EAST FLANK OF THE PERMIAN BASIN IN THE SOUTHWESTERN U.S.A.

KENNETH S. JOHNSON

Oklahoma Geological Survey, University of Oklahoma, Norman, OK 73019 (U.S.A.) (Accepted for publication February 26, 1981)

ABSTRACT Johnson, K.S., 1981. Dissolution of salt on the east flank of the Permian Basin in the southwestern U.S.A. In: W. Back and R. Lktolle (Guest-Editors), Symposium on Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54: 75-93. Hydrogeologic studies prove that natural dissolution of bedded salt occurs at shallow depths in many parts of the Permian Basin of the southwestern U.S.A. This is especially well-documented on the east side of the basin in study areas on the Cimarron River and Elm Fork in western Oklahoma, and on the Red River in the southeastern part of the Texas Panhandle. Four requirements for salt dissolution are: (1)a deposit of salt; (2) a supply of water unsaturated with respect to NaCl; (3) an outlet for removal of brine; and (4)energy to cause water to flow through the system. The supply of fresh groundwater in the region is recharged through permeable rocks, alluvium, terrace deposits, karstic features and fractures. Groundwater dissolves salt at depths of 10-250m, and the resulting brine moves laterally and upward under hydrostatic pressure through caverns, fractures in disrupted rock, and clastic or carbonate aquifers until it reaches the land surface, where it forms salt plains and salt springs. In many areas, salt dissolution produces a self-perpetuating cycle: dissolution causes cavern development, followed by collapse and subsidence of overlying rock; then the resulting disrupted rock has a greater vertical permeability that allows increased water percolation and additional salt diss o h tion.

INTRODUCTION

The Permian Basin is not a single structural basin but is a large region of the southwestern U.S.A. in which Permian salts and other evaporites were deposited with red beds and carbonates. In the eastern part of the basin, that is, in western Oklahoma and adjacent areas, several of the salt units at shallow depth are being dissolved by groundwater, and the resulting brine is being emitted at natural salt plains and salt springs. The current paper (1) summarizes hydrogeologic studies in three separate areas of brine emission, and (2) describes the processes and evidence of salt dissolution in the region. Much of this paper is based upon a long period of collaborative study with geologists and engineers of the U.S. Army Corps of Engineers, Tulsa District, who have been investigating methods that can be used t o control

76

natural salt-water degradation of the Arkansas and Red rivers in Oklahoma (U.S.A.C.E., 1976). Basic studies included detailed field investigation of physical stratigraphy and structure in the brine-emission areas, as well as evaluation of subsurface data (cores, water-well logs, and geophysical and lithologic logs of petroleum tests) t o map regionally the structure and the thickness, depth and distribution of salt units in the vicinity of the emission areas. Western Oklahoma and adjacent areas have a subhumid climate, with precipitation averaging 55 cm/yr.

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HYDROGEOLOGIC STUDIES

General geologic setting Rock units involved in salt-dissolution studies in western Oklahoma and nearby areas are mainly of early Guadalupian (Permian) age. These strata make up a thick sequence of red beds and evaporites deposited in and near a broad, shallow inland sea that extended north and northeast of the carbonate platform that bordered the Midland Basin (Fig. 1) (Mills, 1942; Clifton, 1944; Ham, 1960; Johnson, 1967). Evaporites, mainly salt (halite) and gypsum (or anhydrite), were precipitated from evaporating seawater as layers on the sea floor or grew as coalescing crystals and nodules in a host of mud just below the depositional surface. Thick red-bed shales, siltstones and sandstones were deposited around the perimeter of the evaporite basin, and some of these also extended as blanket deposits across the basin. Many thin red-bed elastic units are interbedded with the evaporites. Salt deposited in the inland sea during Permian time now occurs in several thick salt units that underlie a vast area extending across western Oklahoma COLORADO

I

KANSAS

I

w LOCATION MAP LOCATION MAP

Fig. 1. Paleogeography and principal facies in Permian Basin of the southwestern U.S.A., during deposition of evaporite facies of the Blaine Formation and Flowerpot Shale.

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comRAD0

Oklahoma City

OKLAHOMA

B IIASSIC

Salt beds (cross section) Salt plains and springs Streams degraded 200 KM

I

Fig. 2. Map and scnematic cross-section showing distribution of Permian salt and salt plains in western Oklahoma and adjacent areas. Three squares show outlines (from north) of Cimarron River, Elm Fork and Red River study areas described in this paper.

and adjacent states (Fig. 2). As partial evidence of dissolution of this salt, a number of natural salt plains and salt springs are present on the east side of the salt basin. Individual salt plains release brines with concentrations of 20-340 g NaC1/1, and they each contribute 100-3000 metric tons (t) of NaCl per day that degrades the quality of water flowing in the Arkansas and Red River systems. This paper draws heavily on results of detailed investigations at three separate areas in the region: the Cimarron River area in northwestern Oklahoma, the Elm Fork study area in southwestern Oklahoma, and the Red River study area in the southeastern part of the Texas Panhandle (Fig. 2). Principal stratigraphic units studied are the Flowerpot Shale and the overoverlying Blaine Formation and Dog Creek Shale (Fig. 3 ) . In outcrops the Flowerpot Shale consists mainly of reddish-brown shale with thin layers of greenish-gray shale, gypsum and dolomite; gypsum is more abundant to the

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RED RIVER AREA ELM FORK AREA SUDS1

CIMARRON RIVER AREA

PRINCIPAL LITHOLOGIES

.._1

Sandstone

!

' \

I

E l Shale

SubsurfacelOulcroD

\

\ Gypsum, shale, and dolomite

Salt and salty shale

Fig. 3 . Stratigraphic and lithologic relations between Permian outcrops and subsurface units in salt study areas of western Oklahoma and Texas Panhandle.

south in Texas. The formation is 60-100m thick in most outcrops (Fay, 1964; Johnson, 1967). Salt is present in the upper and middle parts of the Flowerpot Shale at shallow t o moderate depths just back from the outcrop in many areas of western Oklahoma and adjacent states. The total thickness of the Flowerpot salt (the sequence of salt-bearing strata in the Flowerpot Formation) generally ranges from 30 to 9 0 m (Jordan and Vosburg, 1963). The salt unit consists of reddish-brown shale interbedded with transparent to translucent layers, crystals and veins of salt. Commonly the salt is reddish in color, owing to shale impurities, and in many layers it is intimately intermixed with shale. Individual salt beds typically are 0.5-3.0m thick, and halite appears t o make up about one-half the entire Flowerpot salt unit. Outcrops of the Blaine Formation consist of gypsum beds, typically 2-10m thick, that are separated by red-brown shales 2--8m thick; each gypsum bed is underlain by a dolomite bed that commonly is 0.05-2.0m thick. The total thickness of the Blaine ranges from -30 m in northwestern Oklahoma to 60-75 m in southwestern Oklahoma and nearby parts of Texas (Fay, 1964; Johnson, 1967). In the deep subsurface of the Anadarko Basin and in the Palo Duro and Dalhart basins of the Texas Panhandle, the Blaine Formation contains several salt interbeds that are 2-6 m thick. The Dog Creek Shale consists of 1 5 - 6 0 m of red-bed shales in outcrops (Fay, 1964; Johnson, 1967), although gypsum and dolomite beds 0.1-5m thick are present in the lower half of the formation in southwestern Okla-

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homa and in Texas. The formation contains 50-100m of salt interbeds in the deep Anadarko Basin and in the Palo Duro Basin of the Texas Panhandle.

Cimarron River study area The Cimarron River, a major tributary of the Arkansas River, flows through an area in northwestern Oklahoma where the Flowerpot salt is at shallow depth. Natural brine forms when groundwater dissolves the salt, and the resulting brine then seeps from the bedrock into the alluvium, where the brine evaporates to form the two largest salt plains (Big Salt Plain and Little Salt Plain) on the Cimarron River (Fig. 2). These two salt plains emit an average of 5000 t of salt daily to the Cimarron River (Ward, 1961). Thus they represent the greatest source of salt-water contamination anywhere on the Arkansas River and its tributaries, and are evidence of active salt dissolution. Principal outcropping rock units around the two salt plains are flat-lying beds of the Flowerpot Shale and the overlying Blaine Formation (Fig. 4). The Flowerpot is more than 1 0 0 m thick in the area, and the gypsum beds and associated strata in the Blaine are -30m thick. Locally the upper Flowerpot and Blaine strata are mildly t o intensely folded, jointed, fractured and disrupted where they have collapsed and subsided, owing t o removal of underlying salt. Also, the Blaine gypsum beds commonly contain sinkholes, caves, and other karstic features resulting mainly from gypsum dissolution. Quaternary terrace deposits, consisting chiefly of sand and gravel in beds 2-15 m thick, locally mantle the Permian bedrock. The youngest sediments in the area are the sand, silt and clay that constitute alluvium along the flood plains of the Cimarron River, Buffalo Creek, and their tributaries; these young sediments typically are 3-20 m thick. Salt is present in the Flowerpot salt unit at depths of 10-60 m below the surface at Big Salt Plain (Fig. 4) and Little Salt Plain. The total thickness of the salt unit is -60-90m near Big Salt Plain and -75-100m near Little Salt Plain. The top of the Flowerpot salt commonly is 3 0 - 6 0 m below the top of the formation, but the upper surface is irregular as a result of dissolution of the salt by groundwater. The most pronounced dissolution and lowering of the upper surface of the salt is beneath and adjacent to the valleys of the Cimarron River and Buffalo Creek, and in some areas it closely parallels the alluvium-top-of-bedrock contact (Fig. 4). Coincident with this area of pronounced dissolution, the basal gypsum bed of the Blaine Formation dips gently at a rate of 4--8m/km (less than 0.5') toward both the Cimarron River and Buffalo Creek. Clearly, the drape of these outcropping rocks toward the stream channels results from intermittent subsidence of the strata into dissolution cavities developed in the Flowerpot salt. Boreholes drilled in and around the salt plains commonly encounter artesian flows. Solution cavities or zones of lost circulation 0.1-1.0 m thick are also common at the top of the salt unit, but these features generally are

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Outcrop of Blaine Forrnatlon and younger rocks

Big Salt Plain

Line of cross section shown below Outcrop of Flowerpot Shale ~~

~

A

* \ B

1 500

/--

Structure contour line, datum is base of Blaine Formation, interval is 6 meters

~

B &

Subsurface data from seismic profiles ~~

FLOWERPOT s

Fig. 4 . Structure-contour map and cross-section showing drape of base of Blaine Formation toward Cimarron River and Buffalo Creek, owing t o dissolution and removal of upper Flowerpot salt at Big Salt Plain in Cimarron River study area, northwest Oklahoma.

absent within the salt unit. Such features represent the zones at which dissolution is most actively taking place and the pathways whereby high-salinity brine is escaping from the salt beds. The thin, somewhat cavernous zone at the top of the salt is an artesian aquifer wherein the piezometric surface locally rises approximately to, or just above, the elevation of the surface of the salt plains. The overlying Flowerpot Shale in this area is not an aquiclude, but it allows some leakage of brine upward through fractures and thus acts as an aquitard.

TABLE I Chemical analyses of natural brines formed by dissolution of Permian salt beneath salt plains in western Oklahoma Big Salt Plain, Cimarron River, northwest Oklahoma (Johnson, 1970)

Specific gravity at 16OC (g/cm3 1 PH

Kiser Salt Plain, Elm Fork Red River, southwest Oklahoma (Johnson and Denison, 1973)

Chaney Salt Plain, Elm Fork Red River, southwest Oklahoma (Johnson and Denison, 1973)

1.208 6.9

1.203 6.9

1.209 6.8

HCO,

131,600 2,250 340 n.d. n.d. 205,000 48 10 4,040 36

n.d. 338 384 295 2 n.d. n.d. n.d. 759 220

n.d. 379 400 292 3 n.d. n.d. n.d. 656 120

Total

343,324

1,998

1,850

328,700 326,702

337,100 335,250

Concentration (mg/l): Na Ca Mg K Li C1 Br I

so4

Dissolved solids (Na -tC1) (by difference) n.d. = not determined.

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Natural brines flowing through salt-dissolution cavities at Big Salt Plain commonly are saturated with respect to salt. The concentration of NaCl in brine from one well is 337g/l, and (Na -4- C1) constitute 98%of all dissolved solids in the brine (Table I). The large quantity of brine locally present in the cavern system is evident from the fact that this brine is pumped at rates of 2000-4000 l/min. from wells 12-30 m deep for commercial production of salt by solar evaporation (Johnson, 1970). Not all brines in the area are salt-saturated; mixing of brine with near-surface fresh water dilutes the brine in some parts of the salt plains.

Elm Fork study area The Elm Fork of the Red River is one of the main tributaries of the Red River in southwestern Oklahoma (Fig. 2). In northern Harmon County the Elm Fork cuts through an area where the Flowerpot salt is at shallow depth and is being dissolved by groundwater. Adjacent t o the river are three small canyons, named the Chaney (Salton), Kiser and Robinson canyons, that contain salt plains at which brine is being emitted (Fig. 5). Each of the three salt plains contributes -200t of salt daily to the Red River system (Ward, B

M A

m,

Care

57 5

.......................... BLAINE.,..FORMAT!ON.. . . . . .

FLOWERPOT SHALE Elm ,Fork ~

,

2 125

(Salt dissolved)

I

t R 26 W

lWOM

I

I

T

6 N

Alluvium Gypsum and anhydrite

Salt and salty shale I

LOCATION MAP

L ~ W MI

a

Shale, mostly red b m m

Fig. 5 . Generalized cross-section through Elm Fork study area in southwest Oklahoma, showing dissolution limits of Flowerpot salt beneath principal river.

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1961), and they are the outlets for a moderate amount of brine currently produced by natural dissolution of salt in the subsurface. Outcropping rocks at these three salt plains are gently dipping red beds and gypsum units in the Flowerpot Shale and Blaine Formation (Fig. 5). The total thickness of the Flowerpot Shale is -100 m, of which the top 25 m is well exposed in cliffs of the canyons. The overlying Blaine Formation is -60 m thick and consists of nine major gypsum beds, 3-10 m thick, interbedded with thinner layers of shale and dolomite. Sinkholes, caves and other karstic features are abundant in the Blaine Formation, and locally the Flowerpot and Blaine strata are mildly jointed and fractured. Quaternary alluvium is estimated to be 5-20m thick in the bed of Elm Fork and only 1-6 m thick in the floors of the three canyons. The upper part of the Flowerpot contains as much as 60 m of interbedded salt and shale in the subsurface several kilometers to the south, but in the vicinity of the salt plains most of the salt is missing. The top of the Flowerpot salt is -10-15m below the surface of the salt plains, and the depth, thickness and distribution of salt here are erratic. The salt beds contain some natural dissolution cavities that generally are less than 1m high. The cavities contain brine under artesian conditions with a piezometric surface that locally is as much as several meters above the salt plains in the canyon floors. The Flowerpot Shale in the base of the canyons is fractured and acts as an aquitard that permits slow, upward flow of brine along joints and fracture planes. Natural brines produced from shallow wells in the Kiser and Chaney salt plains are saturated with respect to salt (Table I). (Na C1) comprise 327 and 335g/l of brine samples, and they represent more than 99% of the dissolved solids in the brine. Two companies have produced solar-evaporated salt from these brines. The brines are pumped at rates of 300--400l/min. from wells drilled 9-12m deep into natural cavities in salt layers (Johnson and Denison, 1973).

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Red River study area Several of the main tributaries t o the Red River (Prairie Dog Town Fork of Red River, Pease River, and their tributaries) cross outcrops of the Blaine Formation and associated strata in the southeast corner of the Texas Panhandle (Fig. 2). This area, in the eastern part of the Palo Duro Basin, embraces much of Childress, Cottle, Hall and Motley counties, Texas (Fig. 6). Salt is present at depths greater than 150-250m below the surface (Fig. 7) in the Flowerpot, Blaine and Dog Creek formations (these three units are sometimes called the “Blaine of Texas” in outcrops, and they are considered part of the San Andres Formation farther west in the deep subsurface of the Palo Duro Basin). Natural dissolution of these salts in the subsurface here and farther to the west produces low- t o high-salinity brines that emerge in five separate areas’ characterized as salt plains, seeps, or springs. These five

a4

= Sandstone

n Shale

Blaine Formation Salt and salty shale

Flowerpot Salt

Gypsum, anhydrite, dolomite

I

MATADOR

I MOTLEYCO

pRcH I 20KM

COrrLECO’--1

Lines of cross section shown in following figure

J

Fig. 6 . Map and columnar section showing distribution of salt units in subsurface of Red River study area of Texas Panhandle. See Fig. 7 for cross-sections A-B and C-B.

areas are along Pease River, Jonah Creek, Salt Creek, Little Red River, and a long stretch of Prairie Dog Town Fork near and downstream from Estelline Spring. They, along with other less-clearly defined emission areas nearby, contribute -2000t of salt daily t o the Red River system and constitute a major region of modern-day natural dissolution of salt. Principal outcropping units around the various salt plains are moderatelyto highly-disrupted gypsum and dolomite beds of the Blaine Formation and the gypsiferous lower member of the Dog Creek §hale. The sequence consists of interbedded gypsum beds (2-10 m thick), dolomite beds (0.1-5 m thick) and shale beds (0.2-10 m thick). The Blaine is typically 60 m thick, and the gypsiferous part of the Dog Creek is -30 m thick. Although the regional dip is generally 2--10m/km t o the southwest into the basin, the outcropping units in many places are chaotically folded, fractured, and disrupted where they have collapsed owing t o removal of underlying salt and (to a lesser extent) gypsum. Inasmuch as gypsum and dolomite make up -75% of the total thickness of these outcropping strata, sinkholes, caves and other karstic features are abundant throughout the region. Bedrock units are mantled in some areas by 3-15 m of Quaternary terrace deposits consisting mainly of sand and gravel. Along the principal rivers and streams of the region, alluvial sands, silts and clays commonly are 5-20m thick. The Flowerpot salt, the most widespread of the shallow salt units in the four-county area (Figs. 6 and 7), typically is 60-75m thick in the west where the salt is more than 1 5 0 m below the land surface. However, salt is

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B

A 1 700

Mx)

500

400

300 1. Amerada, core 56 2 . Sinclair, Shannon 1 3 Texas Gulf, House 1

200

4. Honolulu Noel 1

5.Constantine, Wilton 1 6 Maguire,SmlthI,&CI

B

30KM 20 10 A - 1 I

C

r

_I

Prairiepn,B rnnrrln Do

1

7J

'FLOW'

200

HOLLIS ANTICLINE

MERKEL BED

Alluvium

Sandstone and shale

7 Humble. Matador C 1 8 Amerada, core 21 9 Amerada. core 17 10 Amerada. core 37

Gypsum, anhydnte, dolomite and shale

J

11 Amerada,core43 12 Pure, Gourd L & C1 13 Corps of Engrs , J o n a h Cr SWD 14 Corps of Engrs , core 1

Salt and salty shale

Shale

Fig. 7. Generalized cross-sections in Red River study area of Texas Panhandle, showing dissolution of salt units in vicinity of principal rivers and remnant of salt in syncline northeast of Plaska Dome (A-El, well 5). See Fig. 6 for lines of cross-sections.

generally absent from the Flowerpot t o the northeast and east where equivalent strata are only 60-150m deep. Particularly striking is the absence of salt along an east-west line beneath and north of Prairie Dog Town Fork. However, an outlying undissolved mass of salt, 5 8 m thick, was penetrated at a depth of 1 0 7 m in the Constantine well drilled in a shallow structural depression between the Hollis Anticline and Plaska Dome (Figs. 6 and 7).

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Salt occurs at one stratigraphic level in the middle of the Blaine Formation in the four-county area. The salt unit is 6-12m thick and consists of interbedded salt and shale layers, 0.3-3m thick. Salt is restricted to the western part of the area (Figs. 6 and 7), where it is more than 150-210m below the present land surface. The present northern and eastern limits of this salt are irregular in plan view, and the absence of salt from this saltshale interval 9-12 m thick within a distance of only 2-3 km at many places along this irregular limit can best be explained as a dissolution phenomenon when compared t o depositional patterns of this and similar salt units in other parts of the Permian Basin. The youngest salts in the area occur in the Dog Creek Formation, where these strata are as much as 150-250 m deep in the southwest (Figs. 6 and 7). The total thickness of the salt-bearing Dog Creek sequence is typically 50100 m, and salt commonly appears to make up 50% of the sequence. Salts in the Dog Creek thin abruptly and are completely dissolved within a short distance in the southwestern part of the Red River study area (Fig. 7, crosssection C-B). As much as 1 1 5 m of salt-bearing strata in well 7 is represented by 6 5 m of non-salty equivalent strata in well 8, only 11km to the northeast. Such abrupt thinning is not consistent with the normal pattern of nearly uniform thickness for these salt-bearing strata beneath thousands of square kilometers farther west, in deeper parts of the Palo Duro Basin. Salt deposits and dissolution phenomena of the Palo Duro Basin are currently being investigated in detail by the Texas Bureau of Economic Geology (Dutton et al., 1979; Gustavson et al., 1980a, b, 1981), and new data from those studies will greatly enhance our knowledge of geohydrologic processes in the region.

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DISSOLUTION O F SALT

Salt (halite) is highly soluble, more soluble than any other rock in the Permian sequence of western Oklahoma and nearby areas. Groundwater in contact with salt will dissolve some of the salt, providing the water is not already saturated with NaC1. For extensive dissolution t o occur, it is necessary for the brine thus formed t o be removed from the salt deposit; otherwise the brine becomes saturated, and the process of dissolution stops. Four basic requirements are necessary for salt dissolution to occur here, or in other evaporite basins for that matter (Johnson et al., 1977): (1)A deposit of salt against which or through which water can flow. (2) A supply of water unsaturated with NaC1. (3) An outlet whereby the resulting brine can escape. (4)Energy (such as a hydrostatic head or density gradient) to cause the flow of water through the system. When all four of these requirements are met, salt dissolution and brine transport can be quite rapid, in terms of geologic time.

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Salt-dissolution process Hydrogeologic studies show that natural dissolution of bedded rock salt occurs at shallow depths at many places in western Oklahoma and adjacent areas. Fresh and saline groundwater moves laterally through aquifers, such as sandstone or cavernous gypsum, dolomite, or salt, and also water moves vertically through fractures, sinkholes and collapse features. Fresh groundwater is recharged generally to the west of the salt plains, in upland areas where unconsolidated sands (Ogallala Formation of Tertiary age) or sandstone, gypsum, dolomite, alluvium, or terrace deposits are at the surface (Fig. 8). This water migrates downward and laterally (eastward) to salt beds, which are 10-250m below the surface, and dissolves the salt to form brine. The resulting brine is then forced laterally and upward by hydrostatic pressure through aquifers or through fractures in aquitards until it is discharged at the surface. There are four principal ways whereby fresh groundwater is recharged in the region (Fig. 8): (1)Water seeps into the ground through permeable rocks and soils, such as in areas where sandstone is at the surface. (2) Water enters the bedrock through highly permeable alluvium and terrace deposits along and near the major streams and rivers.

k

h?L Fresh water

--

-6

f

Brine

Dissolution zones and cavities

Disrupted rock

Fig. 8. Schematic block-diagram showing circulation of fresh water and brine in areas of salt dissolution in western Oklahoma. No scale for diagram, but length may be 1-15 km, and height 30-300 m.

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(3) Water enters the ground through sinkholes, caverns and other karstic features, in areas where gypsum, dolomite, or limestone is at the surface. (4)Water enters the ground through joints and fractures present in the rocks, particularly where underlying salt beds are partly dissolved and the rock is more fractured owing to collapse. After the water has dissolved some of the salt and has become brine, there are six principal ways whereby the brine moves underground and is eventually discharged (Fig. 8): (1)Brine moves through dissolution cavities in the salt or other soluble rocks. (2) Brine moves vertically and (or) laterally through joints and fractures, particularly where the rock is disrupted over dissolution cavities. (3) Brine moves laterally through aquifers consisting of sandstone, siltstone, or other permeable rocks. (4)Brine may be discharged at a point-source as a saline spring. ( 5) Brine may be discharged along the course of a stream bed and become part of the surface flow. (6) Brine may enter the base of an alluvial deposit where it can be forced upward under hydrostatic pressure and then drawn upward by capillary action as the brine is evaporated. A thin crust of salt then is precipitated on the land surface as water is evaporated from the brine. In all cases cited above, the energy needed t o cause flow of the water is the hydrostatic head created in the recharge areas, with brine moving laterally and upward toward the piezometric surface. When dissolution occurs, the resulting collapse, subsidence and fracturing of overlying rock causes a greater vertical permeability along joints and openings. Therefore, salt dissolution can produce a self-perpetuating cycle: dissolution causes cavern development and then land subsidence, with the resulting disrupted rock having a greater vertical permeability that allows increased water percolation and additional salt dissolution. Evidence of salt dissolution A number of criteria have been recognized on the east flank of the Permian Basin indicating that salt is being, or has been, dissolved by natural processes. (1)Salt plains, salt springs and salt seeps in areas that are underlain by, or are near, subsurface salt deposits strongly indicate the likelihood that salt is being dissolved. It is possible that the brine may result from emission of connate waters or other formation waters whose salinity is not derived by dissolution of salt, however, these brines generally can be distinguished chemically. The salt plains and similar features normally are in stream valleys and other topographic lows that have intersected the piezometric surface of a brine-bearing aquifer. (2) The Na/C1 ratio of brines formed by dissolution of salt in western

89

Oklahoma is remarkably close to 0.64, regardless of whether the water is a low-salinity or a saturated brine (Leonard and Ward, 1962). This is because salt (and very little else) is being dissolved from the nearby salt deposits, and the combining ratio of Na and C1 in pure crystals of halite is 0.65. Oil-field brines consistently have Na/C1 ratios of 0.55 or less, and the ratio decreases well below 0.50 as the salinity increases. (3) The Ca, Mg and SO4 concentrations and the Ca/S04, Mg/S04 and (Ca Mg)/S04 ratios of salt-derived brines in this Permian sequence also are generally distinguishable from oil-field brines. The differences in concentrations of these ions between salt-derived brines vs. oil-field brines typically are: Ca, 300-3000 mg/l vs. 5000-15,000 mg/l; Mg, 300-500 mg/l vs. 1000-3000 mg/l; and SO4, 500-6000 mg/l vs. 100-1000 mg/l [oil-fieldbrine data generalized from T.W.D.B. (1972)l. Of greater value, however, are the ratios between these constituents, because a brine diluted with varying amounts of fresh water can vary greatly in the concentration of its constituents but the ratios of these constituents t o each other will vary only slightly. The differences in ratios between salt-derived brines vs. oil-field brines typically are: Ca/S04, 0.20-0.70 vs. 10-100; Mg/S04, 0.10-0.70 vs. 2-10; (Ca Mg)/S04, 0.4-1.4 vs. 10-100. (4) Cavities, caverns, or other openings in salt beds are evidence of past or present dissolution. The cavities commonly are filled with low-salinity t o saturated brines that may or may not be under artesian pressure. The cavities may also be partly or totally filled with clay or other sediment deposited from water that had passed through the openings. Cavities may also contain brecciated rock from overlying formations that collapsed into the openings. (5) The presence of low-salinity t o saturated brines in aquifers that do not contain salt interbeds may still indicate dissolution of salt, particularly where the brine has an appropriate ratio of NaCl to the other constituents [see (2) and (3), above]. The brine would be a saline groundwater plume that extends away from the parent-salt being dissolved. By backtracking along the paths of brine migration-that is, by determining the groundwater flow paths and the direction of increased salinity - it should be possible to locate the area where salt dissolution is occurring. (6) Chaotic structures, collapse features, and other evidence of disturbed bedding can result from collapse of rocks into small or large salt-dissolution caverns formed at shallow to moderate depths. Such features are especially common in the Red River study area, and to a lesser extent on the Cimarron River, but they also are reported over old dissolution areas in the Anadarko Basin (Johnson, 1967). The nature of this disruption is not typical of true geologic structures; instead, the strikes and dips of strata are chaotic and change sharply within short distances. Extreme examples of chaotic rock are the small-diameter breccia chimneys or pipes that extend vertically up from dissolution cavities through several hundred meters of overlying Permian strata, as described by Eck and Redfield (1963), Kirkland and Evans (1976), and Anderson and Kirkland (1980). An example of dramatic, present-day

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+

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development of similar collapse structures is seen at the Wink Sink in Kermit County, West Texas, where a surface sink 1 1 0 m in diameter developed in 48 hr. over a salt-dissolution cavern which is presumably -400 m below the surface (Baumgardner et al., 1980). (7) A pronounced decrease in thickness of shallow salt beds beneath or in the vicinity of major streams or rivers is likely t o result from dissolution. The salt is at shallower depths beneath the valleys than it is beneath nearby divides and uplands, and thus it is more generally accessible t o circulating fresh groundwater. Furthermore, the abundant quantity of water flowing in the streams can exchange freely with groundwater in alluvium and nearby bedrock, thus increasing the availability of fresh groundwater that can accelerate salt dissolution and flush or remove brine before it is fully saturated. Examples of accelerated dissolution adjacent t o rivers are seen in the Cimarron River study area, where the top of the Flowerpot salt is lowered and overlying rocks are draped toward the valleys (Fig. 4),and also in the Red River area, where the limits of several of the salt units are nearly coincident with the course of Prairie Dog Town Fork (Figs. 6 and 7). (8) Outlying patches or masses of salt distant from the main body of salt, where they cannot be explained reasonably as depositional features, probably result from dissolution of salt beneath the intervening area. Such patches are possible mainly where the outlying salt is at a greater depth (beneath a structural low or a topographic high) than equivalent strata nearby, or is protected in some other way from dissolution that has removed salt from the surrounding area. Such an example is present in the Red River study area, where a mass of Flowerpot salt is preserved in a syncline 15 km north of the dissolution front (Figs. 6 and 7). In some areas, of course, an outlying mass of salt may not be a dissolution remnant but may result from deposition in a small basin or pan, separated from the main site of salt deposition. (9) Inliers or windows where salt is locally thin or absent from an area that is surrounded by salt usually indicate dissolution of salt, provided that the feature cannot be explained reasonably as a depositional feature. Examples of such inliers were shown by Jordan and Vosburg (1963) in the Flowerpot salt in Beaver County, Oklahoma, and evidence of such dissolution is found in the breccia chimneys and pipes [see (6) above] . (10) Breccia beds at a stratigraphic position known t o be, or believed t o have once been, occupied by salt are important evidences of dissolution. Anderson et al. (1978) differentiated two types of breccia beds in the Castile Formation of the Delaware Basin: (a) dissolution breccias, consisting of chaotic and subangular fragments of thin non-salt interbeds that remain after the surrounding salt is dissolved; and (b) collapse breccias, consisting of angular fragments of rock units above the salt bed that drop chaotically into dissolution cavities and overlie the dissolution breccias. Such features are identified only in cores or outcrops of the breccia beds. (11)Irregular lateral limits and abrupt thinning of salt beds, in a manner

-

91

that is not consistent with depositional facies changes normally recognized in the unit under study, are indications of salt dissolution. This is common in the Red River study area, where the widespread and uniformly thick salts of the Flowerpot, Blaine and Dog Creek formations thin abruptly and are missing within a few kilometers at a number of places. Also, the northern and eastern limits of these salt beds typically are irregular in plan view. (12) Salt casts and molds, in the form of cubic and hopper-shaped voids, impressions, and fillings in rocks, are evidence of the former presence of salt in the rock. These features are fairly common in each of the study areas and are best seen in the dolomite beds that are capable of retaining the delicate impressions of the salt crystals. The quantity of salt removed from such casts and molds is small in comparison t o that which can be removed from bedded rock salt.

CONCLUSIONS

Hydrogeologic studies in several areas on the east side of the Permian Basin prove that natural dissolution of bedded rock salt is occurring at shallow depths. Fresh groundwater moves downward and laterally under hydrostatic pressure and dissolves salt at depths of 10-250 m t o form brine. The resulting brine then moves upward and laterally under hydrostatic pressure until it reaches the surface. Both the fresh water and the brine can move laterally through aquifers and also can move vertically across aquifers or aquitards through fractures, sinkholes and collapse features. The process of salt dissolution produces cavities, normally at the updip limit or at the top of the salt unit, into which overlying rocks can settle or collapse chaotically. Disrupted rock helps to make salt dissolution a somewhat self-perpetuating process, inasmuch as cavern development followed by collapse and fracturing of the rock will cause a greater vertical permeability, and this allows further access of fresh water to the salt. Among the main criteria for recognizing that salt dissolution is, or has been, going on are salt plains, brines with characteristic ionic ratios, and disrupted or chaotic rock above areas that are, or have been, underlain by salt. Other criteria include cavities or other openings in salt beds and the presence of breccia beds at stratigraphic positions once occupied by salt. Further evidence of dissolution is the abrupt thinning or absence of salt in an area surrounded by thick salt or the presence of an outlying mass of salt where equivalent strata in the surrounding area are devoid of salt. Salt-dissolution studies, such as the ones herein reported, are necessary for determining the sources of natural salt-water contamination in the rivers and streams of the region. Such studies also are needed for unravelling the geology above dissolution areas, and for interpretation of seismic records made across dissolution boundaries. Furthermore, such hydrogeologic data also are of major importance in evaluating the potential for safe storage of

92

radioactive wastes in salt deposits far from dissolution zones (Dutton et al., 1979; Gustavson et al., 1980a, b, 1981). ACKNOWLEDGMENTS

Many of the data presented in this report were collected and developed jointly with S. Thomas Gay and other geologists and engineers of the US. Army Corps of Engineers, Tulsa District, and thanks are extended to Mr. Gay, Phillip E. LaMoreaux, and Thomas C. Gustavson for review of this manuscript. Appreciation also is expressed t o Robert F. Walters, Charles L. Jones, Walter E. Dean, and Roger Y. Anderson, with whom the author has discussed a number of problems on salt dissolution in the Permian dasin. Illustrations were drafted by Marion Clark, cartographic technician with the Oklahoma Geological Survey. Publication of the report is authorized by the Director of the Oklahoma Geological Survey. REFERENCES Anderson, R.Y. and Kirkland, D.W., 1980. Dissolution of salt deposits by brine density flow. Geology, 8: 66-69. Anderson, R.Y., Kietzke, K.K. and Rhodes, D.J., 1978. Development of dissolution breccias, northern Delaware Basin, New Mexico and Texas. In: G.S. Austin (Compiler), Geology and Mineral Deposits of Ochoan Rocks in Delaware Basin and Adjacent Areas. N.M. Bur. Mines Miner. Resour., Circ. No. 159, pp. 47-52. Baumgardner, R.W., Gustavson, T.C. and Hoadley, A.D., 1980. Salt blamed for new sink in west Texas. Geotimes, 25(9): 16-17. Clifton, R.L., 1944. Paleoecology and environments inferred for some marginal middle Permian marine strata. Am. Assoc. Pet. Geol. Bull., 28: 1012-1031. Dutton, S.P., Finley, R.J., Galloway, W.E., Gustavson, T.C., Handford, C.R. and Presley, M.W., 1979. Geology and geohydrology of the Palo Duro Basin, Texas Panhandle. Texas Bur. Econ. Geol., Geol. Circ. 79-1, 99 pp. Eck, W. and Redfield, R.C., 1963. Geology of Sanford dam, Borger, Texas. Panhandle Geol. SOC.,Field Trip Guideb., pp. 54-57. Fay, R.O., 1964. The Blaine and related formations of northwestern Oklahoma and southern Kansas. Okla. Geol. Surv., Bull. 98, 238 pp. Gustavson, T.C., Finley, R.J. and McGillis, K.A., 1980a. Regional dissolution of Permian salt in the Anadarko, Dalhart and Palo Duro Basins of the Texas Panhandle. Texas Bur. Econ. Geol., Rep. Invest. 106, 39 pp. Gustavson, T.C., Presley, M.W., Handford, C.R., Finley, R.J., Dutton, S.P., Baumgardner, Jr., R.W., McGillis, K.A. and Simpkins, W.W., 1980b. Geology and geohydrology of the Palo Duro Basin, Texas Panhandle. Texas Bur. Econ. Geol., Geol. Circ. 80-7, 99 PP. Gustavson, T.C., Simpkins, W.W., Alhades, A. and Hoadley, A., 1981. Karstification of the Rolling Plains of the Texas Panhandle. J. Earth Surf. Process. (in press). Ham, W.E., 1960. Middle Permian evaporites in southwestern Oklahoma. 21st Int. Geol. Congr., Copenhagen, Part XII, Reg. Paleogeogr., pp. 138-151. Johnson, K.S., 1967. Stratigraphy of the Permian Blaine Formation and associated strata in southwestern Oklahoma. Ph.D. Dissertation, Illinois University, Urbana, Ill. (unpublished).

93 Johnson, K.S., 1970. Salt produced by solar evaporation on Big Salt Plain, Woods County, Oklahoma. Okla. Geol. Notes, 30 : 47-54. Johnson, K.S. and Denison, R.E., 1973. Igneous geology of the Wichita Mountains and economic geology of Permian rocks in southwest Oklahoma. Okla. Geol. Surv., Spec. Publ. 73-2-Guideb. Geol. SOC.Am., Field Trip No. 6 (1973 Annu. Meet.), 33 pp. Johnson, K.S., Brokaw, A.L., Gilbert, J.F., Saberian, A., Snow, R.H. and Walters, R.F., 1977. Summary report on salt dissolution review meeting, March 29-30, 1977. Union Carbide Corp., Nuclear Div., Off. Waste Isolat., Rep. Y/OWI/TM-31, 1 0 pp. Jordan, L. and Vosburg, D.L., 1963. Permian salt and associated evaporites in the Anadarko Basin of the western Oklahoma-Texas Panhandle region. Okla. Geol. Surv. Bull. 102, 76 pp. Kirkland, D.W. and Evans, R., 1976. Origin of limestone buttes, Gypsum Plain, Culberson County, Texas. Am. Assoc. Pet. Geol. Bull., 60: 2005-2018. Leonard, A.R. and Ward, P.E., 1962. Use of Na/Cl ratios to distinguish oil-field from saltspring brines in western Oklahoma. In : Geological Survey Research 1962. U S .Geol. Surv., Prof. Pap. No. 450-B, pp. 126-127. Mills, J.M., 1942. Rhythm of Permian seas - a paleogeographic study. Am. Assoc. Pet. Geol. Bull., 26: 217-255. T.W.D.B. (Texas Water Development Board), 1972. A survey of the subsurface saline water of Texas, Vol. 2. Chemical analysis of saline water. Texas Water Dev. Board, Rep. 1 5 7 , 3 7 8 pp. U.S.A.C.E. ( U S . Army Corps of Engineers), 1976. Arkansas-Red River Basin chloride control. U.S. Army Corps. Eng., Tulsa Dist., Tulsa, Okla., Des. Mem. No. 25, Gen. Des. Phase 1,Plan Formul., Vols. I and 11, 7 Appends. Ward, P.E., 1961. Salt springs in Oklahoma. Okla. Geol. Notes, 21: 82-85.

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PATTERNS OF GROUNDWATER SALINITY CHANGES IN A DEEP CONTINENTAL-OCEANICTRANSECT OFF THE SOUTHEASTERN ATLANTIC COAST OF THE U.S.A.

F.T. MANHEIM and C.K. PAULL* U.S. Geological Survey, Woods Hole, M A 02543 (U.S.A.)

(Accepted for publication February 26, 1981)

ABSTRACT Manheim, F.T. and Paull, C.K., 1981. Patterns of groundwater salinity changes in a deep continental-oceanic transect off the southeastern Atlantic coast of the U.S.A. In: W. Back and R. Lktolle (Guest-Editors), Symposium on Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54: 95-105. Investigations of formation-fluid salinities in a transect from western Georgia t o the edge of the Blake Plateau off the coast of Georgia show surprisingly similar hydrochemical features offshore and onshore. A fresh-brackish wedge of groundwater ( l o 0 g/kg) occur in Lower Cretaceous (?) strata. These strata have a pronounced evaporitic (anhydritic) character in the offshore segment. Strong salinity gradients in interstitial waters signify buried evaporite deposits at drill sites beneath the Blake Plateau.

-

INTRODUCTION

Until relatively recently, hydrologists in the U.S.A. focused their main attention upon shallow groundwaters containing potable and near-potable waters. Petroleum exploration firms generally limited their interests in fluidbearing strata to prospective petroleum-producting (reservoir) beds. Consequently, data on regional hydrochemical patterns, migration paths, permeability and sources of chemical constituents in sedimentary basins as a whole were rare. Increased interest in use of subsurface strata as waste-disposal sites, in problems involving seawater encroachment, and in extension of the concept of “freshwater resources” to water containing as much as lOg/kg total dissolved solids (TDS) (Kohout, 1981) has increased interest in the composition and distribution of deeper saline fluids in sedimentary strata. Manheim and Horn (1968) summarized subsurface hydrochemical data from a shoreline transect extending from Long Island t o the Florida Keys and provided a map of inferred salinity distributions from surface to igneous-metamorphic basement. The southeastern Atlantic region from *Present address: Scripps Institution of Oceanography, La Jolla, CA 92093, U.S.A.

96

South Carolina to northern Florida was shown t o be the site of some of the most complex hydrochemical features along the U.S.A. Atlantic seaboard. The deepest freshwater horizons extended more than 1000 m below sea level in the region of South Carolina north of Parris Island. In contrast, in southernmost Georgia, hypersaline brines (i.e. brines containing at least 50-100 gTDS/kg were found at depths below 700m. The distribution of such hypersaline brines, which approached saturation in NaCl beneath much of Florida, was linked to the distribution of evaporitic strata (Manhein and Horn, 1968). Evaporitic facies mainly characterized by dolomite-anhydrite occur in Paleocene rocks in Florida and southern Georgia (Chen, 1965), and thicker strata including some salt are found in Lower Cretaceous and Jurassic evaporites in Florida on land. Salt also occurs in diapirs offshore north of the Blake Plateau (Grow et al., 1977) and has been sampled during oil drilling in the Baltimore Canyon Trough (Oil and Gas JournaZ, 1978). Recently, Brown et al. (1979) evaluated in detail the deep-well wastestorage potential of Mesozoic aquifers in Georgia and South Carolina. They estimated formation-fluid salinity ranges by using electrical-log analysis and water analyses. We have taken advantage of these data, have further analyzed some of the electrical logs, and have incorporated hydrochemical data from available offshore drill holes and other onshore drill holes t o prepare crosssections of subsurface "salinity" (TDS content). Table I contains a list of sites utilized; Fig. 1 shows their locations in Georgia and Florida and offshore regions. The authors' purpose is not to determine fine-scale chemical 84" I

EXPLANATION

34"

-0

82"

c7

I

80"

78"

I

(

ONSHORE DRILL, SITES

34"

32"

30"

28"

Fig. 1 . Location of drill holes utilized in study. Abbreviations as shown in Table I. Bathymetric contour interval, 200 m.

TABLE I List of boreholes Utilized in study Borehole No.

Borehole name

Source of information*

LOW-1 LOW-2 LOW-3 CAL-1 DO-1 MI-1 COL-1 EC-2 EC-5 SCR-1 SAV JAX STM. COL.IS. SUN GE-1 T 1,2,3,4,5, 6 6002 6004

Hunt Petroleum, J.T. Stalvey No. 1 , Lowndes County, Ga. Hunt Petroleum, Langsdale No. 1 , Lowndes County, Ga. Hunt Petroleum, E.N. Murray No. 1, Lowndes County, Ga. Sowega Mineral, J.W. West No. 1 , Calhoun County, Ga. J.R. Sealy, Reynolds No. 1, Dougherty County, Ga. Stanolind, J.H. Pullen No. 1, Mitchell County, Ga. R.T. Adams, D.G. Arrington No. 1, Colquitt County, Ga. Hunt Petroleum, Superior Pine Co. 2, Echols County, Ga. Humble Oil, Bennett and Langdale No. 1, Echols County, Ga. Boenwell Drilling Co., McGain-Pryor No. 1, Screven County, Ga. Savannah Port Authority, Chatham County, Ga. Jacksonville test well, U.S.G.S. U.S. Geological Survey), Duval County, Fla. St. Mary R. Oil/Hilliard Turpentine Co., Nassau County, Fla. Colonel's Island test well, U.S.G.S., Glynn County, Ga. Sun Oil Co., Powell, Volusia County, Ga. COST (Continental Offshore Stratigraphic (Test) well GE-1 Tenneco wildcat well JOIDES (Joint Oceanographic Institutions Deep Earth Sampling Project) sites AMCOR (Atlantic Marine Coring Project) site 6002 AMCOR site 6004

(1*2 (1*2 (1*2 (1*2 (1*2 (1*2 (1*2

1 (2)

(3)

(3) (5) (61, (7) (3),( 8 ) (9) (9)

*1

Sources o f information: (1)= Brown et al. (1979); ( 2 ) = G.W. Leve (1961), cited in Manheim and Horn (1968); (3) = Manheim and Horn (1968); (4)= H. Gill, U.S.G.S. (pers. commun., 1979); (5) = Scholle (1979); (6) = L. Poppe (pers. commun., 1979); (7) = R. Johnston, U.S.G.S., Reston, Va. (pers. commun., 1979); (8) = Manheim (1967); (9) = Hathaway et al. (1979). *2 Re-interpreted in part from original logs. CD -4

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variations, but t o discern broad regional trends that might shed light on fluid history, fluid migration, and fundamental geochemical and hydrochemical processes. METHODS

All available sources of information on the composition of formation fluids in the study are (Fig. 1) were utilized. The most accurate data are potentially those obtained from analysis of: (1) drill-stem tests or other (e.g., reverse flush) fluid tests on permeable strata in land wells (no oil-producing wells exist in the study area); and (2) pore water extracted (by “squeezing”) from cores of unconsolidated and partly consolidated rocks at depths as great as 300m beneath the sea floor (Manheim, 1967; Manheim and Horn, 1968; Hathaway et al., 1979). The validity of drill-stem-test data is normally governed by the care used in sampling fluids t o minimize the influence of drilling mud. A minimum criterion for water.-test samples is that chloride analyses on sequential fluid samples (either flowing water or successive pipe stands) reach a constant value asymptotically. The squeezing-andanalysis methodology was used extensively in Deep Sea Drilling Project studies (Manheim and Sayles, 1974; Manheim and Gieskes, 1981). Wireline samples from which external drill-fluid-contaminated zones had been removed were squeezed through filter paper in a stainless steel press, and recovered fluids were analyzed by microchemical techniques. Data are available from all JOIDES (Joint Oceanographic Institutions Deep Earth Sampling) and AMCOR (Atlantic Margin Coring Project) sites (see Fig. 1 and Table I) (Hathaway et al., 1979; and references cited herein). Much less accurate but indispensable is the technique of estimating fluid resistivity and, through it, formation-fluid “salinity ” by quantitative electrical-log interpretation. For the older drill holes on the continent, the only available method is simple estimation from the spontaneous potential (SP) log, as has been done by Manheim and Horn (1968) and Brown et al. (1979). The technique is described in standard logging references (Schlum berger Well Surveying Corp., 1978, and documents cited therein) :

SSP = -- K log ( Rmf/Rwe) (1) where K is a constant dependent on temperature; R,, is the resistivity of mud filtrate; R,, is the apparent resistivity of formation fluid; and SSP‘is the static spontaneous potential (SP). The SSP is derived from the departure in millivolt scale units from shale baseline of the SP curve; it is corrected where possible for the effects of thin beds, mud resistivity, and fluid invasion of the formation. “Salinity” or the salinity of NaCl solutions having resistivities corresponding t o observed resistivities can then be calculated from R w e . We have used the Arps-Hamilton log analysis slide rule or calculator program (e.g., Schoonover and Fertl, 1979).

99

The SP data are particularly useful in deposits containing fresh and brackish strata, as the errors (to 50%) are still smaller than the salinity variations; the salinity values vary by four orders of magnitude. In our study areas, the simpler SP methods give rise t o serious errors for the deeper strata containing saline water and especially for those containing significant proportions of clay colloids. Further, these methods are not applicable in carbonate strata. For the GE-1 well (Scholle, 1979) that reached 4004m depth below sea floor, -140 km seaward of Jacksonville, Florida, we used alternative methods utilizing deep induction (resistivity) and porosity logs. The basis for such calculations are as follows:

F = R,/Rw

and

F =

where F is the “formation factor”; R, is the true formation resistivity; R, is the true formation-fluid resistivity; 4 is the porosity; and a and rn are constants for given types of strata. An empirical formula (Schlumberger Well Surveying Corp., 1978) for formation-fluid resistivity, given 100%water saturation, can be expressed as: Rw = @ 2 [0*8(Rt- v s h / R s h ) ( l - vsh)l-l where V,h are volume percent of shale within the formation; and Rsh is the “shale” resistivity. The V, and R, figures can be obtained from induction-, neutron- and formation-density logs by cross-plot techniques described in Schlum berger Well Surveying Corp. (1978) and references cited therein. We were most fortunate t o have both porosity and interstitial-salinity values as well as resistivity values for AMCOR sites 6002 and 6004. These permitted direct calibration of eq. 2 and 3 for Tertiary strata penetrated by nearby GE-1. For deeper strata, the estimations are based upon cross-plot data and eq. 4, and inferences of saline residues in proximity t o anhydrites, analogous to conditions beneath the mainland (see Manheim and Horn, 1968). The “salinity” data (NaC1-equivalents) are subject t o variation in ionic ratios. These variations are not discussed here but the error they introduce is less than the uncertainty attributable t o log analysis.

STRATIGRAPHIC NOTES

The geologic cross-sections upon which the salinity data are superimposed (Figs. 2 and 3) are based on available literature (Toulmin, 1955, Herrick and Vorhis, 1963; Chen, 1965; Applin and Applin, 1967; Maher, 1971; Cramer, 1974; Schlee, 1977; Dillon et al., 1979; Hathaway et al., 1979; Scholle, 1979). We divide the stratigraphy of southern Georgia into the units: Neogene-Holocenej Eocene and Oligocene, Paleocene, Upper Cretaceous, Lower Cretaceous (?), and igneous and metamorphis basement.

100

The Neogene-Holocene unit is composed of predominantly clastic sediments, and the carbonate content increases to the southeast. The Eocene and Oligocene unit is dominantly carbonate except for some of the lower Eocene in southwestern Georgia, where clastic deposits dominate. The Upper Cretaceous and Paleocene units are generalized as marl, but carbonate (chalk) and evaporitic facies (anhydrite4olomite) increase southeastward in the Paleocene. The lithologies of the Lower Cretaceous (?) are different in the Southeast and Southwest Georgia embayments; the basins are probably connected only by a thin basal sand. The unit thickens t o more than 800 m of unfossiliferous, immature sandstone in southwest Georgia, where its age is questionable (Gohn et al., 1978) but has historically been labelled as Early Cretaceous. The Southeast Georgia Embayment contains a thick and variable sequence of Lower Cretaceous limestone, anhydrite, sandstone and shale over metamorphic basement (Scholle, 1979).

Paleoenvironment Except for anhydrite in the Paleocene and traces of evaporite units in the lower Eocene, all sediments of the Upper Cretaceous and Cenozoic were deposited in normal marine conditions. The lack of distinguishing fossils in the Lower Cretaceous (?) of the Southwest Georgia Embayment renders paleoenvironment identification difficult. In the offshore Southeast Georgia Embayment, the presence of evaporitic strata including anhydrite probably indicates hypersaline deposition. DISTRIBUTION OF FORMATION SALINITY

The main salinity features are delineated in Figs. 2 and 3 by lines of equal “salinity” (TDS content), or isosalines, at concentrations of 1 , 5,25, 50, 100 and 200g/kg. These values may be converted t o parts per million by multiplying by 1000. The isosalines are superimposed on transects A-A’, and B-B‘, whose locations are shown in Fig. 1. In interpreting the diagrams, we note that the isosalines may be subject to error, particularly where salinity distributions are complex, as between Echols and Mitchell counties, Georgia, (between MI-I and EC-2, Figs. 1and 3), or in the deeper parts of the offshore basin. There may be inliers of different salinity and complex microstructure that cannot be depicted at the scale used here. Moreover, the salinities derived from the SP log in the deeper Southeast Georgia Embayment may be in error because of clay colloids and may understate true formation-fluid salinity. Major relationships are evident from Figs. 2 and 3 : (1)Strata on land have been infiltrated by meteoric (fresh) water (identified as water having a salinity less than 1g/kg) to depths between 350 m in Colquitt County, Georgia, (COL-I), t o more than 500 m near Jacksonville, Florida, (JAX, Fig. 3). Farther northward near the South Carolina--Georgia

101

LEVEL

KILOMETERS

I

t4

Fig. 2. Salinity-stratigraphy transect, A-A' (see Fig. 1for location). Contours are in total dissolved solids, g/kg. Abbreviations for wells are as in Fig. 1.Stratigraphic intervals are: NEO-HOL = Neogene-Holocene; E O - O L I G O = Oligocene and Eocene; PALEO. = Paleocene; UK = Upper Cretaceous; L K ? = units historically referred to as being Lower Cretaceous; DEV. MET. = Devonian metamorphic rocks; PALEOZ. MET. = Paleozoic metasedimentary rocks. Partial limestone symbols in Upper Cretaceous and Paleocene strata represent marls and some evaporitic (anhydritic) sediments. The Neogene-Holocene section is variable but is predominantly of clastic lithology.

border (SAV, Fig. 2), fresh waters extend deeper, to more than 900-m depth. Freshwaters mixed with brackish waters have salinities less than the salinity of seawater (35 g/kg) and extend t o a depth of 1.2 km in Lowndes County, Georgia (LOW-1, -2, Fig. 3 ) . (2) A fresh-brackish-water (significantly less saline than seawater) wedge extends under the Atlantic Ocean as far as 120km from shore, down to depths greater than 600 m. This phenomenon was reported in papers cited earlier and is discussed in considerable detail by Kohout (1981), and Paul1 and Dillon (1981). The evidence suggests that boundaries are much smoother and more regionally continuous than on land. These boundary characteristics are consistent with predictions based on the mobility of water influenced by artesian circulation (land) relative t o the mobility of water in largely diffusive fluxes (beneath the sea). ( 3 ) Only beneath the fresh-brackish-water lens underlying the shelf do we find substantial ,thicknesses of formation fluids having salinities within the approximate range of seawater salinities. On the continent, salinities

102

0

-m I 111 W

t

w

z 0 -J2 Y

I

l-

a

x 3

4 Fig. 3 Salinitystratigraphy transect, B-B‘. For explanation, see Fig. 2. Depth of isosalines in area seaward of the shelf is estimated beneath drillholes partly by salinity gradients in JOIDES holes 4 , 5 and 6.

generally pass relatively abruptly from brackish t o hypersaline levels (>50 g/kg). (4)The presence of hypersaline brines in stratigraphic levels as young as Paleocene appears t o be limited to those areas where the evaporitic rocks are present in the earliest Tertiary; i.e. Echols County, Georgia, and eastward. (5) More concentrated hypersaline brines (>lo0 g/kg) are limited t o Lower Cretaceous (?) strata on land. A lens of such fluids appears t o extend from Calhoun County to Lowndes County, Georgia, in transect B-B’ (Fig. 3, CAL-1 to LOW-5); below the lens is less saline water. However, at these depths, the SP-values yield only semiquantitative “salinity ” values, because the “greasy” (high-clay-content) nature of the varicolored micaceous sands reported in this zone affects the well logs. (6) Offshore, in the GE-1 well, very high salinities are identified in Lower Cretaceous strata, consistent with the presence of evaporitic rocks (anhydrite). Toward the base of the section, low permeabilities render interpretation of salinities more difficult and less reliable, even with the aid of the porosity methods. (7) Previous results (Manheim and Horn, 1968) have shown that along the Atlantic margin, salinity commonly decreases just above the basement. In this study, we have detected indication of this phenomenon at sites in eastern Georgia and possibly in COST GE-I. On the other hand, indications

103

in some sites (LOW-I and St. Mary Hilliard, transect B-B’) are that salinity levels of 50-100 g/kg continue into crystalline or metamorphosed sediments.

DISCUSSION

The existence of fresh water beneath the Atlantic Ocean, to the edge of the U.S.A. continental shelf, has been well documented in the JOIDES and AMCOR drill holes (Manheim, 1967; Hathaway et al., 1979; Kohout, 1981; Paul1 and Dillon, 1981). New electrical-logging data from the COST GE-1 well confirm the existence of brackish water in the upper 900 m. Moreover, a special agreement by Tenneco Oil Co. permitted the U S . Geological Survey to run a drill-stem test at -350-m depth, in a wildcat well (Fig. 1, T ) -85 km seaward of Jacksonville, Florida. A drill-stem test confirmed presence of brackish water having less than half the salinity of seawater (R. Johnston, pers. commun., 1979). The distribtuion of lithologic and stratigraphic units and formation-fluid salinities shows clearly that few if any “paleosalinities” remain in deeper subsurface porous strata. All surficial strata have had the original seawater solutions permeating the interconnected pores of marine sediments flushed out by fresher waters of meteoric origin. We suggest that in the depth range where freshwater influence diminishes sharply, the dominant salting influence is frequently not seawater but brine of hypersaline origin. These brines once originated from seawater, but they have been modified by processes involving secondary interactions with solid phases (Braitsch, 1971; Carpenter, 1978) during deposition and burial. Evaporite brines may permeate not only contemporaneous sediments, but also underlying strata to depths of several kilometers or more (e.g., Manheim and Schug, 1978). Another source of salt is inclusions and microlayers of rock salt that are associated with original anhydritic rocks and are later dissolved, thereby contributing to total formation-fluid salinity. Not infrequently, anhydritic strata are characterized by brine concentrations approaching saturation with respect t o NaC1, even though no appreciable salt bodies occur in the strata. In previous papers, the senior author and his coworkers (Manheim and Sayles, 1974; Manheim and Hall, 1976) pointed out that significant increases in salinity of interstitial waters as depth increases in oceanic strata nearly always point to presence of evaporitic strata at depth. As the present data indicate, “evaporitic strata” need not mean massive halite. Anhydritic rocks also incorporate a sufficient reservoir of brine salt t o permit upward diffusion t o influence fluids of overlying strata during geologic time. The new information can be applied to the interstitial salinity gradients shown in Figs. 2 and 3 to infer the extension of Lower Cretaceous (?) evaporitic facies to JOIDES holes 4 , 5 and 6 (Figs. 1and 3). Pre-existing hypersaline concentrations (if any) in earlier, deeper strata would have merged with the saline-

104

water concentrations contributed by the latest evaporitic sequence and would probably not be discernible in the study area. A further inference may be drawn from the saline gradients. If inorganic ions (Na, C1, etc.) can move upward through the strata in response to concentration gradients, then light hydrocarbons dissolved or otherwise entrained in pore fluids may likewise be able to migrate through the strata. More detailed delineation of saline gradients may help map zones of diffusive or other permeability and “calibrate” surficial hydrocarbon anomalies.

REFERENCES Applin, P.L. and Applin, E.R., 1967. The Gulf Series in the subsurface in northern Florida and southern Georgia. U.S. Geol. Surv., Prof. Pap. 524-G, 24 pp. Braitsch, O., 1971. Salt deposits: their origin and composition. Springer, New York, N.Y., 297 pp. Brown, P.M., Brown, D.C., Reid, M.S. and Lloyd, O.B., 1979. Evaluation of the geologic and hydrologic factors related to the deep-well, waste-storage potential of Mesozoic aquifers in the southern part of the Atlantic Coastal Plain, South Carolina and Georgia. U.S. Geol. Surv., Prof. Pap. 1088,137 pp. Carpenter, A.B., 1978. Origin and chemical evolution of brines in sedimentary basins. Okla. Geol. Surv., Circ. 79: 60-77. Chen, C.S., 1965. The regional lithostratigraphic analysis of Paleocene and Eocene rocks of Florida. Fla. Geol. Surv., Bull. 45, 1 0 5 pp. Cramer, H.R., 1974. Isopach and lithofacies analysis of the Cretaceous and Cenozoic rocks of the Coastal Plain of Georgia. In: L.P. Stafford (Editor), Petroleum Geology of the Georgia Coastal Plain, Symposium. Ga. Geol. Surv., Bull., 87 : 21-43. Dillon, W.P., Paull, C.K., Buffler, R.T. and Fail, J.P., 1979. Structure and development of the Southeast Georgia Embayment and northern Blake Plateau, preliminary analysis. In: J.S. Watkins, L. Montadert and P.W. Dickerson (Editors), Geological and Geophysical Investigations of Continental Margins. Am. Assoc. Pet. Geol. Mem., 29 : 27-41. Gohn, G.S., Christopher, R.A., Smith, C.C. and Owens, J.P., 1978. Preliminary stratigraphic cross sections of Atlantic Coastal Plain sediments of the southeastern United States, Part A. Cretaceous sediments along the South Carolina coastal margin. U.S. Geol. Surv., Misc. Field Stud. Map, MF-1015-A. Grow, J.A., Dillon, W.P. and Sheridan, R.F., 1977. Diapirs along the continental slope off Cape Hatteras. SOC.Explor. Geol. 47th Annu. Int. Meet., Calgary, Alta., Alberta Program, p. 51 (abstract). Hathaway, J.C., Poag, C.W., Valentine, P.C., Miller, R.E., Schultz, D.M., Manheim, F.T, Kohout, F.A., Bothner, M.H. and Sangrey, D.A., 1979. U.S. Geological Survey core drilling on the U.S. Atlantic Shelf. Science, 206(4418): 515-527. Herrick, S.M. and Vorhis, R.C., 1963. Subsurface geology of the Georgia Coastal Plain. Ga. Geol. Surv., Info. Circ. 25, 78 pp. Kohout, F.A., 1981. Reflict fresh ground water of the continental shelf: an unevaluated buffer in present-day salt water encroachment. In: D.D. Arden, B.F. Beck and E.N. Morrow, Proc. 2nd Symp. on Geology of the Southeast Coastal Plain. Ga. Geol. Surv., Info. Circ. 53 (in press). Lefond, S.J., 1969. Handbook of World Salt Resources. Plenum, New York, N.Y., 384 pp. (see especially pp. 82-84). Maher, J.C., 1971. Geological framework and petroleum potential of the Atlantic Coastal Plain and Continental Shelf. U.S. Geol. Surv., Prof. Pap. 659, 9 8 pp., 17 plates.

105 Manheim, F.T., 1967. Evidence for submarine discharge of water on the Atlantic continental slope of the southern United States, and suggestions for further search. N.Y. Acad. Sci. Trans., Sect. 11, 29: 839-853. Manheim, F.T. and Bischoff, J.L., 1969. Geochemistry of pore waters from Shell Oil Co. drill holes on the continental slope of the northern Gulf of Mexico. Chem. Geol., 4: 63-82. Manheim, F.T. and Gieskes, J.M., 1981. Interstitial water methods. In: Initial Reports of the Deep Sea Drilling Project. U.S. Government Printing Office, Washington, D.C. (in press). Manheim, F.T. and Hall, R.E., 1976. Deep evaporitic strata off New York and New Jersey: evidence from interstitial water chemistry of drill cores. U.S. Geol. Surv. J. Res., 4:697-702. Manheim, F.T. and Horn, M.K., 1968. Composition of deeper subsurface waters along the Atlantic continental margin. Southeast. Geol., 9 : 215-236. Manheim, F.T. and Sayles, F.L., 1974. Composition and origin of interstitial waters of marine sediments, based on deep sea drill cores. In: E.P. Goldberg (Editor), The Sea, Vol. 5. Wiley, New York, N.Y., Ch. 16, pp. 527-567. Manheim, F.T. and Schug, D.M., 1978. Interstitial waters of Black Sea cores. In: California University, Scripps Institution of Oceanography (LaJolla, Calif.), Initial Reports of the Deep Sea Drilling Project, Volume XLII, Natl. Sci. Found., Washington, D.C., pp. 637-651. Oil and Gas Journal, 1978. H O & M plugs third Baltimore Canyon dry hole. Oil Gas J., 76(38): 72. Paull, C.K. and Dillon, W.P., 1981. The stratigraphy of the Florida-Hatteras shelf and slope, and its relationship t o the offshore extension of the principal artesian aquifer. In: D.D. Arden, B.F. Beck and E.N. Morrow, Proc. 2nd Symp. on Geology of the Southeast Coastal Plain. Ga. Geol. Surv., Info. Circ. 53 (in press). Schlee, J.S., 1977. Stratigraphy and Tertiary development of the continental margin east of Forida. U.S. Geol. Surv., Prof. Pap. 581-F, 25 pp. Schlumberger Well Surveying Corp., 1978. Long interpretation charts. Schlumberger Well Surveying Corp., Houston, Texas, 8 2 pp. Scholle, P.A., 1979. Geological studies of the COST GE-I Well, United States South Atlantic Outer Continental Shelf area. U S . Geol. Surv., Circ. 800, 114 pp. Schoonover, L.G. and Fertl, W.H., 1979. How t o find temperature, R , and salinity with hand calculators. Oil Gas J.. 77: 109-111. Toulmin, L.D., 1955. Cenozoic geology of Georgia. Am. Assoc. Pet. Geol. Bull., 39: 207-2 35.

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107

CHARACTER OF BRINES FROM THE BELLE ISLE AND WEEKS ISLAND SALT MINES, LOUISIANA, U.S.A.

MADHURENDU B. KUMAR and JOSEPH D. MARTINEZ

Institute f o r Environmental Studies, Louisiana State University, Baton Rouge, L A 70803 (U.S.A.) (Accepted for publication February 26, 1981).

ABSTRACT Kumar, M.B. and Martinez, J.D., 1981. Character of brines from the Belle Isle and Weeks Island salt mines, Louisiana, U.S.A. In: W. Back and R. Lhtolle (Guest-Editors), Symposium on Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54: 107-140. This paper is based on the chemical analyses of 110 samples of brines collected from the Belle Isle and Weeks Island salt mines of Louisiana, U.S.A. These brines are manifested by the active leaks, puddles of water and stalagmitesstalactites in the mines. Most of the waters are a sodium-calcium-chloride brine with minor concentrations of additional ionic species. The ionic compositions of these brines are compared with average seawater, seawater concentrated by evaporation and formation waters (oil-field brines). The ionic distributions of the mine brines resemble those of formation waters except that the mine brines are relatively high in potassium, strontium, boron and bromide. The higher concentrations of potassium and bromide in these brines relative to formation waters indicate a similarity t o a residual brine and suggest some genetic relationship. However, the high concentration of Ca, exceeding the concentration of K and, in some instances, even that of Na in the mine brine is highly indicative of formation water. It is concluded that the mine brines may represent an admixture of (Ca-rich) formation water and (K-Br-rich) residual brine or a formation water of unusual character with high concentrations of Ca, K and bromide. Isotopic data strongly support the formation water hypothesis.

INTRODUCTION

The determination of effectiveness of a salt stock to block inflow of water into and outflow of water from a mined opening in a dome is a critical aspect of the nuclear-waste isolation issue. Impermeability and self-sealingproperties of salt as a rock type make it attractive as an isolation medium. Nevertheless, variable wet conditions of some salt mines in Louisiana may have raised questions concerning the potential hydrologic integrity of a nuclear-waste respository in a salt dome. The evaluation of the degree of hydrologic isolation that a salt dome can afford is one of the tasks undertaken by the Institute for Environmental Studies, Louisiana State Unversity, Baton Rouge, U.S.A. The present paper focuses upon conditions in the Belle Isle and Weeks Island salt mines, two of the several salt mines of the Gulf of

108

Mexico Coast in which mine leaks occur. The purpose of this study is to characterize and interpret the origins of the mine leaks.

THE SALT MINES AND SUBSURFACE LEAKS

Belle Isle and Weeks Island are two salt domes of the well-known fiveisland trend of south Louisiana (Fig. 1).The geological framework of the two domes is indicated in Figs. 2 and 3; in essence, a Tertiary sequence of sandstone and shale surrounds the upper parts of the salt stocks. In the vicinity of these domes there are active petroleum fields. The outlines of mine workings for the Belle Isle and Weeks Island salt mines are shown in Figs. 4 and 5. Mining in the Belle Isle dome has been only on a 355m (1165ft.) level. The Weeks Island mine has two old levels, at 152.5 and 244 (500 and SOOft.) below the surface and a new 91.5m (300ft.) level (Markel mine) being developed. Evidence of wet conditions in the two mines are active leaks, puddles and pools of water on the floor, drippings from the ceiling and stalactites and stalagmites (Figs. 6-10). These are some of the characteristic features of anomalous zones which also exhibit blowouts (gas pressure pockets), oil and gas seeps, unusual variations of structure, texture, clastic inclusions and

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Fig. 1. Location map of Belle Isle and Weeks Island as part of the five-island diapir trend marked by five salt domes with surface expression. (From Kupfer, 1974.)

Fig. 2. A crosssection through the Belle Isle dome (1mi. = 1609 m). Inset shows structure contours on “F-Horizon” at subsea elevations of 286.5-329.2 m (9400-10,800 ft.). (From O’Neill 111, 1973.)

110

Fig. 3. The crosssection (a) is of the north and west flanks of the Weeks Island dome. The map (b) shows structure contours on the “S-Sand” (1 ft. = 0.3048m; 1 mi. = 1609 m). (After Atwater and Forman, 1959.)

111

Fig. 4. Location map of the sampled subsurface leaks from the Belle Isle salt mine, also showing the approximate outline of the mine superimposed upon the structure contours on the salt top of the Belle Isle dome. (Not shown are several past leaks reportedly in the middle of the mine.)

113

Fig. 6. Puddle of water between Rooms 9 and 10 of 8th main entry east, Belle Isle.

Fig. 7. Stalactite (parted off the ceiling) in Room 3 between 0-1st Belle Isle.

main entry west,

114

Fig. 8 . Sampling of active water leak (through a pipe driven into the main face) at the end of 1st main entry south, Belle Isle.

Fig. 9. Active water along the undercut, forming pools at location A-2 (north) on the new level of the Weeks Island salt mine.

115

Fig. 10. A borehole full of water at location A-1 (north) on the new level of the Weeks Island salt mine.

dark-colored salt. These features were sampled and photographed in the course of mapping the mine leaks during 1977-1979. The distribution of subsurface leaks studied in the Belle Isle and the Weeks Island mines is shown in Figs. 4 and 11. Several active leaks in the new Weeks mine (Markel mine) were mapped in 1979. The mine leaks have been described in detail by B. Hoda, M.B. Kumar, J.D. Martinez and R.L. Thoms in Martinez et al. (1977, 1978, 1979). A great many leaks reportedly were of short duration. Such leaks generally appeared as ceiling drips soon after a working was opened up and lasted for a few days to six months. In the process of roofbolting, the mine operators encountered cracks that drained 75.7-378.5 1 (20---lo0gal.) of water. Drill holes encountered water. Several of the water leaks are associated with gas and oil seeps. At three active leaks in the Belle Isle mine, rates of water flow up t o 0.76 l/hr. (0.2 gal./hr.) were observed In the Markel Incline of the Weeks Island mine the initial flows at up to 189.25 l/hr. (50 gal./hr.) was reduced by grouting to -18.9 l/hr. (- 5 gal./hr.) in 1978. NATURE OF DATA AND GENERAL CHARACTER OF MINE BRINES

110 samples of the water leaks in the Belle Isle and Weeks Island salt mines were collected in 256 ml polyethylene bottles. They were chemically analyzed

t

117

at Petroleum Laboratories, Inc., Lafayette, according t o A.P.H.A. (1975). The resulting analytical data were reported in Martinez e t al. (1978, 1979). The average total amount of dissolved salts in the mine waters is -290 g/hr. or 350,000 mg/l, a t room temperature (maximum 380 g/kg or 485,570 mg/l). Their specific gravities range between 1.195 and 1.310 g/cm3, with their pH-values mostly between 1.5 and 5.5. The ionic distributions in the brine samples are shown in the Figs. 1 2 and 13. The major cations (exceeding 1000mg/l) in the brines are Na, Ca, K and Mg. Sr is also a major cation in most of the brines. The chloride ions account for over 90% of the anions in the mine brines. Bromide is the second most abundant ion. Minor anions

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TECHNIQUES OF DATA ANALYSIS

The basic approach to the understanding of the character and origin of the mine brines adopted here is to compare them with the standard watersource types, namely, average (modern) seawater, evaporating concentrated

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seawater, oil-field brines or formation waters of Louisiana. To this end, the following types of graphical plots of the analytical data were prepared: (1) Ternary plots of percentage reaction values of major ions (Na, K, Ca and Mg) in brine samples: (reaction value) = (ion concentration, [mg/l] ) x

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(2) Log plots of concentrations of important ions in brine samples, using analytical data normalized to 19,000 mg C1/1 (modern seawater) (3) Log-log plots of concentrations of “marker” constituents (chloride and bromide separately) vs. those of other constituents of brine samples. (4)Plots of ionic distribution (with average and maximum concentrations of ions) of the mine brines and formation waters of various ages.

121

Fig. 14. Ternary plots of relative concentrations (in reaction-value percentages) of the major cations in salt mine brines, average seawater, and formation waters of Louisiana (data mostly from Collins, 1975).

IMPLICATIONS OF ANALYTICAL PLOTS

The four types of plots prepared and indicated above are shown in Figs. 14--20. Their implications relative to the chemical character of the mine brines are discussed in the following sections.

Ternary plots of major cations Relative to the major cations (Na, Ca, K and Mg) the domain of mine brines is depicted in Fig. 14 which is located close to the Na-Ca side of the triangles, indicating that they are essentially Na-Ca-Cl brines. The cationic character of seawater at the various stages of evaporation is also plotted in Fig. 14. These plots fall along the Na-Mg and Na-K sides of the triangles, away from the domain of mine brines. This shows that they are very different from seawater at any stage of evaporation. The domain of Louisiana formation waters as worked out by M.B. Kumar in Martinez et al. (1979) is also indicated in Fig. 14. The domain of the formation waters falls within that of mine brines, which indeed extends towards the 100% Ca end-point of the triangle. In other words, a large number of mine leaks have Ca contents much in excess of those of typical formation waters. The ternary diagrams also demonstrate that the mine brines are more enriched in K than are formation waters.

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123 Fig. 15. Chemical composition of brine samples (solid dots) from: (a) Belle Isle mine; (b) new Weeks Island mine (Markel Mine); ( c ) upper level Weeks Island mine; and (d) lower level Weeks Island mine. All data normalized to 19,000 mg/l chloride. Bars and open circles denote average seawater and formation waters of Louisiana, respectively.

124

Plots of normalized ionic concentrations Fig. 15 shows the plots of concentration of important ions of the mine brines normalized to 19,000 mg C1/1 (for average modern seawater). For comparison purposes, relevant information on average seawater and Louisiana formation waters has been included in those plots. From them the following observations are in order: (a) In comparison to average seawater the mine brines are highly enriched in K, Ca and Sr, and moderately enriched in Mg, Rb, Li and Ba, and are relatively low in Na, sulfate and bicarbonate. The mine brines are considerably altered, assuming that they originally were seawater. (b) The ionic distribution of the mine brines bears a strong similarity to that of formation water relative to Na, Ca, Mg and sulfate. The concentrations of Sr, bromide and boron in the mine brines fall in the upper ranges of values for the formation waters. The mine brines, however, distinctly differ from the formation waters in that they (mine brines) are relatively high in K, Rb, Li and nitrate, and are generally low in iodide and sometimes in bicarbonate. Chloride and bromide plots In order to trace the origin of formation waters, the waters have been compared with normal evaporite curves of Collins (1970, 1975) and Carpenter (1978). Collins used log-log plots of chloride concentration vs. concentration of other ions. Carpenter adopted a different approach by plotting log of bromide concentration against log of concentration of other constituents. Similar approaches have been utilized for a study of the mine brines and are discussed in the following sections. Chloride plots. The concentration of chloride is plotted against the concentration of Na, K, Ca, Mg, Sr and B, respectively, of the mine brines in Fig. 16. For comparison purposes, the normal evaporite curve and the domain of some Louisiana oil-field waters from Tertiary, Cretaceous and Jurassic rocks have been shown on each of the plots. Fig. 16a presents log-log plots of chloride concentration us. sodium concentration of the mine brines. The plots, which scatter parallel to and above the normal evaporite curve, indicate that the waters are affiliated with an evaporation process. The upward shift of the plots is suggestive of some dissolution (addition) of halite into the waters. Fig. 16b shows log-log plots of chloride us. potassium. They show that the Louisiana oil-field waters are depleted in K, relative t o sea brine subjected to evaporation. This is indicative of the loss of K of the associated sediments during diagenesis. The plots of the mine brines are suggestive of the late phase of seawater evaporation and a slight enrichment of K, except for a K depletion in the Markel Incline of the Weeks Island mine.

125

Fig. 16c has log-log plots of chloride us. calcium. They indicate that all of the mine waters, except those from the Markel Incline, are enriched in Ca, relative to the evaporating seawater. This is in conformity with the trend of the Louisiana oil-field waters, which show Ca concentration increasing with increased salinity. The Markel Incline waters appear somewhat depleted in Ca, relative to the evaporating seawater.

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5000

I 1 1 1 1

10,000

I

20,000

BROMIDE, m g / l

Fig. 16. Comparison of: (a) sodium; ( b ) potassium; ( c ) calcium; (d) magnesium; ( e ) strontium; ( f ) boron; and (g) bromide concentrations, respectively, of mine brine samples with an evaporating seawater and Louisiana formation waters (dotted outline) from Collins (1975).

Bromide p b t s . Zherebtsova and Volkova (1966) demonstrated that during the evaporation of seawater essentially all of the K, Rb and Li and bromide remain in solution until K-rich salts begin to precipitate and that most of the Li and bromide survive in solution during K-rich salt deposition. Of these elements, only bromide does not precipitate significantly in diagenetic reactions (Carpenter, 1978). The bromide concentrations have already been employed successfully in the investigation of the origin of some oil-field

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300

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EVAPORATING SEAWATER

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200

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Fig. 1 7 a - c . For caption see p. 131.

I

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5000

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1

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20,000

130

brines. Thus, bromide is significant as a “marker” constituent. The log-log plots of bromide vs. Na, K, Ca, Mg, chloride and sulfate, respectively, in the mine brines are compared with Carpenter’s standard curves in Figs. 1 7 and 18 as discussed below. Fig. 17a presents the plots of bromide us. sodium, which show that the mine brines plot near the standard trend line for seawater. This strongly suggests that these brines are genetically related t o concentrated seawater. Fig. 17b has plots of bromide us. potassium. The mine brines follow the seawater trend line, most of them clustering primarily around the upward end of the trend line. This is suggestive of brine near the stage of K-rich salt precipitation. Fig. 17c shows plots of bromide us. calcium. The mine brines plot far to the right top of the seawater trend line which strongly indicates a high enrichment in Ca concentration, relative to seawater subjected to evaporation.

I00

L t

+

New Level W e e k s 0 Upper Level Weeks Lower Level Weeks 0 Belle Isle

20

131

000

0

BROMIDE, ( r n g / l )

100,000

I

0 0 0

500 0

0

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0

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N e w Level Weeks Upper Level Week6 Lower Level Weeks Belle Isle

0 0

0 0

0

0

20

I

I

I

I

I I I I I

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I

I l l l l l l

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Fig. 17. Relations of: (a) sodium; (b) potassium; (c) calcium; (d) magnesium; (e) chloride; and (f) sulfate vs. bromide, respectively, in the mine brine samples, compared to the trend (solid line) of an evaporating seawater (after Carpenter, 1978).

132

Fig. 17d presents plots of bromide us. magnesium. The mine brines plot far below the seawater trend line, indicating an appreciable depletion in Mg concentration, relative to seawater evaporation. Fig. 17e has plots of bromide us. chloride. The mine brines plot near the upper end of the curve, suggesting the late stage of halite precipitation. Fig. 17f shows plots of bromide us. sulfate. The mine brines plot far below the standard trend line, which is suggestive of an excessive deficiency of sulfate. According to Carpenter (1978), the parameter MC1, [which equals milli-equivalents per liter of (Ca Mg Sr - SO4 - HCO, ) ] appears to be of some value in determining the chemical history of C1-rich brines. A plot of log MClz us. log Br in evaporating seawater (to the point of carnallite precipitation) is a straight line with a 1:1slope, symbolized as follows:

+

log MC1, = log Br

+

+ 0.011

Most of the mine brines plot just below the straight line (Fig. 18),suggesting

5000 -

t 0 0

0

New L e v e l W e e k s Upper L e v e l W e e k s

Lower Level Weeks Belle Isle

BROMIDE, ( mg / I )

Fig. 18. MClz (meq. M2+/1)-bromide relationships in the mine brine samples, compared to the calculated concentration trend (solid line) of an evaporating seawater (after Carpenter, 1978).

that they are slightly depleted in MC1, or enriched in bromide, with respect to brines resulting from the evaporation of seawater. The Markel Incline waters plot to the left bottom (outside) of the graph, suggesting an entirely different origin of the waters.

Summary. To summarize these observations, all the mine brines (except those from the Markel Incline of the Weeks Island mine) appear in some respects

133

genetically related to the evaporation of seawater affected by dissolution of halite from the salt deposit. These brines have concentrations similar to the late stage of halite precipitation prior to the precipitation of K-rich salt. Relative to evaporating seawater, the brines are considerably enriched in Ca, Sr and Br, slightly enriched in K and are depleted in Mg, B and SO4. In view of these characteristics, the mine brines resemble formation waters and residual evaporite brines in varying degrees.

Plots of ionic distributions o f mine brines and formation waters of various ages It is instructive to compare the distribution of all the important ions in the mine brines and formation waters from Tertiary, Cretaceous and Jurassic systems, since the mine brines could be related to the rocks of these various ages. To this end, graphical plots of the average and maximum concentrations of Na, K, Ca, Mg, Sr, Ba, B, C1, Br, I, SO4 and HCO, in those waters were constructed, and the average concentration graphs are presented in Fig. 19. These graphs show that the composition variation patterns of the mine brines, except those from the Markel Incline, are generally similar to those of the formation waters, although the mine brines have much higher contents of K, Sr, B and bromide, and are conspicuously low in iodide, relative t o the formation waters. The dominance of Ca over Mg is obvious. The Markel Incline brines contrast with the formation waters in having high concentrations of Na and C1 and low concentrations of other ions. As determined by Morton Salt Company, U.S.A., using the dye-injection method, the water in the Markel Incline is from ponds on the surface. Thus the Markel Incline brine clearly has a meteoric source. This water, which was originally fresh, became highly saline through dissolution of salt as it penetrated into the mine incline.

DISCUSSION AND CONCLUSIONS

As compared to the Markel Incline brine which is clearly meteoric in origin, all other mine brines are non-meteoric in character. The higher concentration of bromide and K in these brines relative to formation waters indicates a similarity to a concentrated relict or residual brine, and suggests some genetic relationship. On the other hand, the high concentration of Ca (exceeding the concentration of K and, in some instances, even that of Na) in the mine brines, is suggestive of an affinity with formation waters (Fig. 14). Characteristically, a relict (residual) brine lacks Ca, contains Mg salts, bromides and iodides, and represents the very late phase of evaporation of seawater, as illustrated in Fig. 20. This distinguishes the mine leaks from relict brines, even though the brines bear some similarities t o both formation waters and concentrated relict brines. Thus the mine brines bear some

134

1 Zt-

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I

/

1

1 1

t

c

-

@

AVERAGE

-

CONCENTRATION, I m g

136

,":,":::: ; 100.000 -

50,000 -

-

10.000

,

I

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rn

-E z 0

2

1,000 -

F

8

500-

z

s W 4 0

; t

100 -

50

-

20 10

0

-

1

Fig. 19. Comparison of the average chemical character of brines from: (a) Belle Isle mine (B.I.M.); (b) Markel Incline of Weeks Island mine (M.Z.W.Z.); (c) the new Weeks Island mine (N.W.Z.) or Markel mine; ( d ) the upper level Weeks Island mine (U.L.W.Z.); and (e) the lower level Weeks Island mine (L.L. W.Z.), respectively, with the formation waters (from Collins, 1975) of the Tertiary, Cretaceous and Jurassic systems.

similarities to both formation waters and concentrated relict brines. The high Ca/Na ratio, however, would appear to rule out an origin from relict brines. On the other hand, the relatively high Br/K concentration could represent a contamination by concentrated relict brines or bitterns through a mixing phenomenon or an unusual type of formation water. Some components rich in Mg and K of the brines may have been acquired through dissolution of late-stage or secondary evaporite minerals such as bischofite, carnallite, etc. However, the crux of the problem is related to the factors or processes responsible for the enrichment of Ca in the mine brines. These processes, which are briefly reviewed in the following paragraphs, could have occurred at the time and site of salt deposition, or in an environment outside the salt deposit.

137 STAGES OF EVAPORATION OF SEAWATER I

II

rn

m

P

1,000,000

"

CHLORIDE

-

z 0 c

i

c z W

z 0

8

\I

I

II

m

E

I

P

Fig. 20. Variations in concentration of important constituents in the average seawater subjected to evaporation in five stages (data plotted from Collins, 1975).

At the time of salt deposition, the initial seawater may have been richer in Ca than the modern ocean (Spiro and Vouk, 1961; Kramer, 1963). However, this possibility has been ruled out by White (1965) whose analysis proves that such a simple relation to time is untenable. Thus, this possibility merits no further consideration. From the pattern of salt precipitation in average (modern) seawater (Fig. 20), it is apparent that a significant concentration of Ca exists in the residual brine at the stage of halite precipitation. This residual brine may have been trapped in the initial salt mass which was mobilized during subsequent diapirism. This is a distinct possibility assuming that the Ca-rich brine was emplaced as part of the diapiric system. This will be elaborated on later in this discussion (p. 138).

138

Some noteworthy processes which may have operated outside the site of halite precipitation, or even after the onset of diapirism, are: (1) the dolomitization of metastable polymorphs of calcium carbonate; (2) the ultra-filtration of shale membrane; and (3) the alteration of clay minerals in the sediments surrounding the salt stock. The mechanism of dolomitization (Mg reacting with limestone), which is typical of the coastal sabkha environment (Bush, 1979), removes a significant amount of Mg in seawater and releases a great deal of Ca, increasing its concentration in the groundwater a t the time of dolomitization. This water may be the formation water which entered the salt stock or was incorporated somehow into the salt diapir. With regard to the process of shale-membrane ultra-filtration, White (1965) suggests that:

cccaz+(mostly from calcite?) is less mobile than C1-, probably because of the double charge of Ca2+,and is enriched in the retained brines.” However, the mechanism of membrane-filtration may not hold much promise (Manheim and Horn, 1968) in accounting for the Ca content of mine brines. In addition to these processes, alterations of clay minerals may result in the depletion of Mg from the original brine. The formation of chlorite from montmorillonite, for example, requires -9.2 mol MgO per mol of chlorite (Eckhardt, 1958). Such a reaction could remove large amounts of Mg from brines. Hiltabrand (1970) has shown that contemporary argillaceous sediments can remove 100mg Mg/l from seawater. The scope of mineral alterations resulting in the enrichment of Ca in brines was investigated by Chave (1960) and von Engelhardt (1960). They compared ocean water with subsurface brines with high Ca content and demonstrated that dolomitization cannot account for all of the Ca in the brines. von Engelhardt (1960) noted that even the formation of chlorite utilizing Mg with an exchange of Na and Ca does not account for all of the soluble Ca; however, exchange reactions with other clays were not considered, as Collins (1975) observed. In view of their chemical character, discussed earlier (p. 133), the mine brines may be the (Ca-rich) formation waters which were originally in the sediments adjacent to the dome, which entered the salt stock and later mixed with and incorporated the residual brine contained in the salt. This residual brine could have been trapped in the original salt mass which may have porosities in excess of 50% (Landes, 1960). For example, the porosities of the playa salt deposit of Uyuni and Coipasa, Bolivia, average 20-30%, and the coarse-grained intervals of salt have porosities as high as 40% (Ericksen et al, 1977). The residual brine trapped after the halite precipitation, may have some Ca concentration (100 mg/l in an evaporating average seawater), but not as much as the mine brines with a Ca content ranging between 9800 and 121,200 mg/l. This implies that there has t o be a richer source of Ca in addition to, or other than, the residual brine. This additional source may be gypsum/anhydrite of salt mass and/or formation water in the sediments surrounding the salt body. On the basis of oxygen-hydrogen isotope ratios, Knauth et al. (1980) have concluded that it is possible, but unlikely, that the

139

mine waters have been derived from the dehydration of gypsum, and that a contribution of anhydrite t o the composition of the mine brines appears very unlikely. Thus, the alternative to the possibility of the admixture of the formation water and the residual brine is actually a formation water alone with unusual concentrations of K and bromide. In fact, the isotopic studies of Knauth et al. strongly support this possibility. Thus, the migration of formation water from the deep sediments into the salt during upward growth of the stock appears t o be a reasonable possibility.

ACKNOWLEDGEMENT

This work was supported by the U S . Department of Energy. By acceptance of this article the publisher and/or recipient acknowledges the U S . Government’s right to retain a non-exclusive royalty-free license in and to any copyright covering this paper.

REFERENCES A.P.H.A. (American Public Health Association), 1975. Standard Methods for Examinations of Waters and Waste Waters. Am. Publ. Health Assoc. - Am. Water Works Assoc. - Pollut. Control. Fed., New York, N.Y., 14th ed., 1193 pp. Atwater, G.I. and Forman, M.J., 1959. Nature and growth of southern Louisiana salt domes and its effect on petroleum accumulation. Am. Assoc. Pet. Geol. Bull., 43: 2592-26622. Bush, P., 1979. Carbonate coastal sabkhas -the precursor of Mississippi Valley type lead zinc deposits. Am. Inst. Min. Eng., Annu. Meet., Feb. 18-22, 1979. Carpenter, A.B., 1978. Origin and chemical evolution of brines in sedimentary basins. In: Proceedings, 13th Industrial Minerals Forum. Okla. Geol. Soc., Norman, Okla., pp. 60-77. Chave, K.E., 1960. Evidence on history of seawater from chemistry of deeper subsurface waters of ancient basins. Am. Assoc. Pet. Geol. Bull., 44: 357-370. Collins, A.G., 1970. Geochemistry of some petroleum-associated waters from Louisiana. U.S. Dep. Inter., Bur. Mines, Washington, D.C., 31 pp. Collins, A.G., 1975. Gepchemistry of Oilfield Waters. Elsevier, Amsterdam, 496 pp. Eckhardt, F.J., 1958. Uber Chlorite in Sedimenten. Geol. Jahrb., 75: 437-474. Ericksen, G.E., Vine, J.D. and Ballon, A.R., 1977. Lithium-rich brines a t Salar de Uyuni and nearby Salars in southwestern Bolivia. U.S. Geol. Surv., Open-File Rep. 77-615, 47 PP. Hiltabrand, R.R., 1970. Experimental diagenesis of argillaceous sediment. Ph.D. Thesis, Louisiana State University, Baton Rouge, La., 152 pp. Knauth, L.P., Kumar, M.B. and Martinez, J.D., 1980. Isotope geochemistry of water in Gulf Coast salt domes. J. Geophys. Res., 85(B9): 4863-4871. Kramer, J.R., 1963. History of the composition of sea water -liquid inclusions compared with a chemical equilibrium model. Geol. SOC.Am. Spec. Pap. 73, 190 pp. Kupfer, D.H., 1974. Boundary shear zones in salt stocks. In: 4th Symp. on Salt, N. Ohio Geol. SOC.,1: 215-225. Landes, K.K., 1960. The geology of salt deposits, In: D.W. Kaufman (Editor), Sodium Chloride. Am. Chem. SOC.,Monogr., 145: 28-69.

140 Manheim, F.T. and Horn, M.K., 1968. Composition of deeper subsurface water along the Atlantic continental margin. Southeast. Geol., 9: 215-236. Martinez, J.D., Thoms, R.L., Smith, Jr., C.G., Kolb, C. R, Newchurch, E.J. and Wilcox, R.E., 1977. An investigation of the utility of Gulf Coast salt domes for the storage or disposal of radioactive wastes. Inst. Environ. Stud., Louisiana State Univ., Baton Rouge, Rep. No. Y/OWI/Sub-4112/37,475 pp. Martinez, J.D., Thoms, R.L., Kolb, C.R., Kumar, M.B., Wilcox, R.E. and Newchurch, E.J., 1978. An investigation of the utility of Gulf Coast salt domes for the storage or disposal of radioactive wastes. Rep. to U.S. Dep. Energy, Inst. Environ. Stud., Louisiana State Univ., Baton Rouge, La., Vol. 1, Rep. No. EW-78-C-05-5941/53,390 pp. Martinez, J.D., Thorns, R.L., Kolb, C.R., Kumar, M.B., Wilcox, R.E. and Newchurch, E.J., 1979. An investigation of the utility of Gulf Coast salt domes for the storage or disposal of radioactive wastes. Rep. to U.S. Dep. Energy, Inst. Environ. Stud., Louisiana State Univ., Baton Rouge, La., Rep. No. E 511-02500-A-1, 572 pp. O’Neill 111, C.A., 1973. Evolution of Belle Isle salt dome, Louisiana. Gulf Coast Assoc. Geol. Sci., Trans., 23: 115-135. Spiro, N.S. and Vouk, T.L., 1961. Changes of the salt composition in the world ocean. Tr. Nauchno-Issled. Inst. Geol. Arktiki Min. Geol. Okhrany Nedr. S.S.S.R., 119: 23-27. von Engelhardt, W., 1960. On the chemistry of the pore solution of sediments. Uppsala Univ. Geol. Inst. Bull., 40: 189-204 (in German). White, D.E., 1965. Saline waters of sedimentary rocks. In: A. Young and J.E. Galley (Editors), Fluids in Subsurface Environments. Am. Assoc. Pet. Geol. Mem., 4: 342366. Zherebtsova, I.K. and Volkova, N.N., 1966. Experimental study of the behavior of trace elements in the process of natural solar evaporation of Black Sea water and SasykSivash brine. Geochem. Int., 3: 656-670.

141

SULFUR AND OXYGEN ISOTOPES AS TRACERS OF THE ORIGIN OF SULFATE IN LAKE CRETEIL (SOUTHEAST OF PARIS, FRANCE)

A. CHESTERIKOFF, P. LECOLLE, R. LETOLLE and J.P. CARBONNEL

Dkpartement d e Giologie Dynamique, Universitk Pierre e t Marie Curie, 75230 Paris Ckdex 05 (France) (Accepted for publication June 23,1981)

ABSTRACT Chesterikoff, A., Lecolle, P., Letolle, R. and Carbonnel, J.P., 1981. Sulfur and oxygen isotopes as tracers of the origin of sulfate in Lake Cr6teil (southeast of Paris, France). In: W. Back and R. L6tolle (Guest-Editors), Symposium on Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54: 141-150. Lake Cr6teil is located “10 km SE of Paris and the preservation of its water quality has prompted a thorough interdisciplinary study, for the lake itself and for the nearby feeding aquifer. Although previous hydrogeological and hydrochemical studies have given indications of the mechanisms involved, uncertainties remained as for the origin, path and history of the waters. This isotopic study of dissolved sulfate (34S, 0)has been carried out in order to find new information about the system.

LOCATION AND GEOLOGICAL SETTING OF THE LAKE

Lake Creteil is located SE of the confluence of the rivers Marne and Seine (Fig. l), in the small alluvial plain limited by these two rivers and the Tertiary outcrops of the Mt. Mesly hill (NE) and the Bois de la Grange Plateau (SE). The geological cross-sections (Fig. 2) show several superposed aquifers: the perched aquifer of the karstic Brie Limestone, then the semiconfined aquifer of the Champigny Limestone which is hydraulically linked with the alluvial aquifer of the small River Yerres in the south and less sharply with alluvium of the Cr6teil plain [these are presently investigated (Fig. 2)]. The “Calcaire Grossier” (Lutetian) aquifer is found -30 m below the latter plain. The lake and other ponds in this plain are located in a former sand pit. Its area is -0.4 km2, the mean depth 4 my and the total volume 1.6 Mm3. At site 2 a small trough exists (-1000 m2 ), which is 6 m deep (Fig. 3). For several years development programs have intended for it to be arecreation area. However, the alluvial aquifer which feeds the lake is completely devoid of oxygen. Natural alluvium, dug out in the past, has been replaced by various infillings, made of clinker, spoiled earth and refuses, which produce a strongly reducing medium.

142

1

Dissolution

( M' Merly

OF

Montmartre's

, plaster ).

Mixing op groups

&

1

and

2

gypsum

.

Sulfate OF Seine , Morne ond /or Chompigny limestone aquifers

Fig. 1. Principal features o f the Creteil area.

Waters are reoxygenated in the lake, while at the bottom and especially during summer the oxygen deficiency still exists: the reducing character of bottom waters is not due to an excess of organic matter ( 5 d a y biochemical oxygen demand, BOD,, and 5-day chemical oxygen demand, COD,, are low, -3 mgl-' and 30 mgl-l , respectively), nor t o an eutrophic state of the lake, which is in fact mesotrophic (Chesterikoff and Testard, 1981).

HYDROGEOLOGICAL DATA

From piezometric studies (Chesterikoff, 1980) it is known that the major

143

W

m y1

N

SENART

FOREST

1

kF G=

c

Loom.

36

Stampion marl . Sonnoision limestone

Ludian

Alluvlum Lower Ludban Monlrnortr? gypsum (in Mt Mesly ).

.

_ _ -

m

Lower

marl.

Ludion

Champigny limestone

Upper Lutetian marl

iyc Aquifer Lutetian PL

limestone “Calcaire grassier*'

81

Fig. 2 . Geologic cross-sections along lines indicated in Fig. 1.

water input comes from the superficial aquifer in the alluvium and recent infillings, from east to west around the south of Mt. Mesly (Fig. 1).It has been shown that there is no contribution from the nearby confined Lutetian “Calcaire Grossier” aquifer. The water itself in the superficial aquifer comes from several sources: (a) The River Marne, of which the water level is rigorously regulated for navigation to the 31.4m MSL height mark. The level of Lake Creteil oscillates from -I-30 to 30.5 m within the season. (b) The semiconfined Champigny Limestone aquifer in the south is a possible contribution, although the input is minimal during dry periods, as most of the output discharges in the River Yerres.

+

144

w

E

ARTIFICIAL

201 I - '

AQUIFER

NNW

PROGRESSION

2

1

SSE

Fig. 3 . Cross-sections of Lake Creteil.

(c) Rains in the Creteil and La Fosse aux Moines alluvial plains southeast of the lake (6 km2 ), which are mostly responsible for the water-table fluctuations. A part of rain water running down the slopes of.the surrounding hills may also contribute to the recharge of the superficial aquifer. Waters from the superficial aquifer and from the lakes in the Creteil plain flow to the west and NW, towards the River Seine, of which the mean watermark level is -2 m lower than the level of the River Marne.

CHEMICAL DATA

Table I presents the mean chemical composition of the lake for the major ions. In this paper only the sulfate ion will be considered. The Mt. Mesly sequence contains the southernmost occurrence of the Bartonian gypsum beds, formerly excavated for plaster manufacturing. Dissolution of gypsum considerably enhances the sulfate concentration of waters in the northern part of the alluvial Creteil plain. Mean samples were taken from piezometers which are shown in Fig. 1 (A-G, L and 201). It was possible to sample separately surface and bottom water from piezometer 201 (201 and 201 ) as well as from piezometer L (surface Ls , bottom L, and medium height L , ). TABLE I Mean chemical composition of Lake Creteil waters

so:c1-

HCO;

590 230 120

20.8

2.0

Ca2+ Mg2+ Na+ K+

250 31 116 38

12.5

52:;

1 .o

1 J

21.0

145 TABLE I1 Characteristics of the waters of this study

Marne Darse Marne

20 40

6 8

9.4 11.6

Seine

18

5

10

Lake Crkteil

1 2

M B

590 620

26.6 32.5

15.7 16.7

Piezometers

201

S B M S B M M M M M S M B

1,760 1,780 1,400 165 785 295 500 1,120 265 460 415 67 0 710 -

27.2 28.0 23.4 15.6 15.3 13.8 12.5 17.0 8.6 12.8 17.6 25.5 26.4

15.2 16.2 15.6 14.4 12.8 11.1 11.3 13.8 11.2 10.7 16.3 17.8 17.0

16.2

14.9

A B

Mt. Mesly gypsum

B = bottom; M = medium height of the water column; S = surface.

TABLE111 Sulfate content of piezometers 203, 206,209 and 211 Piezometers SO:-

(mg 1-'

203

206

209

211

1,150

1,510

820

1,500

Water from Lake Crkteil sampled at site 1 where the water depth is of the uniform type (4m), and at the 6 m deep site 2. An aliquot of samples was precipitated in the form of BaS04 for isotope analyses. Sulfate concentrations are shown in Table 11. Chemical data do not show seasonal variations except for dilution by rain. Piezometer data may be divided into three groups: (A), SO4 concentration >1400 ppm; (B), from 600 to 1120ppm; and (C), l o 0 m), Nabicarbonatesulfate and Na-bicarbonate-chloride types were present. TDS was as low as 400 mg/l in the shallow drift aquifer, generally constant at -1000 mg/l in the deep drift and shallow bedrock aquifer, and over 1700mg/l in the deep bedrock aquifer system. Chemical evolution of groundwater in the area appears to be dominated by two depth zones having different types of water-rock interaction. In the shallow drift zone, the dissolution of soil C O , in infiltrating groundwater, oxidation of organic carbon, sulfur and pyrite result in the formation of carbonic and sulfuric acids that attack carbonate and silicate minerals. On the basis of X-ray diffraction analysis, these minerals were calcite, dolomite, plagioclase feldspar, and smectite clays. However, in the deep regional bedrock aquifer, conditions are reducing (presence of methane), groundwater is alkaline (pH 8.610.3), and the Na-bicarbonate-chloride composition of groundwater is believed t o result from the hydrolysis of volcanic glass or feldspar crystals of oligoclase-andesine composition under conditions of very slow leaching of reaction products and low partial pressure of COz. Under such conditions, calcite and possibly Ca-zeolite are sinks for Ca ion, but Na can accumulate in the pore water. As a result of groundwater movement induced by the postglacial hummocky topography, water from the drift aquifers mixes with water from the deep bedrock aquifers in the groundwater discharge area, yielding a range of intermediate compositions that may be explained by dilution and calcite precipitation, using the MIX2 chemical equilibrium model. Chemical diffusion was shown t o be of negligible importance in comparison with mechanical dispersion to explain the mixing effect.

246

INTRODUCTION

Background and purpose With the institution of large-scale surface mining operations in the province of Alberta and elsewhere on the interior plains of North America, many responsible citizens are concerned that the quantity and quality of local groundwater resources will deteriorate. The excavation and exposure of organic-rich fine-grained overburden materials to oxygen and low humidity can lead to groundwater and soil salinization problems. Questions are posed regarding the length of time required for the groundwater regime t o reach a new state of equilibrium and how the groundwater chemical composition relates to the bulk mineralogy of the exposed overburden. Although answering questions of this kind is fundamental t o the development of predictive models for mining impact studies, the time available in a typical research project is generally too limited to get a good handle on the problem inasmuch as slow reaction rates, low precipitation and pronounced heterogeneity of materials introduce a serious noise problem. The key to answering questions regarding the chemical evolution of groundwater in surface-disturbed areas is t o locate and study a natural surface disturbance of known age. In a field situation of this type, chemical patterns of the groundwater would have developed since the time of the disturbance and the mineralogy of the sediments would reflect the weathering and leaching that had taken place. In this respect, Alberta has a landscape that was covered by more than 1.5 km of ice and re-appeared at the end of the Pleistocene some 10,000 yr. ago. So shallow groundwater flow systems were essentially created at the beginning of the Holocene. In addition, the author had studied an hydraulically closed groundwater drainage basin in east-central Alberta that had been excavated by glaciers into coal-bearing fine-grained rocks similar to those in most strip-mining areas. By radiocarbon dating of the organic carbon in lake sediments cored from a closed lake in the basin, the age of the basin was verified t o be 10,250yr. Also analogous to the strip-mine setting was the thick hummocky moraine deposits that could be compared with spoil piles. The purpose of this paper is to show how chemical evolution of groundwater in a prairie setting that is a strip-mine analogue may be understood by considering chemical weathering reactions involving organic matter, common carbonate and silicate minerals, and transport of weathering products in groundwater flow systems. Much of the material presented here is of a conceptual rather than a quantitative nature, although there is little to prevent the incorporation of the concepts into a hydrogeochemical model.

Previous work A number of significant studies, at least partially devoted to the problem

247

of chemical evolution of groundwater, have been conducted on the interior plains of Canada and the northern great plains of the U.S.A. In Alberta, Le Breton and Jones ( 1962) recognized that groundwater chemical composition could be broadly correlated with changes in geology and climate. Toth (1966, 1968) employed flow-system and hydrochemical-facies concepts to determine the origin of groundwater chemistry in sinall drainage basins in central Alberta. Vanden Berg and Lennox (1969) considered base-exchange, sulfate-reduction and membrane-filtration processes to be important in the chemical evolution of groundwater in south-central Alberta. In a broad-based study of formation waters associated with oil exploration and development, Hitchon et al. (1971) concluded that the ultimate origin of dissolved salts in deep formation waters in the Alberta basin was seawater, and that dilution by freshwater recharge and concentration of ions by shale-membrane filtration are the major factors that control composition. Other modifying processes included dissolution of evaporites, precipitation of minerals, cation exchange on clays, desorption of ions from clays and organic matter, and control of solubility by equilibrium with sparingly soluble salts. In the other areas of the western Canadian plains and in the northern great plains of the U.S.A., similar studies were made, for example, Rutherford (1966), Rozkowski (1967), Freeze (1969), and Davison and Vonhof (1978) in Saskatchewan; Charron (1969) and Cherry (1972) in Manitoba; and Moran et al. (1977) in North Dakota. The processes invoked by these workers t o explain the changes of groundwater chemical type with increasing depth or distance along a flow path included the following: (a) Reaction of carbonic acid with limestone and/or dolomite in the glacial till, leaving Ca2+,Mg2+,and HCO, in solution:

+ CaC0, + Ca2++ 2HCO; 2H2C03 + CaMg(CO,), + Ca2++ Mg2++ 4HCO;

H2C03

(b) Oxidation of pyrite in the presence of calcite, under alternate wet-dry conditions : FeS,,,

+ $0, + H,O

+

Fe2+

+ 2SO:- + 2H+

F e 2 + + 4 0 2+ H + + F e 3 + + & H 2 0 Fe3+

+ 3H20

+ CaC03 Ca2++ SO:-

H+

.+

-+

-+

+

Fe(OH),@,, 3H+

+

Ca2+ HCO; CaSO,(,,

(c) Dissolution of gypsum t o produce CaZ+and SO:-:

+ SO:- + 2 H 2 0

CaS04-2H20-+ Ca2+

(d) Loss of Ca2+and Mg2+ and gain of Na+ by cation exchange with Narich smectite clays:

248

RNa;

+ Ca2+or Mg2+ =+ Rca~+,Mg2+ + 2Na+

(e) Once the partial pressure of oxygen decreases t o a level sufficiently low for sulfate reduction to occur, loss of sulfate and gain of CO, take place in the presence of Desulfovibrio desulfuricans bacteria: SO:-

+ 2CH,O

+

H S - + HCOi

+ HzC03

where CH,O represents organic matter (HS- is removed from solution through reactions with ferrous iron t o form FeS and then pyrite, FeS,). (f) Addition of C1-through ionic diffusion from deeper more saline formation waters or through dissolution of halite. With the exception of cation exchange on clay minerals, none of the previous workers considered the role of chemical weathering of silicate minerals in the chemical evolution of groundwater on the plains. Silicate mineral weathering is believed t o be a key-process in the chemical evolution of groundwater in a surface disturbed area where the bedrock contains little carbonate. According t o Curtis (1981), the key weathering reactions involve the formation of hydrogen ions in subsurface water through dissolution of CO, and production of sulfuric acid by sulfide oxidation. Further, in wet, cool climate (such as on the interior plains), carboxylic acids and phenols associated with degradation of organic matter can also contribute significantly to hydrogen ion production. Curtis (1977) stressed that anion production reactions, especially for bicarbonate and sulfate, are the limiting factors in chemical weathering. Hydrogen ions are used up in reactions of the type:

+ H+

silicate+-M

+

silicate+-H

+ M+

where metallic cations are removed in solution; silicate=O

+ H, 0

-+silicate=(OH),

where the silicate “core” is hydrated, and: silicate=Si=O

+ 3H20

silicate=(OH),

-+

+ Si(OH)4

where silicic acid is produced. According to Curtis (1977), the products of chemical weathering which are the soils (or any altered parent-rock regardless of depth) and aqueous solutions form mainly as a result of three processes: (1)acid-induced metalcation leaching; (2) hydration; and (3) oxidation of ferrous t o ferric iron. Plan o f attack With the latter background t o the problem and the framework for chemical weathering in mind, the author would like to outline the plan that will be followed for the remainder of this paper: (1)the study area will be described from the standpoint of hydrogeologic environment (geology, topography and climate) and the hydrodynamics and general aspects of the groundwater

249

chemistry will be given; (2) the mineralogic composition of the bedrock and drift sediments will be presented, sources of carbon and sulfur and hydrogen ion will be specified based upon carbon and sulfur stable-isotope ratios, and the various silicate and carbonate mineral weathering reactions that consume hydrogen ions will be discussed in detail by means of stability-field diagrams; and (3) the chemical evolution of groundwater in the bedrock and in the drift aquifers will be treated individually, and in the context of a mixing model to explain the variability of groundwater chemical types encountered in the basin. DESCRIPTION OF THE STUDY AREA

Location The location of the study area is shown in Fig. 1A. The greater area, which includes the associated regional and local groundwater basins, is bounded by 110°17'-111"10' W long. and 52°05'-52036'N lat. In all, the size of the regional basin (Fig. 1B) is -500 km2 in contrast t o the size of the local basin (Fig. 1C) which is 33 km2 . The local basin is located within a large region of interior drainage that straddles the provinces of Alberta and Saskatchewan (Fig. 1A).

-

Climate and topography Mean daily temperature of the study region (Fig. 1B) ranges from -15°C in January t o +18"C in July. Potential evapotranspiration of 530 mm/yr. greatly exceeds the mean annual precipitation of 373 mm/yr. These data are from a compilation of hydrogeologic data for the area by Hackbarth (1975). The area therefore has a cold semi-arid continental climate. The general nature of the regional topography is shown in Fig. 1B. The Neutral Hills end moraine in the southwest and west (elevations t o 830m above sea level) slopes northeastward onto a broad undulating plain averaging 650 m above sea level in elevation. As previously noted, surface drainage in the area is poorly integrated or non-existent, especially in areas of knoband-kettle topography. A few of the larger kettles are shown on Fig. l b by the hatched depressions. The direction of the last ice advance during the Wisconsin was from northeast t o southwest, an observation based upon the trend of the long axes of landforms such as the Neutral Hills end-moraine, and the capturing of Ribstone Creek in the north-central portion of the map area. Arcuate landforms, glacially contorted bedrock, and closed depressions in the bedrock surface indicate that glacial ice thrusting was an important process in generating the topography, criteria given by Christiansen and Whitaker (1976). A blowup of the central area in Fig. 1 B is shown in Fig. 1C. Although only the water-table contours are given in this figure, the essential features of

-

250 T o p q rap hy

Water T a b l e

0

1

2

3

4

LEGEND

KH

-6823-8'

CONTOURS I N METERS ABOVE SEA L E V E L HYDROGEOLOGICAL CROSS S E L I I U N L O C A T I O N OF WELL

OR

P I E Z O H E T E R NEST

Fig. 1. A. Location; B. regional topography; and C. local water-table topography.

the topography are nonetheless clearly visible inasmuch as the water table is a subdued replica of the surface topography. A closed depression is flanked on the west and east by highlands. Topographic relief in this locality is substantial for a plains' setting. Attention is drawn to the trends and locations of two cross-sections: section A-A' which is aligned parallel to the direction of glacial ice movement, while section B-B' follows the predominant local west-east topographic slope.

251

Geology

The area portrayed in Fig. 1B is underlain by bedrock of Late Cretaceous age, which in turn is overlain by glacial till and surficial deposits. The bedrock geology of the area was summarized by Green (1972). Two bedrock units subcrop in the area: the Bearpaw Formation and below this, the Belly River Formation. The strata dip gently t o the southwest at -1 m/km. Lithologic descriptions for the two units are as follows: Bearpaw Formation: dark-grey blocky shale and silty shale, greenish glauconitic and grey clayey sandstone; thin concretionary ironstone and bentonite beds, marine (Green, 1972) Belly River Formation : complexly interbedded pale-grey, bentonitic sandstone, laminated siltstone, and medium- to dark-grey carbonaceous claystone, thin-bedded darkbrown weathering sideritic ironstone and coal; coastal plain or deltaic, notable gas producer (Locker, 1973)

The Lea Park Formation is the lowermost unit of hydrogeological interest in the area but it is covered by at least 300m of Belly River sediments. The strata are notable for their low permeability and consist of marine deposited grey silty shale, thin sandstones, and common sideritic ironstone concretions (Green, 1972).

701 610

457 LEA PARK FM.

0 A

884 162

610

701 610

457

457

8

0

16 km

LEGEND WATER TABLE

HYDRAULIC HEAD .---701----

(METRES ABOVE SEA L E V E L )

-1700-

TOTAL D I S S O L V E D S O L I D S mg/l

Fig. 2. A. Regional stratigraphy and hydraulic-head distribution. B. Regional distribution of total dissolved solids (TDS).

252

The bedrock sequence is everywhere overlain by glacial drift, ranging in thickness from 1 to 6 0 m . Glacial deposits consist of interbedded till, sand and gravel, and lacustrine clays derived primarily from glacial and glaciofluvial erosion of the poorly competent Belly River Formation. Although the bulk of the drift material is of local origin [as much as 85% according to Bayrock (1967)], it nevertheless contains a calcite and dolomite component derived from the Lower Paleozoic limestone and dolomite deposits that surround the Precambrian Shield. The direction of last ice movement during the Wisconsin Epoch was from northeast to southwest, or from right to left along section A-A’ (see Fig. 2A). Massive deformation of the topography and stratigraphy by the moving ice mass is indicated by the “gouge-and-step” nature of the land surface and by the southeastward displacement of the Bearpaw Formation to form the Neutral Hills end-moraine. H y d rogeology

The regional distribution of hydraulic head was determined from surveys of water-level elevations in domestic wells, Research Council of Alberta test wells, and piezometers located along section A-A‘. As shown in Fig. 2A, hydraulic gradients are downward everywhere along the section except near the closed lake where upward flow gradients persist to a depth of -100m below land surface. The apparent vertical hydraulic gradient decreases from southeast t o northwest, from left t o right on the diagram. As shown in Fig. 2B, total dissolved solids (TDS) contours are generally parallel to the hydraulic-head contours. The fact that the 700-mg/l contour is deeper in the left-hand half of the cross-section than on the right-hand half indicates that recharge to the regional flow system occurs primarily under the area of the high plateau. The patterns are consistent with the expectation that TDS increase both in the direction of groundwater movement and with increasing depth. The local distribution of hydraulic head and the approximate flow pattern along section B-B’ are given in Fig. 3. The section traverses a closed groundwater basin (note closed 670-m contour in Fig. 1C) with a closed saline lake occupying the lowest elevations of the basin. The flow of groundwater is generally downward through the glacial drift into the Belly River Formation except at the closed lake where groundwater moves toward the surface. Based upon empirical head data and surface groundwater features such as springs, seeps, marshes, hydraulic character of ponds (sloughs), three orders of groundwater flow systems were deduced in the area: (1)shallow and transient flow systems associated with knobs and potholes of the hummocky glacial topography in the west half of the basin; (2) an intermediate flow system on the west and a local flow system on the east that terminate at the closed lake; and (3) a deep regional system that bypasses the lake and flows toward the northeast, parallel to the regional topographic slope. It is also a characteristic within the basin that bulk hydraulic conductivity

253

decreases exponentially with increasing depth. Based upon a suite of aquifer tests on each of the piezometers, the hydraulic conductivity was seen to vary cm/s in the most productive aquifer at the eroded driftfrom -4 bedrock contact zone, to -7 * cm/s in the sandy lenses of the upper Belly River Formation, and t o -2 cm/s in the deep sands of the Belly River that constitute the regional aquifer system. The distributions of TDS and hydrochemical facies reflect the groundwater recharge flux. Where recharge is good, the flushing of weathering products is most active and TDS are apt to be low. In recharge areas, groundwaters of low TDS will persist t o greater depths than in zones of lateral flow or discharge. The distribution of TDS and hydrochemical type of groundwater along section B-B' is shown in Fig. 4.In the west half of the section, water storage in depressions resulting from the hummocky topography contributes to effective recharge of the local and intermediate flow systems, and therefore, the TDS in the groundwater at a given depth below water table is less than in the east half of the section, which is not an effective recharge area. The marked increase in TDS at a depth of -60m below the water table in the west half of the cross-section coincides with the boundary between the intermediate and regional groundwater flow systems. The several orders of groundwater flow system are reflected in the distribution of chemical types of groundwater. In the shallow drift, Ca-MgHC0,-type groundwaters are most common, whereas in the deep drift and shallow bedrock, the groundwater is generally of Na-HC03-S04 type. In

-

-

701

701

670

670

Y

640

640

610

610

P 579

579

4, >

G

549

549

5181 488

1518

'

, 1.6

4.8

3.2

6.4

I

488

8.0

(km)

- LEGEND --- -

- 667

D R I FT/BEDROCK

CONTACT

HYDRAULIC HEAD (METRES ABOVE SEA L E V E L ) INFERRED GROUMDWATER FLOU D I R E C T I O N

0

FLOW I N T O PLANE OF CROSS S E C T I O N GLACIAL D R I F T

_4 WATER TABLE

Fig. 3 . Local hydraulic head and approximate flow pattern.

2 54

I LRR

I .'6

3.2

418

6.4

8.0

(km)

LEGEND

WATER TABLE

-

TOTAL D I S S O L V E D S O L I D S (rngl9.)

-600-

....

~

-

~

HYDROCHEHICAL ZONES

I

NaHCO ( G I )

I1

NaHC03S04

111

3

HCOj { Shallow CaHg nco3s04 ~~~p ~a

YELL/PIEZOHETER POINT

Fig. 4 . Local distribution of chemical type and total dissolved solids.

the deep bedrock, the groundwater most often is of the Na-HC03 or NaHC03-C1 type. Essentially, this is the common progression of groundwater chemical facies observed in most studies on the interior plains. The following section examines the origin of these groundwater chemical types by considering chemical weathering and leaching of the rock-forming minerals of the aquifers and aquitards of the drainage basin.

ROCK-WATER INTERACTION

Mechanisms by which groundwater chemistry changes with depth or along flow path can be elucidated by considering the bulk mineralogy of the sediments and the major weathering reactions that take place within the flow region. Bulk mineralogy includes not only the present mineral assemblage (and organic matter, e.g., coal) but also the suite of primary minerals that were originally deposited t o produce the parent-rock. Whereas in wellleached highly-permeable groundwater systems, the relatively slow diagenetic reactions can be ignored without serious consequences, this is not true when chemical weathering is slow due t o cool climate, and groundwater circulation is limited by shallow relief and a semi-arid climate. Insofar as the study area was glaciated and new groundwater flow patterns developed as a consequence of the altered topography, one can surmise that the observed groundwater chemical patterns result from chemical diagenetic processes taking millions of years (regional groundwater flow system - NaHC03 type) and

255

weathering, leaching, and ionic transport processes that began only 10,000 yr. ago (local and intermediate groundwater flow systems30,320 (c) >30,330 (c) 29,990 k 1,930 (c)

- 6.2 -12.4 - 7.6 - 7.9 -11.4

+ 0.1 ( d ) f 0.1 f 0.1 f 0.1 f 0.1

(d) (d) (d) (d)

f 0.83

-18.9 + 0.2 -20.6 f 0.2 -23.0 f 0.1 -16.8 + 0.1 -21.8 f 0.2 -19.9 + 0.3 -21.3f 0.1 -21.1 f 0.2 -21.2 + 0.1

f4

+4

X = -141

SMOW)

98 f 5 228 f 4 158 + 6 Of3 4 f 3 45 f 4 Of4 6+4 3f4

+ 1.8

Laboratories: (a) Atomic Energy Canada Ltd., Chalk River, Ontario, courtesy R.M. Brown; (b) Department of Earth Sciences, University of Waterloo, Waterloo, Ont., courtesy P. Fritz; (c) Saskatchewan Research Council, Saskatoon; (d) University of Saskatchewan, Saskatoon, Sask., and remaining analyses performed at Weizmann Institute of Science, Rehovot, Israel.

N D

w

TABLE I11

N

6 %S in sedimentary sulfate

N

Lithology

cn

6% ("/&I,

Depth (m)

Lithology

(ft.)

OBS-1: Greengrey clayey sand Green-grey clayey sand Greengrey clayey sand Green-grey clayey sand Greengrey clayey sand Green-grey clayey sand Dark-brown clayey silt, coal, Fe-nodules Dark-brown clayey silt, coal, Fe-nodules Dark-grey clayey sand Dark-grey sandy clay Dark-grey sandy clay Grey clayey sand and gravel Green grey clayey sand Green-grey to dark-brown shaley sand and siltstones, coal, limonite staining Dark-brown silty clay Dark-brown silty shale Dark-brown silty clay

f0.9 -5.1 -2.2 $0.5 -1.1 4- 6.8

1.5 3.0 4.6 6.1 9.1 10.7

5 10 15 20 30 35

OBS-2 : green-grey sandy clay green-grey clayey sand green-grey clayey sand green-grey clayey sand green-grey sandy clay green-grey sandy clay

-11.0 -16.5 -12.0 -10.9 - 11.3 -6.6

$2.1

12.2

40

green-grey sandy clay

-6.6

15.2 16.8 22.3 24.4 27.4 30.5 36.6 39.6 42.7 61.0 64.0 65.5

50 55 73 80 90 100 120 130 140 200 210 215

-1.4 -2.8 -10.8 -16.9 -

-

$7.8 f4.8

green-grey sandy clay green-grey clayey sand dark-brown/green silty clay dark-green/brown silty clay green clay, Fe staining, coal dark-greenlbrown clay dark-brown/green clay green-grey sandv clay dark-brown silty clay, coal green-grey clay green-grey sandy clay green-brown clayey medium gravel

-

~

-10 .a -10.9 -40.1 -29.2 -12.5 -13.4 $1.4 -4.9 -8.7 -10.4 -36.2

'

i = [--15.1 f10.31

mean 3c = [-12.1 + i i . a ]

68.6

RAH-4 : Green-grey coarse sandy till Green-grey silty shale Green-grey/brown shaley sandstone Grey/greenish-brown shaley sandstone and coal Grey silty shale, coal Grey fine sandstone and coal Grey-green shaly siltstone and coal Grey silty shale Grey-brown slightly silty shale Grey shaley fine sandstone

225

+2.6 +9.2

1.4 1.7

4.66 5.67

+7.2

2.3

7.67

+8.1

3.1

9.00

3.1 3.3 3.6 4.1 5.4 7.7

10.33 10.83 11.92 13.33 17.83 25.08

f9.6 +11.5 +11.6 +4.9 f7.3

+ 12.0

Green-brown clayey very coarse sand HSL-2 : Soft greenish black To grey organic clay

+12.4 +12.2 +13.2

muck with medium to coarse crystals of salt dense crystal and black mud and crystal with pockets of H2S greenish grey very soft organic muck with a base of organic sandy clay

+17.2 +5.8 +15.7 f15.6 +10.9 f11.9

+4.9

264

Analyses were carried out in the laboratory of H.R. Krouse, Department of Physics, the University of Calgar (Table 111). The average value of 634S for sulfate leached from the bedrock sediments is 8.4 -+3%,, while the average value for the drift is -13.9 k10.7?oo. The value for the bedrock therefore differs substantially from the + 30.3YoO mean value reported by van Everdingen and Krouse (1977) for marine evaporite gypsum in the Lower Devonian Bear Rock Formation in the MacKenzie District of the Northwest Territories and the value of +34%, for SO:- in groundwater passing through the Middle Devonian evaporite in the Athabasca Oil Sands area (H.R. Krouse, pers. commun., 1980).The mean value of 634S is also considerably more negative than that of seawater sulfate, reported by Sakai (1957) as +20.77yoO. Recent analyses of 634S of organosulfur compounds in Alberta coal samples collected by the Alberta Research Council ranged from -1 t o 13yoO , and averaged 5yoo (H.R. Krouse, pers. commun., 1980). The average value of 634Sfor the bedrock is therefore more similar to that of coal rather than that of marine evaporite or seawater sulfate. The likely source of the sulfate in groundwater is therefore the organic matter and the very finely disseminated pyrite that are present in the thin coal seams and clayshale of the Belly River Formation and are incorporated into the drift. Oxidation of these compounds in the presence of calcite and/or dolomite gives rise t o the formation of Ca2+ and SO:- ions or secondary gypsum that dissolves during flow of groundwater through the soil and unsaturated zones during a recharge event.

+

+

Silicate mineral reactions Because analyses for alumina in groundwater were not available, it was not possible t o calculate saturation indices for congruent reactions by means of SOLMNEQ (Kharaka and Barnes, 1973). However, the ion-activity ratios computed in SOLMNEQ for groundwaters from drift and bedrock aquifers were plotted on a number of stability-field diagrams (Helgeson et al., 1969a). It should be noted that all of the diagrams that are presented in this paper are for systems at 0°C. Inasmuch as the groundwater temperatures ranged from 3 to 18"C, and averaged 11.8 k 5"C, a small error is associated with the use of the diagrams. According to Garrels and Christ (1965, p.261), a temperature change of a few degrees does not alter the activity diagrams more than the width of the line used t o indicate the phase boundaries. They gave as an example the equation that describes the Eh-pH boundary between magnetite and hematite. The equations for the boundaries are: Eh = 0.221 - 0.059pH,

at 25°C;

Eh = 0.227 - 0.061pH,

at 35°C

The differences are quite negligible.

and

265 I

6l og

LNa+l [H+l

0

GIBBSITE

5-

I,

. 0 BEDROCK

0 DRIFT

I

-6

-5

.

-4

-3

log I S i O * l

Fig. 10. Stability-field diagram (Na20-Al~03-Si0~-H20).

Na20-A1203-Si02-H20 system. As shown in Fig. 10, groundwater compositions from drift aquifers fall exclusively in the kaolinite field, whereas the groundwater compositions from the bedrock aquifers are in both the gibbsite and kaolinite fields. The cluster of points grouped in category III represents samples obtained from the deepest bedrock aquifers. These points fell farthest into the gibbsite field. Samples grouped under the R category were obtained from water-table wells in the glacial drift in the recharge area (viz. left-hand half of section B-B') and these are displaced farthest from the montmorillonite stability boundary. Samples that plotted close t o the montmorillonite stability-field boundary were generally from drift and shallow bedrock wells situated in the groundwater discharge area near the closed lake.

Discussion. The latter distribution of points may be compared with trends GH and IMJ shown in Fig. 11 (after Helgeson et al., 1969b). Trend GH follows the distribution of groundwater compositions obtained from aquifers in the Sierra Nevada Mountains, reflecting the weathering of albite t o kaolinite in a well-leached environment. This trend is similar to the one observed for groundwater from the drift and shallow bedrock aquifers in Fig. 10. Essentially, Na is increasing while the concentration of silica remains limited within half a log cycle, and the pH is increasing with higher Na concentration as groundwater moves from the recharge area to the discharge area. Kaolinite is the solid phase that is in equilibrium with these groundwater compositions. Trend IMJ describes the path followed by the reaction of clay

266 9

1

' I "

KAOLlNlTE,

Fig. 11. Compositions of waters in the Na20-A120~-SiO~-H~0 system at 25OC, unit activity of water, and 1atm. (after Helgeson et al., 1969b). The stability-field boundaries shown for montmorillonite are thermodynamically consistent with one another, but they are based on analyses of waters issuing from sediments that reportedly contain coexisting montmorillonite and kaolinite. Irreversible reaction paths (dotted and dashed lines) are shown in the diagram for the hydrolysis of albite (ABCDEF)and coexisting K-feldspar and albite with relative reaction rates of 1:I (A'B'C'D'E'F'G'H'I') and 0.1 :1 (A"B"C"D"E"F"), weathering of Sierra Nevada rocks ( G H ) , and reaction of clay minerals with Bermuda seawater ( I J ) . The area labeled M designates the composition of surface seawater and point N represents the average composition of world streams.

minerals with Bermuda seawater, and as such, reflects the type of reaction that occurs in a sluggish chemical diagenetic system. This trend is similar to that observed for the groundwater compositions from the bedrock aquifers that are part of the regional aquifer system. To explain why the groundwater compositions from bedrock aquifers fall within the gibbsite field, it is necessary t o consider the manner in which diagenesis affects the relative mass of the secondary minerals as a function of reaction progress (that in turn is a function of time, depth, pressure and specific surface area). Fig. 1 2 (after Helgeson et al., 1969b) is a schematic diagram that shows the paragenesis and relative mass of authigenic minerals produced by the hydrolysis of coexisting K-feldspar and albite. The figure shows that as the time of burial of the minerals increases, the principal type of authigenic mineral formed changes from gibbsite t o kaolinite t o K-mica to Na-montmorillonite. It is implicitly assumed that the sediment-water system is closed; that is, that no intermediate reaction products are physically removed from the sediment-water system. In the early stages of diagenesis, gibbsite is the phase that will precipitate and enter into chemical

267

OVERALL EQUILIBRIUM ESTABLISHED GIBBS I T E

1

C'

B' -H'mOLINiTE K-MICA K- F E LDS PAR Na-fiONTHORILLONITE

Fig. 1 2 . Paragenesis and relative mass of authigenic minerals in hydrolysis of coexisting K-feldspar and albite (after Helgeson et al., 1969b). A L P I N E M O U N T A I N MEADOW S O I L

Fig. 1 3 . Co-evolution of mineralogy and groundwater chemistry (after Kovda and Samoilova, 1969).

equilibrium with the pore water. In eastcentral Alberta, the sediments of the Belly River Formation are unconsolidated, poorly cemented, and generally immature in terms of diagenetic alteration. Therefore, the presence of the groundwater compositions from bedrock aquifers in the gibbsite field is consistent with the theory that diagenetic changes in the sediments are controlling the chemical composition of the pore water in the bedrock. Differences in weathering regimes between well and poorly leached parts of a groundwater flow system have been noted previously in the literature. For example, Kovda and Samoilova (1969) presented a diagram (Fig. 13) showing how kaolinite tends t o accumulate in uplands and montmorillonite in lowlands as a consequence of groundwater flow. The formation of montmorillonite results from the increased availability of silica so that a reaction of the following type prevails: 1.17Al2Si2O,(OH), kaolinite

+ 0.167Ca2+ + 1.33H4SiO4+

268

Cao.167A12.33Si3.67010(OH)2 + 0.33H'

+ 3.83H.20

Ca-montmorillonite

Whereas in humid environments the Na is flushed from the soil and unsaturated zones, in more arid climates the concentration of Na in the groundwater is high due to the precipitation of calcite, cation exchange, and the higher ionic mobility of Na with respect to Ca, and Na-montmorillonite tends to be the stable solid phase in lowland areas and with increasing depth in the groundwater system.

CaO-A120,-Si02-H,0 system. The compositions of groundwater from drift and bedrock aquifers are plotted in Fig. 14, a stability-field diagram for the CaO-A1203-Si02-H20 system. The groundwater compositions occur in the leonhardite (variety of laumontite) field with water from the drift aquifers tending t o be undersaturated with respect t o calcite at the atmospheric partial pressure of CO,, and water from the bedrock tending t o be oversaturated. The data indicate that a chemical potential exists for the precipitation of leonhardite and calcite.

Fig. 14. Stability-field diagram (Ca0-A120~-Si02-H20).

Discussion. Oki et al. (1977) explained the presence of high-pH groundwaters (pH9--10.3) in the Tanzawa Mountains of Japan as being the result of the hydrolysis and incongruent dissolution of Ca-plagioclase with minor supply of C 0 2 in the deep subsurface system and the very slow rate of formation of Ca-zeolites such as laumontite. The authors stated that the groundwaters were extremely undersaturated with respect t o the natural atmospheric partial pressure of CO, . Formation of laumontite was attributed to the further reaction of secondary montmorillonite:

269

+ 6H4SiO4 + 1 6 H 2 0

3Cao.33A14,67Si7.33020(OH)4 +- 6Ca2+

=$

Ca-montmorillonite

7CaAl2Si4OI2* 4H2O

+ 12H'

laumontite

It was shown that laumontite rather than anorthite is stable at ordinary temperatures (25°C). Although laumontite can crystallize at ordinary temperatures, the rate of this reaction is slow enough so that hydrogen ions released are adsorbed on the montmorillonite reactant, thereby resulting in alkaline groundwaters. Samples of drill cuttings from the glacial drift and the Belly River Formation were analyzed by means of the X-ray powder camera method in an attempt t o find laumontite. The results were negative. The only location in Alberta where laumontite has been found is in the Cretaceous Blairmore Group in the folded foothills in the southwestern corner of the province (Miller, 1972). What these observations indicate is that somewhat higher temperatures and pressures are required t o form significant quantities of the zeolite laumontite than are available in the study area.

system. In the stability-field diagram for the Mg0-K20-A1203-Si02-H20 system (Fig. 15),it may be noted that all of the compositions of groundwater from the drift aquifers fall within the kaolinite stability field. However, compositions of groundwater from the bedrock aquifers also plot in the illite and Mg-chlorite fields. One possible interpretation of the distribution of these points is that a continuous spec-

MgO-K20-AE20,-SiOz-HzO

KAOLINITE

0 B E D R OC K

6.0

.DRIFT

4.0

2.0

6.0

4.0 log

0.0

I0 0

[Ktl [H+l

Fig. 15. Stability-field diagram (Mg0-K2O-M2O, -Si02 -H20).

270

trum exists between illite breakdown in the weathering and leaching process, and illite formation during diagenesis. This is t o say that in the more active regions of a flow system, illite would tend t o weather t o kaolinite, while in the deeper more stagnant zones, illite is forming in the aquifer/aquitard sediments. Two reaction mechanisms are postulated as being important in the release and the solubility control of K and Mg in the groundwater: the first is the formation of illite by the reaction of biotite, K-feldspar and montmorillonite during diagenesis. Illite is a very stable secondary mineral, based upon the fact that K-Ar ages of illites tend t o remain constant during weathering and erosion in temperate climates, indicating either little K loss, or loss of K and Ar in their whole mineral ratio (Garrels, 1976). Therefore, the following reaction proceeds in the forward direction only: KMg3A1Si3010 (OH),

+ 5co2

18H2O

+ KA1Si308 + NaG.128Ca0.079A12.33Si3.67010 (OH), K0.6Mg0.25A12,3Si3.s0i0(OH),

1.4K+

+ 2.75Mg2+

+ 0.128Na' + 0.079Ca2+ + 2.03AI(OH), + 6.17H4Si04 + 5.16HCO; The mold K/Mg ratio in pore waters where the transition is taking place would be 0.51. The second reaction consists of the weathering of illite to kaolinite :

+

K 0 . 6 M g 0 . 2 5 ~ 2 . 3 0 S i 3 . s 0 1 0 ( O H )1.1H' 2 4-3.15H20 =+1.15A12Si205(OH),,

+ 0.6K' + 0.25Mg2' + 1.2H4Si04 The K/Mg ratio in pore waters where the latter reaction occurs would be 2.40. The distribution of values of the K/Mg ratio in groundwater from the drift aquifers is presented in Fig. 16. The mean ratio is 0.18 kO.12. For ground-

I8 16

f

14 12

GROUNMIATER FROM DRIFT AQUIFERS

10

x

-

0.18 f 0.12

8 6

4 2

. . .

N * U )

A A . &

(K+)/ (Hg++) (k)/ (I%++) Fig. 16. Ratio of K+/MgZ+(epm) in groundwater from bedrock and drift aquifers.

271

water in aquifers from the bedrock, the ratio is -1.0, as shown in Fig. 16. The ratio of exchangeable K t o Mg in the drift averages 0.4 kO.08. These observations suggest that the solubility of K t o Mg is probably related to clay-mineral transitions although the exact nature of the reactions remains unknown. It may be noted that the concentration of K+ does not differ significantly between groundwater from the drift and the bedrock aquifers (5.7 t 2.6 mg/l vs. 4.8 ? 2.7 mg/l, respectively). The higher Mg concentration in the groundwater from the drift aquifers, which are known t o contain abundant dolomite, presumably accounts for the lower value of the K/Mg ratio.

Breakdown of plagioclase to montmorillonite. The breakdown of plagioclase crystals or glass fragments in volcanic ash during chemical weathering may be described by the reaction:

1.37Nao.62Cao.38A11.38Si2.6208 + l.66C02 + l . 6 6 H 2 0 =+

o.8~Nao.,28Cao.079A12~33Si3~67010(OH)2 + 0.745Na' + 0.46Ca2+

+ 0.61Si02 + 1.66HCOi In the above equation, the composition of the montmorillonite was adjusted so that the Na/Ca ratio is the same as the parent-feldspar. This assumption is based upon the observation that the mold ratio of exchangeable Na t o Ca in the montmorillonite of the Belly River Formation is 1.5, a value derived from the data reported in Table IV. When the reaction occurs in the sediments, pore waters will be enriched in Na', Ca2+,HCO; and S i 0 2 . Inasmuch as the groundwater chemical analyses presented in Table I indicate that the concentrations of Ca2+and SiO, are TABLE IV Exchangeable cations in drift*' and in the Belly River Formation*2

AH-2 (9-1.4 m) AH-5 (16.3 m) AH-6 (13.1 m) (14.6 m ) (19.2 m) AH-7 ( 0 . 9 m ) Belly River Fm.

*'*2

Ca2+

Mg2+

Na+

K+

CEC*3

5.2 40.2 30.5 28.2 38.9 2.8 16.5

1.4 3.4 3.6 3.5 3.5 0.8 5.1

0.04 0.04 0.13 0.13 0.12 0 12.4

0.2 0.7 0.72 0.9 0.69 0.15

5.6 7.0 7.3 7.0 7.0 3.7 28.8

-

Analyzed by Soils Division, Alberta Research Council. Average of seven samples from Locker (1973). *3 Cation exchange capacity. Note that CEC is not equal to the sum of the exchangeable cations for most of the drift samples because of contribution from calcite and dolomite to the exchangeable Ca2+.

272

+I

I

I

0 0

O -0

ooo -1

0

" O

DL""

0 DRIF

I

0

-5

-4

-3 log

-2

-I

pco2 ( a m )

Fig. 1 7 . log [CaZ']/[Mg2+] vs. logPC0, for groundwater from bedrock and drift aquifers.

low compared with those of Na' and HCO;, the question arises as to what is the fate of Ca2+and SiOz in sediment-water systems. One mechanism for the removal of Ca2+ is the precipitation of calcite, already noted as one of the principal cementing minerals. in the bedrock. The fact that groundwaters in the drift and bedrock aquifers are generally oversaturated with respect to calcite (see Fig. 5) supports the existence of this mechanism. The saturation indices indicate that calcite either precipitates from solution or simply does not dissolve. Fig. 17 is a plot of the log of the activity ratio of [Ca2'] /[Mg2'] vs. log of the equilibrium partial pressure of CO, , using computations made with SOLMNEQ. Apparently, the [Ca2+]/ [Mg"] ratio of groundwater from the drift aquifers is controlled by calcite and dolomite in as much as similar concentrations of Ca2+and Mg2+are in solution. In the bedrock aquifers, precipitation of CaC03 likely accounts for the decrease in the [Ca"] /[Mg2'] ratio and the drop in partial pressure of CO,. Another important mechanism for removal of Ca2+ion from pore waters is cation exchange. It is true that the montmorillonite in the bedrock is not especially "Na rich" and the montmorillonite in the drift is actually enriched in Ca, according to the exchangeable cation data in Table IV. So, how does the exchange loss of Ca and gain of Na take place? One point that must be raised immediately is that the mass ratio of ions in solution t o ions on the exchange complex is very small, and therefore considerable exchange can take place without noticeably altering the clay composition. Another point is that in the course of chemical weathering and redistribution of reaction products by fluid circulation, the Na+ ion are more mobile than the Ca2+ion. Then, similar to the situation described by Kovda and Samoilova (1969), a segregation between Ca-rich montmorillonite in the leached recharge areas and Na-rich montmorillonites in the discharge areas will take place. The exchange capacity of the montmorillonite therefore becomes a function of the age of the groundwater flow system and the position within the flow system. In view of the latter considerations, further study is required before an understanding of the relative importance of cation exchange vs. calcite precipitation can be assessed in the area.

273

Removal of silica from solution likely occurs by precipitation as amorphous silica. From Table I, groundwater from the drift aquifers averaged 12.8 5 4 mg/l silica while groundwater from the bedrock averaged only 3.8 4 mg/l silica. This difference probably reflects precipitation loss of silica when groundwater from the drift enters the bedrock.

*

Origin o f groundwater composition Na-HC03-type groundwater from the bedrock aquifers is seen as the composition that would result from the consumption of hydrogen ions by the chemical weathering of feldspar t o clay plus calcite/siderite. Hydrogen ions are introduced t o the sediment-water system at depth by means of sulfate reduction and methanogenesis, or in the near-surface environment by oxidation of organic matter and pyrite. The rate of chemical weathering is considered t o be controlled by the rate of hydrogen ion production. In hilly areas such as in the hummocky moraine of the study area (or in spoil piles of surface mine sites) the unsaturated zone is thickest; oxygen can reach the fresh rock materials, and CO, and SOz can dissolve in water t o yield carbonic and sulfuric acids. In relatively flat areas, the unsaturated zone is fairly thin and hydrogen ion production will be less. So, in the study area, the rate of chemical weathering in the bedrock aquifers is thought to be a function of the surface topography. The Ca-Mg-HC0,and Ca-Mg-S04 -type groundwater from the glacial drift aquifers reflects the dissolution of calcite and dolomite by carbonic acid formed in the soil zone, and the production and leaching of secondary gypsum through oxidation of sulfide in the presence of calcite or dolomite under conditions of partial saturation. In cases where the content of carbonates is low, silicate mineral weathering potentially occurs.

MIXING O F DRIFT AND BEDROCK GROUNDWATER

Conceptual model of mixing The present-day topography and distribution of drift sediments are comparatively recent (3D,OOO YEP

-16.8 t o -22.9

-125

to -167.8

na

Poorly mixed

- 1 8 . 9 t o -20.3

-150.2

to

-153.9

nb

Yell mixed

-19.9 t o -22.9

-141.4 t o -167.8

ma

P o o r l y mixed

mb

Yell Mixed

Fig. 19. Local distribution of stable isotopes.

1

23