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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
<65 20-500
14 180
10 300
4 700
3 750--1,000
0.7 2,000
<39 100-500
soils
strongly desaturated ferrallitic
Bedrock
argillaceous sands upper Tertiary
chloritic schists 4
0.5 3,000
medium desaturated ferrallitic
tropical ferruginous
granites
granites
granites granites (pegmatite (quartz lodes) lodes) Precambrian
w
* PET = potential evapotranspiration. w
6
Organic carbon is determined by the oxygen which is consumed during oxidizing of organic matter with KMn04. Total nitrogen is titrated by the Kjeldhal method. The other used are the classical methods.
RESULTS
For several thousand samples the representation of eighteen parameters needs treatment and simplification of the data. A preliminary analysis of the results of the individual samples shows their relative homogeneity : all titrated waters, including seepage (drainage) and phreatic waters, have low solute concentrations, generally between 0.2 and 4 meq./l and high resistivities, between 5 * l o 3 and l o 5 C2 cm. Statistical treatment of the results has been executed for the results of two stations (Adiopodoume and Korhogo) : element-element bivariate relations, distribution of the parameter values depending upon the seasons and the volume of flow (runoff, seepage, etc.). Fig. 3 indicates the distribution of the resistivity values of the oblique drainage waters of Adiopodoume according to the seasons; the abscissa give the cumulated heights of rainfall since the beginning of the year: the first drainage waters, at the beginning of the main rainfall season, appear t o have lower resistivities -therefore t o be more mineralized -than the other waters; the chemical charge of the water is minimal at the middle of the rainfall season. This process of seasonal dilution occurs more or less marked at all sampling stations. We can see in Fig. 3 that low values of resistivity may exist during the whole rainfall season. However, the main factor is in fact not the seasonal influence but the volume of flow: small volumes and consequently low dilutions occur also at the middle of the rainfall season, when low-intensity rainfalls are occurring. On the other hand, the volumes of flow are always small and the resistivities show always minima at the end and particularly at the beginning of the rainfall season. Fig. 4 gives the distribution of Ca2+concentrations in oblique drainages waters of Adiopodoume; it shows clearly the influence of the volume of flow on concentrations: for high volumes of flows, characterizing for the middle of the rainfall season, the lowest concentration values always appear. This influence also exists, more or less, for the other major anions and cations; but it is unappreciable or nonexistent for the less mobile elements, such.as SiOz and the sesquioxydes of Al and Fe (Fig. 5). Generally, at each station, concentration values decrease when the volumes of flow are increasing. Consequently, and in view of determining the flux of dissolved products in waters, mean concentration values, balanced by flow volume (MPV: “moyenne pond6r6e par les volumes”), have been calculated: MPV
=
CVc/ZV
where V is the flow volume (1); and c is the concentration (mg/l).
7
630
540
1 450
1 1
1 1
1 1.
2
2
1
11 1 2
1 1
360
2
1
1
'.
1
1
* E
-
270
1
:
1 1 1 1
1
v) c
I1 1
a
1 1
180
1;
1
1
1 .. 1'
90.
2 11 11
z
I
1 I 2.2
1
11 2 2
1
1
11 1
1
?
1
11 I 1 1 1 1 :3 1 1 31 21 1'
11 ;
1
2
2
2 1 1 2
1 111 1 11 1
l! 1 1 1 1 1 1 12 1 1 1 z 12 271 71 1 1 1 ii 171 1 1 1 1 7 1 ? 112 1 3 1
2
1
11
1 3
2
1 21 1 1 11 1
1
li! 1
1
.................................................................... . I 1
0.0
1 11
z
1 1 11 I 11
i -z
1
w VI
Y
1 11
1 2
1 1
1
150.
0.00
750.
450.
300.
600.
1350
1050
900.
1200
19!
1650
1500
1800
MEASON
Fig. 3. Seasonal variations of resistivity in oblique drainage waters at the Adiopodoume station. The resistivity values, in cm, are plotted as a function of the cumulated heights (mm) of rainfall ( M M S E A S O N ) .
The MPV-values obtained for 18 parameters in the eight stations are presented in Table 11. The alkalinity values, which probably correspond with nearneutral pH-values of the HCO; species in these waters, have been calculated with reference to ionic balance; this constituent is prevailing, so the approximation is good. The MPV-values of the major ions expressed in mg/l are low, similar to those which have been obtained in springs and phreatic waters of the crystalline area, with a humid, temperate, or tropical, climate (Tardy, 1969; Bourrib, 1978). In general, concentration values increase from rainfall and throughfall waters to runoff &d drainage waters, down to 2 m depth, but they decrease
8 72
.................................................................... .
I
63 .1
54
+
1
.l .1 45
.
36
a
.
1
3
I
u
.1
A
a
u
.
111
.I 27
. 12 +
1
1 1
.31
.1
18
.3
1
1 1 .ll 1 1 .74 YZS 1 1 1
1
*32
.'t5
9.
1
1
1 1 1?2
1
1 2
1
*
1
*33 1 1 1 1 1 ? 1 .763311 I ,4451 2 4 ? l f > 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/Mg
l, 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 -
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.
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mCa2’0.4-
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u
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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.
-
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.
77
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
78
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-
79
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
-
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.
82
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.
83
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).
+
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
85
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).
86
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.
-
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.
87
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.
88
(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
+
+
90
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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95
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 (<25 g/kg total dissolved solids) lies beneath the shelf to a depth of 900 m. On land, brackish waters extend t o a maximum depth of -1.2 km below sea level in Lowndes County, Georgia. In deeper horizons, hypersaline brines ( > 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
98
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
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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
JEFFERSON I S
New Iberia
0
WEEKS I S
Frank Iin
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(Residual High) BAYOU SALE' SALT STRUCTURE
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Atchafalaya Bay
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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126
Fig. 16d presents log-log plots of chloride us. magnesium. They indicate that the mine brines are deficient in Mg relative to evaporating seawater. A similar trend of depletion is reflected by the Louisiana oil-field waters. Fig. 16e shows log-log plots of chloride us. strontium. They indicate that, like the Louisiana oil-field waters, the mine brines are enriched in Sr, relative to seawater.
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127
Fig. 16f has log-log plots of chloride us. boron. The Louisiana oil-field waters are usually high in B concentration, relative to evaporating seawater. Collins (1975) attributes this t o the survival of B in solution until late-stage crystallization. The mine brines are generally deficient in B, relative to evaporating seawater, which may be due to the dissolution of halite from the salt deposit. Some of the mine waters do plot near the upward tip of the standard curve, indicating a relatively high concentration of B. Fig. 16g presents log-log plots of chloride us. bromide. The mine-brine plots cluster near the upward end of the evaporite curve, suggesting the late stage of halite precipitation for these brines. Also, the plots appear as the upward extension of the trend of the oil-field waters.
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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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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.
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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:
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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
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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
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CHLORIDE
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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), <500 ppm. The low concentrations for piezometers B, and L, relative to B , and LB respectively, are due to dilution by rain, as shown by the concentrations of other cations which are insensitive t o microbial action (C1; Na', etc.).
146
In Table I11 sulfate concentrations are shown for some of the piezometers which were destroyed before sampling for isotope studies could be performed.
ISOTOPE DATA
The BaS04 precipitate has been processed following the classical methods in stable-isotope geochemistry (Mizutani, 1971; Sakai and Krouse, 1971; Filly et al., 1975) which will not be discussed in detail here. Data are given in per mil deviation from an international standard, which is the standard mean ocean water (SMOW) for '*O, and Canon Diablo Troilite (CDT) for 34S: (isotope ratio for sample) (isotope ratio for standard)
-
j
3. x 1000
-
Accuracy and reproducibility are ? 0.2yoOfor both' "0 and 34S. Variation in the heavy-isotope content of sulfate ion varies with the origin of sulfate: (a) Marine sulfate shows large variations through the geological epochs and even in the same basin it may vary in its 634Scontent (Claypool et al., 1980) and t o a lesser extent in l80(Holser et al., 1979). The present authors have data for the Bartonian gypsum beds (NW of Paris) (Fontes and Letolle, 1976) of the same age as the Mt. Mesly gypsum, for which they had the possibility to study one sample from a trench. This sample is shown in Fig. 3 together with the position of isotope sulfate values for all gypsum samples from the area (data taken from Fontes and Letolle, 1976). (b) Sulfate from fertilizers: this comes from the sulfuric acid used in the chemical treatment of phosphorites, and its isotope content is lower than for marine sulfate. The authors possess some data from a previous study of a nearby aquifer in the Brie and Beauce areas, south of Paris (Berger et al., 1977). (c) Sulfate from sulfide reoxidation: sulfides (or elemental sulfur), restored under oxic conditions, are oxidized and the sulfate formed usually shows a low isotope content. In the studied area, such a possibility occurs in the top soil, since waters become quickly anoxic with depth. The dissolved S2- content in the aquifer is highly variable and may be as high as 34 mg 1-' (piezometer 201), but the highest value observed in the lake is only 300 pgl-' , even at the deepest site (2). This implies that local reoxidation of hydrogen sulfide is not significant. (d) Sulfate from rains: this source cannot be discounted in this area since Megnien (1976) has shown that local rains contain from 7 to 1 6 mg 1-' SO:-, with an isotopic composition between -3 and 7%0. This source may contribute from 1%to 10% of the sulfate content of the aquifer, and may represent a potential component, a t least for the less mineralized waters.
+
147 I
20
0
10
6“s
I
I
20
30
I
%o/CDT
Fig. 4.6180-634S diagram for SO$- samples.
(e) Finally, water from the nearby rivers, of which the 634S fluctuates from 3 to t 8yoO(Mkgnien, 1976; unpublished data of the present authors, 1977), does not intervene locally, which is demonstrated by piezometry. I t is known that the lighter-isotope sulfate is more easily reduced by anaerobic bacteria than the sulfate of the heavier isotope. This effect has been studied by various authors (see, e.g., Thode et al., 1961, for 34S; Mizutani and Rafter, 1969,1973, for l80). I t has been shown that, on the whole, the sulfate remaining through reduction becomes enriched in l80as well as in 34S. The change may be described, in a 6180-634S diagram (Fig. 4), as going from left t o right and from bottom to top in Fig. 3; the slope being taken as first approximation between 2 and 4 (see Zak et al., 1980, for comments). Finally, let us mention that sulfate of a given isotopic composition may result from the mixing of initial batches with different isotopic composition. In a 6180-634S diagram, this appears as straight “mixing lines”. A t first glance, the isotopic data, as shown in Fig. 4,would favor a straightforward interpretation based on the reduction process from two sources: the first should be the gypsum bed from Mt. Mesly (or plaster imbedded in recent infillings), the second from an unknown source of sulfate with low l80 and 34S contents, which could be fertilizer sulfate or sulfate coming from the superficial oxidation of sulfur, always present in clinker. However, this interpretation does not fit with the distribution of sulfate concentrations (Table I) and a more elaborate model must be found. Fig. 4 shows that samples may be divided into four groups: (a) Group 1 comprises the lake samples 1 and 2 and piezometer samples A and 201, for which it is clear that sulfates come from Mt. Mesly.
+
148
(b) Group 1’ is for piezometer L alone. Sulfate comes from gypsum, but not necessary from the above sedimentary source, as this piezometer is located on a small mound built from refuses containing plaster indisputably manufactured years ago, from Parisian gypsum (“platre de Paris”) (roasting gypsum to make plaster does not modify the isotopic composition of the SO:- ion) (Fig. 3). (c) Group 2: piezometers C, D ,F and G show 6 l 8 0 - and 634S-valueswhich cannot be ascribed to a regional sedimentary origin, whatever the mechanism is involved t o modify the isotope content. Intrusion by Marne river water for piezometers C, F , D, and by the Seine river water for piezometer G, cannot be excluded on the basis of isotope data. Here the hypothesis of the intrusion of fertilizer sulfate, from the south may be introduced, as well as sulfate from sulfide reoxidation. Data from Berger et al. (1977) show 634Sof fertilizers with a mean value of +8%0. Waters from the Beauce aquifer (SW of Paris) exhibit a mean SO:- concentration of 40 ppm (+40),with a large scatter of 634S-values from -10 t o +12x0. Whatever the exact sources of this sulfate (fertilizer, rain, or pyrite reoxidation) may be, these values could be considered as representing approximatively the second end-member for the mixing process postulated for the waters sampled with the southeastern piezometers. (d) Group 3 ( B andE) could correspond to a mixing of waters from groups I and 2. Although the isotopic composition would indicate the same origin as for group 1 (dissolution of sedimentary gypsum, but without further
30
I-
n
20
0
u)
$ W
10
PARIS
0
500
GYPSUM
1000 b4=1
1500
2000
m9/e
Fig. 5. Interpretation of data through the 6?3---[
SO:-] representation.
149
evolution), the direction of the flow of the aquifer, as indicated in Fig. 1 from piezometric measurements, prevents us from such an interpretation. In defining group boundaries, both isotopic composition and concentration of sulfate were considered, so that they correlate to establish genetic relationships: group A group I group B groups 1' group C + group 2 --f
--f
+3
Moreover, the reduction phenomenon is clearly shown, since there is no possible local source with 634Shigher than 20yoO(see Fig. 4). In Fig. 5 the model built upon the above observations is summarized. l8O measurements fit exactly the same model.
+
CONCLUSION
Two end-members at least may be defined:
Source I : Interstitial water or a confined low-conductivity aquifer in which solutions develop to high sulfate concentrations. The sulfate which may have originated in Paris (Montmartre) gypsum has a heavier isotopic composition, being probably the sulfate remaining partly through reduction. These are found t o the west of Mt. Mesly, down the piezometric level. Source II: Waters of low salinity and isotopic light sulfate are found down the Bois de la Grange Plateau (sample F , group 3 ) . This light sulfate should originate from oxidation of sulfides or fertilizers. The piezometry forbids the possibility of sulfate originating from lateral infiltration from the two nearby rivers. The data of the isotope study have permitted the improvement of the hydrogeological model of the Creteil plain and ascribe more precisely the origin of water and of its mineral content. Moreover, it has been shown that a study based on isotope data alone can lead t o improper conclusions without the information concerning other parameters. ACKNOWLEDGEMENTS
Thanks are due to G. Shearer and I. Zak for their most helpful comments and criticisms, and to A. Dindeleux and M. Grably for technical assistance. This work was performed through a grant of C.N.R.S. (ERA 604).
150
REFERENCES Berger, G., Bosch, B., Desprez, N., Letolle, R., Marce, A., Mariotti, A. and Megnien, C., 1977. La mineralisation des eaux souterraines de la Beauce et le renouvellement de la nappe: application des marquages isotopiques (3H, "N, 34S). In: Colloq. Protection des eaux souterraines captees. Orleans. Bur. Rech. Geol. Min., Spec. Publ., 2: 21-34. Chesterikoff, A., 1980. Les eaux du lac de Creteil; leur origine; leurs caracteristiques. Actes 256me Congr. Assoc. Fr. Limnol., pp. 134-146. Chesterikoff, A. and Testard, P., 1981. Le lac de Creteil. Etude d'un plan d'eau artificiel en milieu en voie d'urbanisation. (In prep.) Claypool, G.E., Holser, W.T., Kaplan, I.R., Sakai, H. and Zak, I., 1980. The age curves of sulfur and oxygen isotopes in marine sulfate and their mutual interpretation. Chem. Geol., 28: 199-260. Filly, A., Letolle, R. and Pusset, M., 1975. L'analyse isotopique du soufre; aspects techniques. Analusis, 3( 4) : 197-200. Fontes, J.C. and Letolle, R., 1976. l80and %S in the Upper Bartonian gypsum deposits of the Paris Basin. Chem. Geol., 18: 285-296. Holser, W.T., Kaplan, I.R., Sakai, H. and Zak, I., 1979. Isotope geochemistry of oxygen in the sedimentary sulfate cycle. Chem. Geol., 25: 1-17. Megnien, C., 1976. Hydrogdologie du Bassin de Paris. Mem. Bur. Rech. GQol. Min., 98, 532 pp. Mizutani, Y., 1971. Improvement in the carbon reduction method for the oxygen isotopic analysis of sulphates. Geochem. J., 5 : 69-77. Mizutani, Y. and Rafter, T.A., 1969. Bacterial fractionation of oxygen isotopes in the reduction of sulphate and in the oxidation of sulphur. N.Z. J. Sci., 1 2 : 60-68. Mizutani, Y. and Rafter, T.A., 1973. Isotopic behaviour of sulphate oxygen in the bacterial reduction of sulphate. Geochem. J., 6 : 183-191. Sakai, H. and Krouse, H.R., 1971. Elimination of memory effects in 180/'60 determination in sulfates. Earth Planet. Sci. Lett., 11: 369-373. Thode, H.G., Monster, J. and Dunford, H.B., 1961. Sulphur isotope geochemistry. Geochim. Cosmochim. Acta, 25: 159-174. Zak, I., Sakai, H. and Kaplan, I.R., 1980. Factors controlling the l 8 0 / l 6 O and % S / j 2 S isotope ratios of ocean sulfates, evaporites and interstitial sulfates from modern deep sea sediments. In : Isotope Marine Chemistry, Uchida Rokakuho, Tokyo, pp. 339-373.
151
THE MADRID BASIN AQUIFER: PRELIMINARY ISOTOPIC RECONNAISSANCE
FERNANDO LOPEZ VERA', JUAN CARLOS LERMAN' and ANTHONY B. MULLER'
'
Department of Geology and Geochemistry, Universidad Autbnoma de Madrid, Madrid 34 (Spain) Laboratory of Isotope Geochemistry, Department of Geosciences, University of Arizona, Tucson, A Z 85721 (U.S.A.) (Accepted for publication May 11, 1981)
ABSTRACT L6pez Vera, F., Lerman, J.C. and Muller, A.B., 1981. The Madrid Basin aquifer: preliminary isotopic reconnaissance. In: W. Back and R. Lhtolle (Guest-Editors), Symposium of Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54: 151-166. The northwestern part of the Madrid Basin is underlain by an aquifer 6000 km2 in surface and having a thickness of 200-2000 m of unconsolidated silt and clayey sands, overlying a fractured basement complex which forms groundwater basins and barriers. The basin's fill materials are Miocene-Pliocene and were deposited in a continental environment of alluvial fans. Important urban areas such as Madrid, Toledo and Guadalajara are within the limits of this aquifer. The aquifer is characterized by a high degree of heterogeneity. Horizontal hydraulic conductivity is "0.5-1 m/day. The corresponding vertical conductivities are "100 times less. An attempt has been made t o corroborate a three-dimensional groundwater flow model proposed previously for this aquifer by using the environmental isotopes of carbon ( 14C) and oxygen ( lSO). Preliminary work, which the authors completed to date, has strongly supported the model but only in some areas. The homogeneity of the stable-isotope composition observed in the groundwater and its similarity t o the composition of recent precipitation indicate that the environmental and meteorological conditions during infiltration, even of the oldest waters, were not very different from those of the present.
INTRODUCTION
This paper deals with the aquifer of the Madrid Basin. The authors' scope is to update previously published, more preliminary studies on the basin (Lopez Vera, 1977a,b,c; Sastre, 1978; Herraez et al., 1979). The basin has an area of -6000 km2 and is formed mainly by detrital materials of Miocene age. The underlying fractured basement forms several sub-basins which depths range up t o 2000m. These sub-basins are separated by raised blocks which do not crop out. Madrid, Toledo and Guadalajara are the largest cities within the limits of the basin (Fig. 1).Although the principal water supply of these cities is surface water, groundwater is providing the industrial and
152
U .-
0
n
A=
I5
E
W
c
P c
v
a
e o
a$
6 U
..c 0
E
c
V c
153
agricultural users in this area. The smaller towns in the basin completely rely for their water supply on the aquifer being investigated. Several governmental organizations and individual researchers have studied the hydrogeology of the area (for a description see, among others, Llamas and Lopez Vera, 1975; Lopez Vera, 1977a, b, c). Despite these extensive studies there remain several unknown features of the hydrogeology of the Madrid aquifer which Llamas and Cruces (1976) have pointed out in their model, such as (a) the identification and physical separation of the different flow systems; (b) the residence time of the water in the aquifer. In the present paper, we make a preliminary attempt t o determine these features by means of studies based on the environmental isotopes 14Cand '$0. The aquifer consists of arkosic sand with variable amounts of silt and clay, deposited in alluvial fans during the Miocene and Pleistocene. The texture of the grains is finer and more homogeneous in the northern part of the basin than in the southern part. The basin is bound by the Sistema Central and Montes de Toledo mountain ranges on the northern and northwestern sides, and by gypsum and marly deposits t o the south. All these boundaries are thought to be impermeable. The aquifer is -200 km long and 30 km wide, the main rivers in the region are the Henares, Jarama, Manzanares, Guadarrama and Alberche. All of them are permanent tributaries of the Tajo River which runs along most of the southern boundary of the aquifer (Fig. 1).The basement complex is tilted towards the southwest. Sediments fill the basin t o a maximum thickness of 1800m at the Montes de El Pardo and decrease to 200 m to the south and southwest of Madrid (Pedraza, 1978). In this aquifer groundwater flows from the interfluvial areas where recharge takes place to discharge areas situated in valleys (Llamas and Lopez Vera, 1975; Lopez Vera 1977a, c). The flow system is controlled largely by the anisotropy and heterogeneity of the aquifer materials. Based on the results of a digital model constructed along vertical cross-sections (L6pezCamacho and Lopez Garcia, 1979), the anisotropy ratio of the aquifer would be above 100 while the thickness of its saturated zone does not have much influence on the flow system when the ratio of the horizontal to vertical dimensions is less than 30. When the thickness decreases, the intermediate flow tends to disappear, the regional flow tends to decrease and the local flow to increase. The heterogeneities of the aquifer medium are represented in the model by several layers which are less permeable than the rest of the medium. These zones significantly affect the flow pattern when the ratio of their permeability t o that of the bulk medium is less than 100. The horizontal permeability of the aquifer is -0.5m/day for the area between the Jarama and Henares rivers, on the eastern part of the basin, and 1m/day for the area between the Jarama and Alberche rivers, southwest of Madrid. The vertical permeabilities are 100 times less. According to Herraez et al. (1979) the aquifer is subject to two distinctive types of climate: (a) Humid or semi-humid areas with elevations of 800-900 m above mean
155
sea level. Most of the Sistema Central has this type of climate, with mean annual rainfall ranging from 500 t o 800 mm/y. and potential evapotranspiration within a range of 550-750 mm/yr. (b) Semi-arid areas at elevations between 400 and 900 m. A large part of the aquifer has this type of climate with a mean rainfall of 400-600 mm/yr. and potential evapotranspiration of 750-800 mm/yr. Most of the precipitation occurs during fall and winter with a mean temperature of 9°C. The summers are dry with a mean temperature of -2lOC. Most of the infiltration occurs during the cold season. From studies on the evolution of the soils, Vadour (1977) concluded that, in this region, the climate during the Pleistocene was 6-8°C colder than at present. This cooler period would probably have been more humid than the present-day climate.
SAMPLES: DESCRIPTION AND ANALYSIS
Samples for isotope dating and tracing were taken under several programs and analyzed in several laboratories. Beginning in 1974, F. Lopez Vera sampled the groundwaters of the Madrid Basin for I4C and tritium analyses and in 1976 A. Sastre sampled $or "0 analysis. These series consisted of 1 2 and 16 samples, respectively, taken from existing wells. Samples were analyzed in Madrid for 14C, tritium and l80. Our study is based on 102 samples: 74 of groundwater and 28 of surface water. For the geographical distribution of the samples, see Fig. 2. The results of the analyses are given in Tables I-VI. The 6 180-valuesare referred TABLE I l 8 0 analyses
of groundwater Depth (m)
6l80
(5)
Altitude (m) (6)
(7)
(8 )
1 0 1 0 0 0 0 0 1 0 0 1 0 0
64 9 654 680 619 602 600 615 618 740 6 10 590 715 637 621
0-23 0-130 38-185 26-108 117-162 57-276 0-3 85-1 05 230-250 80-1 62 75-128 0-64 3 7-1 43 20-1 23
-8.0 -8.5 -7.9 -8.6 -8.7 -9.2 -8.0 -8.1 -8.0 -7.7 -9.1 -8.9 -9.5 -9.5
Sample No. in Fig. 4
Sample No. in Herriez et al.
Laboratory*'
Aquifer*3
Aquifer zone*4
(11
(1979)*' (2)
(3)
(4)
1
1 1 1 1 1 1 1 1 1 1 1 2 2 2
1 2 3 4 5 6 7 8 9 10 11 12 13 14
0 3 1 4 0 0 0 5 2 6 7 8 0
11
4 3 4 1 1
1 3 3 3 3 2 2 2
(%o
1
156
TABLE I (continued) Sample No. in Fig. 4
Sample No. in Herriez et al.
Laboratory*2
Aquifer*3
Aquifer zone*4
(1)
(1979)*' (2)
(3)
(4)
(5)
15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52 53
10 9 0 12 0 0 0 0 0
13 15 0 14 0 0 16 0 0 0 0 0 0 17 0 0 0 24 23 0 18 21 0 22 0 0 0 19 0 20
2 2 3 3 3 2 2 1 1 2 2 1 2 1 2 2 2 2 2 2 2 2 2 2 2 2 3 3 4 2 2 2 2 2 2 2 2 2 2
2 2 2 2 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3
0
0 0 0 1 0 0 0 1 0 0 1 0 1 1 1 1 1 1 1 1 1 1 1 0 0 0
0 1 1 1 1 1 1 2 2 0 0 0
Altitude (m) (6)
Depth (m)
6l80
(7)
(8)
630 588 625 641 661 560 508 460 500 456 420 500 420 560 650 660 675 672 680 67 2 64 5 560 590 470 438 456 440 440 462 475 482 440 450 420 574 600 500 520 444
16-104 54-63 60-118 55-1 22 60-90 66-1 38 0-62 0-120 0-0 0-100 0-44 0-8 0-40 0-0 43-169 11-89 61-211 18 2-2 50 102-232 9 5-4 1 2 18-110 0-0 55-65 0-90 0-50 8-102 170-200 9 0-1 20 20-84 12-24 10-130 0-140 50-150 54-90 0-61 2 9-1 1 4 0-101 40-104 0-110
-9.4 -9.3 -8.8 -8.5 -7.7 -8.8 -8.1 -8.2 -7.7 -7.6 -7.2 -7.0 -7.3 -7.2 -7.9 -7.8 -7.5 -8.8 -8.1 -7.9 -8 .O -8.1 -7.6 -7.6 -8.2 -8.5 -8.2 -7.6 -7.9 -7.2 -7.2 -8.4 -9.1 -7.2 -7.2 -7.8 -8.9 -8.2 -8.8
(%o
1
*' 0 = not listed by Herriez et al. (1979); otherwise, sample number in that publication. *'
1 = Madrid; 2 = Paris; 3 = Tucson, Ariz.; 4 = average of two laboratories. 1 = between Manzanares and Jarama rivers; 2 = between Jarama and Henares rivers; 3 = Alberche river basin. *4 0 = discharge; 1 = recharge; 2 = other. *3
157
TABLE I1
“0 analyses of surface water - AITOYOS Sample No. in Fig. 4
Altitude (m)
6l80
1 2 4 5 6 7 9 10
1,040 850 800 670 620 700 690 810 465 560
-8.0 -7.6 -7.4 -7.6 -8.0 -7.7 -7.8 -7.6 -7.6 -7.5
I1 12
(Yo0 1
Sample No. in Fig. 4
Altitude (m)
6l80
13 14 15 16 17 18 19 20 21
780 780 450 440 600 440 420 750 400
-7.8 -7.9 -7.8 -7.9 -7.9 -7.6 -7.3 -7.7 -7.2
Sample No. in Fig. 4
Altitude (m)
6l80
6 7 8
550 530 490 500
-8.2 -8.7 -8.3 -7.5
(Yo0
1
TABLE 111
“0 analyses of surface water - Guadarrama River and its tributaries Sample No. in Fig. 4
Altitude (m)
6l80
1
660 620 630 560 580
-8.9 -8.7 -7.8 -8.6 -7.4
2 3 4 5*
(”bo
1
9*
(Ym1
* Tributaries. to the international water standard SMOW and the radiocarbon ages are given both in (a) the conventional scale, using the so-called “Libby half-life”; and (b) as “corrected ages”, assuming an original content of 14C equal to 75% of modern. The implications of this normalization are discussed in the corresponding section of this paper. The methods used here are classical in the application of environmental isotopes to the study of the aquifers’ characteristics, especially those like geographical recharge zones, groundwater flow velocity, etc. (see, e.g., Lerman, 1968a, b; Vogel et al. 1972,1975; and bibliography therein). The main purpose of this paper is t o update the information given in the above-mentioned, more preliminary papers and t o revise some of their conclusions on basin of both new and corrected isotopic data. The discussions deal respectively with l80and 14C in groundwater and in surface water.
l80 l80 was assayed both in surface waters and in groundwater to trace the geographical origin of the recharge. Since isotope analyses of seasonal and
158
TABLE IV l80in groundwater samples - summary
Samples from recharge or discharge zones
Median 6 l80
Between the Manzanares and Jarama rivers
both zones only recharge only discharge
Between the Jarama and Henares rivers Alberche river basin
Aquifer region
Midspread 6 "0 (%o )
Number of samples
-8.1 -8.0 -8.55
0.7 0.1 0.85
11 3 8
both zones only recharge only discharge
-9.3 -8.9 -9.35
0.7 0.7
7 1 6
both zones only recharge only discharge
-7.9 -7.75 -8.2
0.6 0.7 0.9
33 20 13
(Yo0
1
GROUNDWATER Zone 1 (between the Manzanares and Jarama rivers):
SURFACE WATER Zone 2 (between the Jarama and Henares rivers):
All samples -7 -7 -8 -8 -9 -9
79 0001 567 12
All samples -7 -7 -8 -8 -9 -9
Recharge zone -7 -7 -8 -8 -9 -9
9 00
7 01 567 12
589 34 55
-7 -7 -8 -8 -9 -9
Recharge zone -7 -7 -8 -8 -9 -9
Discharge zone -7 -7 -8 -8 -9 -9
Zone 3 (Alberche river basin):
9
0222223 56666778999 011122224 58889 1
-7 -7 -8 -8 -9 -9
234 56666677888999 00
4 58 23 6779
Recharge zone -7 -7 -8 -8 -9 -9
02222 566778999 0114 8 1
-7 -7 -8 -8 -9 -9
-7 -7 -8 -8 -9 -9
23 66 12222 5889
-7 -7 -8 -8 -9 -9
Discharge zone -7 -7 -8 -8 -9 -9
58 34 55
Fig 3. Stadistif diagrammes.
8 23 6779
159
TABLE V
‘*oin surface water samples - summary Source
Arroyos (several) Guadarrama River Guadarrama plus tributaries
Median 6 l 8 0
Midspread 6l8O
(%a
(%o
-7.7 -8.6 -8.3
0.3 0.4 0.9
Number of samples
19 7
9
annual precipitation were, and are not yet, available for the Madrid Basin region, a proxy type of samples was secured. This sampling was from a head of several “arroyos” (flowing streams) and from the Guadarrama River and two of its tributaries, which evaporation has been taken into account. These surface water samples provide us with an approximate idea of the isotopic composition of the possible recharge water. These isotope compositions are needed t o compare them with the &values of the groundwater. The wells sampled for this comparison were selected to represent several of the different zones of the Madrid aquifer, as described in the following sections of this paper. To facilitate the discussion of the data given in Table I, we extracted several groups of results, following the grouping suggested by Herraez et al. (1979). We then performed some exploratory data analysis [as designed by Tukey (1977) and implemented by Lerman (1977) and McNeil (1977)] on these groups of data. Table IV and Fig. 3 show the results of this exploratory analysis. The groups of samples consist of three aquifer regions which in their turn have been identified as recharge and discharge zones. The more preliminary publication by Herraez et al. (1979) was based on 24 groundwater samples while the present study is based on 53. The graphs shown in Fig. 3 are histogram-like representations of the data, known as “stemleaf displays” (Tukey, 1977; Lerman, 1977; McNeil, 1977) while the parameters given (Table IV) to represent the groups of data are the “median” and the “midspread”. These are considered, respectively, to provide a better graphic representation and a more robust set of parameters than are used traditionally, due to two principal reasons: (1)in our particular case, none of the distributions of the &values is Gaussian; and (2) the number of analyses is not large enough to warrant a “normal distribution approach” even if the distribution were “normal”. Examination of Tables IV and V and Fig. 3 show two main features: (1)While (a) the recharge water in the Alberche Basin has a &-valuesimilar to the 6-values of the arroyos and (b) the recharge water in the other two regions, between the Manzanares and Henares rivers, has &values which are similar to those of the Guadarrama river water which precipitates at a higher elevation catchment basin.
160
The assumed discharge aquifer water is -0.5?00 lighter than the assumed recharge water. Although this last might be explained by a hypothesized difference in 6-values between the older (discharge) and younger (recharge) water, an examination of the radiocarbon ages given in Table VI seems to indicate that there is no clear-cut correspondence between 6 180-values and groundwater age of recharge and discharge zones. It ought to be explicitly indicated here that nearly one-half of the isotope values listed by Herraez et al. (1979) were affected by a rather large systematic error (-1.56~00).All the samples analyzed at Tucson, Arizona, had been affected by a calibration error which has since then been corrected. This accounts for the new values listed in Table I and for the drastically different conclusions reported in the present paper, especially concerning the hypothesized difference in 6 l 8 0 related t o climatic changes (Herraez e t al., 1979), which is not valid on basis of the corrected data. Until a more complete survey of the isotope values of the surface water is done, covering sampling in different seasons and for several years, it does not seem warranted to advance a more detailed model about the recharge of the Madrid aquifer apart from indicating that a fair agreement exists between the 6-values of the different regions of the aquifer and the altitude of the basins where their water infiltrated. Thus the observed difference between the NW regions (samples 1 and 2 in column 4 of Table I) and the rest of the basin (sample 3 in column 4 of Table I) is fully accounted for in the present model. The correlation of 6-values with altitude reported by Herraez et al. (1979) is not any more apparent when the newly corrected results (Table 11) are plotted vs. altitude. The lack of the correlation between the altitudes of the collecting points and the 6-values of the water from the arroyos is probably due to the fact that each sampling point represents water collected in a catchment basin, and this is not necessarily in a linear relationship, with the elevation at the sampling point.
Thirteen additional radiocarbon analyses are now available to supplement the eight results reported by Herraez et al. (1979). In the present discussion we will use the “corrected” ages appearing in the last column of Table.VI which were computed based on the conventional radiocarbon dates (Table VI, column 11). The calculation of “conventional” dates assumes: (a) the initial radiocarbon content of the water was 100pmC. (percent modern carbon), i.e. an activity equal t o the U.S. National Bureau of Standards’ Activity; and (b) the half-life or radiocarbon equals the “Libby half-life” of 5568 yr. Since the actual initial radiocarbon activity of the infiltrating water was not 100 pmC, these “conventional” dates must be corrected. This initial
TABLE VI I4C groundwater dates
Sample No. in Fig. 4
Sample No. in Herraez e t al.
Aquifer*3
Aquifer zone’4
Altitude (m)
Depth ( 4
I4C modern (%)
0(I4C
story*'
(1)
(1979)*’ (2)
(3)
(4)
(5)
(6)
(7)
(8)
(9)
1
I
1
I 1 1 2 2
2 0 2 1 0 0 1 2 I 0 0 0 1 1 1 0 1
624 615 618 740 790 558 500 440 440 615 6 15 615 736 736 670 638 665 668 800 655 452
10-130 176-1 8 1 85-105 230-250 10-64 59-63 0-8 170-200 10--140 170-206 15-52 87-143 449-451 511-512 2-9 2-1 3 2-26 2-65 2-33 2-16 4-85
29.2 42.3 55.5 37.3 78.0 4.8 110.2 74.8 73.8 20.6 84.7 30.7 72.9 36.5 90.6 114.5 101.9 99.8 78.2 105.3 13.5
4.9 9.8 8.6 8.5 11.2 1.4 8.6
1 7 8
9 12 16 26 41 45 54 55 56 5 7a 57b 58 59 60 61 63
65 66
0 0 0 0 0 9 0 24 21 27 25 26 28 28 0 0
0 0 0 0 0
Labora-
I 1 I 3 1 1 1 1 1 I I 1 1 1 1 1 1 I I
3 3 3 1 I 1 1 1 2 3 1 1 2 1 3
1
1
1 0
*’ 0 = not listed by Herriez et al.; otherwise, sample number in that publication. *2
*3 *4
(%)
-
11.9 13.2 10.8 7.9 10.4 9.2 7.9 10.3 8.9 11.2 12.0 12.6 3.0
mcdern)
I4C conventional age (yr. B.P.) (10)
o(.‘C conven-
I4C
tional age) (yr. B.P.) (11)
age (yr. B.P.) (12)
“corrected age” (yr. B.P.) (13)
11,397 6,911 4,730 7,922 1,996 2,439 -780 2,332 2,440 12,691 1,334 9,486 2,539 8,096 793 -1,088 -151 16 1,975 -415 16,086
1,600 1,900 1,300 1,800 1,200 3,000 600
9,086 4,600 2,419 5,611 -315 22,081 3,091 21 129 10,380 -977 7,175 228 5,785 -1,518 -3,399 -2,462 -2,295 -336 -2,726 13,775
9,100 4,600 2,400 5,600 5,600 22,100 0 0 0 10,400 0 7,200 0 5,800 0 0 0 0 0 0 13,800
1 = Madrid; 2 = Paris; 3 = Tucson, Ariz. 1 = between Manzanares and Jarama rivers; 2 = between Jarama and Henares rivers; 3 = between Guadarrama and Alberche rivers 0 = discharge; 1 = recharge; 2 = other.
-
1,300 5,200 1,000 2,100 1,200 2.000 7 00 700 700 1,000 1,200 1,000 2,000
“75%”
162
activity depends primarily on the relative contributions of: (1)active C 0 2 from plant respiration in the soil zone; and (2) inactive carbon from the dissolution of mineral carbonates to the reservoir of bicarbonate ions in solution. This is discussed, among others, in more depth by Lerman (1972a, b) and Fontes and Garnier (1979). The proportions of these contributions may be estimated from 14C and 13C content of the mineral carbonate reservoir at present and of the soil-C02 reservoir at the time of infiltration. Estimates or measurements of these parameters are not yet available. As we lack this information, we chose the “average” value (valid for many aquifers) of 75% modern as original activity. The calculated “75% age” is listed in Table VI, column 22. Despite the uncertainties in the initial activities, and the consequence as for the knowledge of the absolute groundwater ages, the differences in ages between different wells as used for the velocity calculation should not be affected by the initial-value problems. It appears, from Table VI, that many samples are obviously “recent” because: (a) they show negative ages meaning that they contain 14C, generated as consequence of the nuclear bomb explosions; and (b) they are of age.zero or slightly positive meaning that they contain “modern” 14C. All these samples considered “recent” are listed with zero age in the last column of Table VI. The ages in this last column have also been rounded off. We will describe and discuss in the following sections some of the characteristics of the aquifer as they were studied with radiocarbon. We selected samples from the regions which were better analyzed for this discussion. We selected two regions, in particular: (1) between the basins of the Henares, Jarama and Manzanares rivers; and (2) the region between the Guadarrama and Alberche river basins. These regions differ in their geology, hydrogeology and hydrogeochemistry (Lopez Vera, 1977a). The sediments of both regions have different lithologic characteristics as well as different hydrogeologic characteristics such as their hydraulic conductivities (0.5 and 1m/day, respectively). We selected a few sections containing wells for which analyses can provide some information about the groundwater flow. Fig. 4 is a water-level map of the Madrid aquifer and Fig. 5 shows three sections, along directions which are parallel to the water flow as deduced from the equipotential lines in Fig. 4. Section 1-1’ shows the lines of flow between samples 63 and 16 (Table VI). From their corresponding radiocarbon ages, 0 and 22,000 yr., a groundwater m/day] is estimated. From velocity of -1.1 0.2 m/yr. [or (3.1 k 0.4) Darcy’s law, we calculate that for a hydraulic gradient of 9.2yoO(Fig. 4)and an effective porosity of 0.3 (such as usually determined for granular aquifers) the mean hydraulic conductivity between both wells is 0.10 f 0.03 m/day. A similar calculation was done for section 11-11’, incorporating samples 58 and 16 (Table VI). Using the groundwater ages, also of 0 and 22,000 yr., and (Fig. o 4), we obtain a radiocarbon velocity an hydraulic gradient of l l ~ o of 0.45 f 0.06 m/yr. or 1.2 f 0.2 m/day and calculate an average hydraulic
*
-
Boundary of aquifer 0
10
20
30
40
SOkm
-600-
Water-level contour. Shows approximate altitude of water-level spring,1975, in the upper part of the aquifer Contour interval 50 meters
............-................. Surface
water divide
Line of section
Fig. 4. Water-level map of detrital Madrid aquifer.
164 63
-
800
a
I
w
?
K
330
600
500 K W
K
400 m-
K
> w
K
In W
In
55 K
: o 2
W
2 K
700
56
m'
-
600
BOO
400
63
0 300 m.
&
0
5
10
15
20
Fig. 5. Cross-section (section line in Fig. 4).
25Km.
Sarnple/l4C "corrected" age (see Table PT) Water level
165
conductivity of -0.03 f 0.004m/day. Samples 8 and 9 in the same section must represent water from different flow lines as represented in Fig. 5. Section 111-111’ incorporates the samples 8 and 57b (Table VI) as well as 55, 56 and 57a. Fig. 5 shows the flow lines assumed in this case. For a gradient of 5%0 and a radiocarbon velocity of 1.2 f 1.4m/yr. (or 3.2 f 3.7m/day) we calculate an average hydraulic conductivity of 0.20 k 0.23 m/day. CONCLUSIONS
From the “0 analyses we learned that both “types” of water, recharge and discharge, as well as old and recent, have &values close enough to indicate that the climatic conditions in the Madrid Basin have not essentially changed during the time span represented by the water flow in the aquifer. From the 14C analyses we learned that the hydraulic conductivities calculated from the radiocarbon ages are of the correct order of magnitude as determined in local analyses done by aquifer tests. From both types of isotope analyses, we must conclude that there is no evidence for the large regional water flow that had been assumed until these analyses were performed and discussed. We must now accept that the radiocarbon age of the water in those wells whose analyses have been discussed in this communication can only be explained by the existence of local flow. Other factors possibly affecting the radiocarbon ages are being investigated in this aquifer, such as, e.g., carbon precipitation or dilution (Oeschger, 1974); and subsurface radiocarbon production (Zito et al., 1980). ACKNOWLEDGEMENTS
This research is part of a larger cooperative project undertaken by the Spanish Seccion de Investigacion de Recursos Hidraulicos, Consejo Superior de Investigaciones Cientificas, Madrid, and the Department of Hydrology and Water Resources of the University of Arizona, Tucson, Arizona, U.S.A. Principal investigators of this cooperative project are, respectively, M.R. Llamas and S.N. Davis. We acknowledge their kind support of our work. We wish to thank for their contribution to several aspects of this paper: S.N. Davis, C. Bajos Parada, J. Turner, A. Plata Bedmar, A Sastre; and for preparation of the manuscript, E. Zaval. This research was funded by grants t o the above-mentioned institutes. Thanks also are due to Laboratoire de Gbologie Dynamique, Universite Pierre et Marie Curie, Paris, where samples were analyzed for l80. REFERENCES Fontes, J.Ch. and Gapier, J.M., 1 9 7 9 . Determination of the initial 14C activity of the total dissolved carbon: a review of the existing model and a new approach. Water Resour. Res., 1 2 : 399-413.
166 Herraez, I., Lopez Vera, F., Plata Bedmar, A., Baonza del F’rado, E. and Llamas Madurga, M.R., 1979. Aplicacion del analisis del oxigeno-18 y carbono-14 a1 estudio del acuifero terciario detritico de Madrid, 4. Mem. I1 Simp. Nac. Hidrogeol., Pamplona, Publ. Assoc. Ge6l. Esp., Mus. Nac., Cienc. Nat., pp. 703-723. Lerman, J.C., 1968a. Agua subterranea en Bahia Blanca: investigacion con isotopos. Cienc. Invest ., 24 :315-319. Lerman, J.C., 1968b. Reconocimiento isotopico de acuiferos: Region de Cuyo. Cienc. Invest., 24: 368-373. Lerman, J.C., 1972a. Carbon-14 dating: origin and correction of isotope fractionation errors in terrestrial living matter. In: T.A. Rafter and T. Grant-Taylor (Editors), Proc. 8th Int. Conf. on Radiocarbon Dating. R. SOC.N.Z., Wellington, pp. H16-H28. Lerman, J.C., 1972b. Soil-C02 and groundwater: carbon isotope compositions. In: T.A. Rafter and T. Grant-Taylor (Editors), Proc. 8th Int. Conf. on Radiocarbon Dating. R. SOC.N.Z., Wellington, pp. D93-D105. Lerman, J.C., 1974. Discussions on stable isotopes in groundwater. In: Isotope Techniques in Groundwater Hydrology, 1974, Int. At. Energy Agency, Vienna, pp. 224-258. Lerman, J.C., 1977. Interactive exploratory data analysis: New workspaces in the UCC APL public library. Newsl., Univ. Arizona, Comput. Cent., l l ( 2 ) : 19-22. Llamas, M.R. and Cruces, J., 1976. Conceptual and digital model of the groundwater flow in the Tertiary basin of the Tajo River (Spain). I.A.H. Conf., Budapest, Int. Assoc. Hydrogeol., Mem., 11:186-202. Llamas, M.R. and Lopez Vera, F., 1975. Estudio sobre 10s recursos Hidraulicos subterraneos del Area Metropolitana de Madrid y su zona de influencia. Avance de las caracteristicas hidrogeol6gicas del Terciario del Jarama. Rev. Agua, 88 : 36-55. Lopez-Camacho, B. and Lopez Garcia, T., 1979. El flujo subterraneo en medios heterogdneos y anisotropos. Aplicaci6n de 10s modelos digitales a secciones verticales, Vol. 5. Mem. I1 Simp. Nac. Hidrogeol., Pamplona, Publ. Asoc. Ge6l. Esp., pp. 91-108. Lopez Vera, F., 1977a. Geoquimica de las aguas del “Terciario detritico” de la Fosa de Madrid, en relacion con el flujo subterraneo. Estud. Geol., 33: 525-534. Lopez Vera, F., 1977b. Estudios geologicos e hidrogeologicos sobre la Fosa media del Tajo. Bol. Geol. Min., 88(5): 401-416. Lopez Vera, F., 1977c. Hidrogeologia regional de la cuenca del rio Jarama, en 10s alrededores de Madrid. Mem. Inst. Geol. Min. Esp., 91: 227. McNeil, R.D., 1977. Interactive Data Analysis: A Primer. Wiley, New York, N.Y., 186 pp. Oeschger, H., 1974. Discussions on radiocarbon dating of groundwater. In: Isotope Techniques in Groundwater Hydrology, 1974. Int. At. Energy Agency, Vienna, pp. 000-000.
Pedraza, J., 1978. Estudio geomorfol6gico de la zona de enlace entre las sierras de Gredos y Guadarrama (Sistema Central espanol). Ph.D. Dissertation, Universidad Complutense, Madrid. Sastre, A., 1978. Hidrogeologia regional de la cuenca terciaria del rio Alberche. Ph.D. Dissertation, Universidad Complutense, Madrid. Tukey, J.W., 1977. Exploratory Data Analysis. Addison-Wesley, Reading, Mass., 688 pp. Vadour, J., 1977. La Region de Madrid. Alteraciones, Suelos y Paleosuelos. Editorial Ophrys, Paris. Vogel, J.C., Lerman, J.C., Mook, W.G. and Roberts, F.B., 1972. Natural isotopes in the groundwater of the Tulum Valley, San Juan, Argentina. Hydrol. Sci. Bull, 1 7 : 8 5 4 6 . Vogel, J.C., Lerman, J.C. and Mook, W.G., 1975. Natural isotopes in surface and groundwater from Argentina. Hydrol. Sci. Bull., 20: 203-221. Zito, R., Donahue, D.J., Davis, S.N., Bentley, H.W. and Fritz, P., 1980. Possiblesubsurface production of carbon-14. Geophys. Res. Lett., 7: 235-238.
167
RADIOCARBON DATING OF GROUNDWATER OF THE AqUIFER CONFINED IN THE LOWER TRIASSIC SANDSTONES OF THE LORRAINE REGION, FRANCE BERNARD BLAVOUX and PHILIPPE OLIVE Centre d e Recherches Giodynamiques, 7 4 2 0 3 Thonon-les-Bains (France)
(Accepted for publication November 24, 1980)
ABSTRACT Blavoux, B. and Olive, Ph., 1981. Radiocarbon dating of groundwater of the aquifer confined in the Lower Triassic sandstones of the Lorraine region, France. In: W. Back and R. Letolle (Guest-Editors), Symposium on Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54: 167-183. Stable-isotope and groundwater chemistry analyses on samples from about 40 boreholes in the Lorraine region (eastern France) are used to evaluate the paleohydrological regime of the Lower Triassic sandstone aquifer in the study of the mechanisms of the general water dynamics of the aquifer at present. In this sandstone aquifer, the estimation of the 14C activity, A14, in the total dissolved carbon (TDC) to 100% of modern carbon is justified by the absence of carbonates in the reservoir and by the 13C contents in the TDC. A significant relation could be determined between the 14C activity in the TDC and 13C: A14 = -7.7fj13 - 59.6%. This gradual enrichment of the TDC with 13Cresults from the escape of the dissolved COz . The meaning of the radiometric age is discussed as a function of piston and exponential-type flows. The study of the stable isotopes l80 and *H revealed two water families. The most recent waters have 6180-values higher by 1o/oo than the more than lo4 yr. old waters. This 1o/oo discontinuity corresponds to the improvement of weather conditions during the Holocene, which amounts to 2.5OC. The isochrons of the aquifer reveal the existence of hydrogeological sections bounded by faults where the rates of circulation are different.
HYDROGEOLOGICAL BACKGROUND AND PHASES OF THIS STUDY
In eastern France, the lower section of the Triassic is mainly represented by sandstones and conglomerates. These rocks crop out along the western border of the Vosges Mountains; their area extends towards the F.R.G. and Luxembourg where they disappear in a subsurface level south of Ardennes massif (Fig. 1). Being 500 m thick in the outcropping areas, they become gradually deeper lying and thinner towards the centre of the Paris Basin (see Fig. 2 for cross-section). These permeable formations are covered by Muschelkalk clays and marls (Middle Triassic). They usually overlie Permian formations consisting of argillaceous cemented sandstones and volcanic flows, north of a line running from Vittel to kpinal with the exception of the
168
Fig. 1. The extent of the Lower Triassic sandstones: lower Triassic sandstone outcrops (heavy-dotted area); probable maximum range of the Voltzia sandstone facies (broken line); and probable maximum range of the Vosges sandstone facies ( d o t t e d line) (after S.C.G.A.L., 1972).
Saar-Lorraine anticline crest where they lie over Carboniferous provinces. Near Vittel and Plombihres, the frequently arenated crystalline rocks occur just below the sandstones. This general structure, however, is affected by anticlinal and synclinal folds arranged in the direction of the Variscan orogeny (SW-NE) and is cut by SSE-NNW faults. From bottom to top, the Lower Triassic sandstone formation consists of the following:
+ + + + + + + + + + + + + + + +
+
t
+
+
+
+
+
+
+
+
+
+
+
M
170
(1)Vosges sandstones: 300-400 m thick pink sandstones, locally and irregularly feldspathic, iron cemented, with an uppermost conglomerate facies (max. 25 m thick main conglomerate). (2) Up to 6 0 m thick intermediate layers consisting of alternating coarse feldspathic sandstones and red or green clay beds and of compact dolomites. (3) Voltzia sandstones: 20 m thick fine-grained, micaceous and feldspathic sandstones, often containing pyrite. Most of the drillings have revealed Vosges sandstones below the main conglomerate. We may thus consider this trapped groundwater reservoir as a non-carbonate aquifer. The water in that aquifer is partly fresh water and salt water. The water-table piezometric map (Fig. 3) has been drawn from the permanent levels (draining rivers), the first levels recorded on bores drilled before the water table was disturbed by large-scale pumping, and from oil drillings at large distances from developed areas. It can be found that the water table was divided into two major sections: (1)A groundwater runoff concentration zone extending from the Vosges border to the coal trough (natural outlet and water pumped out of coal pits). (2) A dispersal zone with water streamlets flowing from the Cpinal-Vittel area towards Luxembourg and centre of Paris Basin. In view of taking part into a hydrodynamical study of this groundwater trapped below a confining bed, we have carried out an isotopic study of the tritium, radiocarbon, 13C, I 8 0 and deuterium contents in about forty deep boreholes (Fig. 4). In the field, we have recorded the temperature (fO.l"C), the pH (k0.05pH unit) and alkalinity (+1%). The total dissolved carbon (TDC) has been precipitated in situ. In the laboratory, we have analyzed the major elements dissolved in water and estimated the following values: 3H (expressed in TU), "0 (+O.ly,, vs. SMOW) and 2H (fo.5%, vs. SMOW). We have recorded the following contents of TDC: 13C (+0.2~,, vs. PDB) and I4C (percent modern carbon). The TDC is expressed in millimol per litre and, in the pH range of these waters (7.3-7.9) is equal to: TDC N mCOzaq
+ mHCOiaq
(1)
The main results are shown in Table I.
RADIOCARBON WATER DATING
If A , stands for the initial 14Cactivity of the TDC at the time when water was infiltrating and A , for the recorded activity, then the time t between these two moments (radiometric age) is equal to:
t (yr.)
= 80331nA,/A, (2) Since 1957 when K.O. Munnich suggested this sort of dating, many tech-
171
Fig. 3. Piezometric map before large-scale pumping and drilling began.
niques have been put forward in view of estimating the initial activity, in which usually, carbonate aquifers are considered. For the latest estimations, the reader is referred to Mook (1976), Wigley et al. (1978), Reardon and Fritz (1978), and Fontes and Gamier (1979).
172
d
NEU FCH AT E AU
k o r m e s
a Mlrecourt
,
~
IIII
~
P
0
10
20
~
~
30 40 50km
Fig. 4. Location of the studied deep boreholes.
Estimating the TDC initial activity A .
Since the present aquifer is not consisting of carbonate, the reactions in the infiltration zone can be written as follows:
+
+
COz H, 0 Me-silicate =+Me' 4- HCO; 4- H4Si04 where Me is a metallic ion.
+ ...
(3)
~
173
TABLE I Analytical results TDC (mmol/l)
Bining Rahling Basses Vigneules Achen Wittring Macheren Grosbliederstroff Sarreinsming Weldferding Frbcourt Faulquemont Holacourt Niedervisse Bouzonville M on dorf Arracourt Languimberg Moussey Puttigny Tomblaine Dieuze Champigneulles Langatte Heming Gblacourt Reding St. C16ment Morhange Rodalbe Ligneville Bulgn6ville Ville sur Illon ContrexbviIIe Damblain Charmes Mirecourt Norroy
Ageof groundwater
6"O
62H
6nTDC (o/oo vs. PDB)
AI4TDC (pmC)
7.40 6.91
-12.7 -11.8
68.3 44.4
3,100 6,500
3,700 10,100
-8.50 -8.65
-55.5
6.95 5.87 6.00 5.57
-13.0 -11.1 -11.9 -11.8
44.9 18.4 30.2 41.4
6,400 13,600 9,600 7,100
9,900 35,600 18,600 11,400
-8.40 -9.60 -8.80 -9.60
-54.4
5.62 5.55 4.18 4.15 3.14 3.40 6.75 4.33 3.09 4.10 1.84 2.17 2.87 4.62 2.19
-10.6 -10.7 -8.2 -8.4 -8.9 -9.6 -9.8 -9.6 -7.9 -10.9 -12.9 -11.5 -9.2 -8.9 -11.8
19.6 13.3 8.3 1.1 2.8 1.8 29.9 1.0 3.6 56.4 47.7 42.4 25.4 22.8 17.9
13,100 16,200 20,000 36,200 28,700 32,300 9,700 37,000 26,700 4,600 5,900 6,900 11,000 11,900 13,800
33,000 52,400 88,800 722,000 279,000 438,000 18,800 795,000 215,000 6,200 8,800 10,900 23,600 27,200 36,800
-9.65 -9.65 -10.3 -10.05 -10.20 -9.75 -8.60 -9.80 -9.00 -9.50 -10.00 -10.40 -9.90 -9.80 -9.90
3.9 2.82 2.54 3.49 3.19 4.86 2.95 2.45 4.76 4.12
-9.8 -7.9 -12.4 -9.2 -11.8 -9.7 -10.5 -10.5 -14.2 -12.1
16.9 15.1 11.7 11.0 10.2 6.9 15.2 2.4 60.5 43.0
14,300 15,200 17,200 17,700 18,300 21,500 15,100 4,000 6,800
39,500 45,200 60,600 65,000 70,700 108,000 44,800 327,000 5,200 10,600
-9.80 -10.00 -9.45 -9.80 -9.65 -9.80 -10.00 -10.00 -8.50 -8.70
3.11 3.98 1.91 4.44 3.38 4.50
-13.7 -13.6 -11.3 -9.7 -8.8 -8.9
34.8 33.4 23.6 1.7 1.1 <1
8,500 8,800 11,600 32,700 36,200 >37,000
15,000 16,000 26,000 465,000 722,000 795,000
-8.60 -8.70 -9.70 -9.70 -10.2 -9.50
piston flow
30,000
exponential
("/oo vs.
SMOW)
("/oo vs. SMOW)
-63.1
-69.0
-55.9 -63.9 -59.3 -62.6 -66.5 -70.0 -66.6 -64.9 -67.7 -64.1 -67.6 -65.2 -65.8
-55.8 -57.9 -57.5 -57.0 -65.5 -70.1 -64.7
174
The COz results from mineralization of the ground organic matter and is built up by partial pressures which are sometimes between ten and hundred times as high as that of atmospheric COz. Such biologically generated COz has the following isotopic characteristics (Troughton, 1972): 613CO2
N
-21 to -35Yo0,
the average value being -27Yo0
A ~ ~ C= O100% ~ As a matter of fact, tritium contents cannot be detected in either of the waters mentioned in Table I, so we can assume that these waters are not contaminated by thermonuclear 14C. This C 0 2 reservoir may be considered as infinite in relation to the dissolved bicarbonates. The isotope content of bicarbonate merely depends on the e I 3fractionation factor between bicarbonate and COz , say:
+
or 813HC03 2: tjI3CO2 9yoo (4) 613HCOi = 613CO2 + e l 3 with regards to the water temperature in this aquifer (Mook, 1976). Since the 6,,C02 ranges between -21 and -35Yo0 the 6,,HCO, consequently ranges between -12 and -26yoO, the most probable value being -18%0. Because the fractionation in 14C is about two times as much as in 13C (Salihge, 1979): €14 fi
2.3 x
€13
* 0.2
(5)
i.e. 2.3 x 9%o
€14
21%0
2.1%
and is therefore negligible as compared to the 100% of carbon initial activity. So we will assume: A0
fi
(6)
100%
Chemical and isotopic evolution of the TDC of the sandstone aquifer Once the TDC is chemically and isotopically balanced (eqs. 3 and 4) with the confining sandstone in the recharge area, it is generally admitted that the only evolution within the confined aquifer is the loss of 14C through radioactive decay (eq. 2). The structure is then said t o be closed. However, when plotting the 6 1 3 and A14 of the TDC in the 37 samples. (Fig. 5), a significant correlation ( r = 0.71) can be found: A:'
TDC = -7.7b13TDC
59.6%
(7) I t should be noted that with a 100% initial activity of 14C,the associated 6 13 is equal to -20.7%0, a figure which looks very much like those obtained by Troughton (1972) and Mook (1976), which confirms our preliminary hypothesis (eq. 6). It appears from eq. 7 that the TDC gains much 13C when time passes, -
175 I
I
*
I
I
I
I
I
60
c
E, LO
v
u 30 a
k-
u
%
20 10
0
- 1L
- 12
-0
-10
6,,TDC (X,vs.
PDB )
Fig. 5. 14C activity in the total dissolved carbon (TDC) as a function of the I3C content.
which may be accounted for by a loss of C 0 2 from TDC due to bacterial hydrogenation (Wigley et al., 1978). The outcome is a gradual enrichment of the TDC, since the escaping C 0 2 is isotopically lighter (eqs. 1and 4).Such a process can be compared with a Rayleigh distillation and may be written approximately in this form: aI3TDC- 6y3TDC N
~ 1 x3 In
(rnTDC/rn,TDC)
(8)
6 ,3TDC is the initial isotopic composition of the TDC containing rn, mmol/l, at the time when the structure becomes closed. € 1 3 is the fractionation between the evading C 0 2 and the TDC of the system, i.e. = -9yoo. Thus, if the original TDC is given the following characteristics are yielded: 6y3TDC = -13y0,
and
moTDC = 8mmol/l
and the evolution will take place according to the upward curve in Fig. 6:
(9) The loss of C 0 2 , which modifies the balances, may block the precipitation of carbonate. In such a case, the fractionation between the precipitating carbonate and the remaining TDC is low: N lYo0 ; the resulting TDC is almost unaffected. Taking the same initial conditions, the evolution will follow the downward curve (Fig. 6): 613TDC = 5.7-9lnTDC
+
hI3TDC = -15.1 4-1nTDC
(10) The same result can be obtained after gypsum dissolution (through the effect of cation in common with the carbonate), which sometimes occurs in the southern areas. It can be found, from Fig. 6 where eqs. 9 and 10 are plotted, that all samples but two are distributed within the two above-mentioned evolutions. The particular evolutions inside the defined system (loss of C 0 2 and/or precipitation of carbonate) depend on the respective kinetics of either re-
* I
-8
I
I
-10
6,,TDC
1
3Cc
I
1 . I
-12
-14
( % vs. . PDB )
Fig. 6. Evolution of the 613C-values as a function of the total dissolved carbon (TDC) content. The stars show the boreholes which have reached the upper sandstones.
action. The graph shows a qualitative difference between most of the boreholes with regard t o the bottom of Vosges sandstones where the evolution is primarily effected by the CO, escape, and the few drillings which involve the Lower Triassic sandstones as a whole, the evolution of which is essentially brought about by a precipitation of carbonates, probably due to the deposition of gypsum. In both cases, the carbon evasion, whether in the form of COP or of carbonate, does not modify the 14C activity of the TDC, except for the fractionation which is negligible with regard t o I4C (eq. 5). Thus, in the given instance of this non-carbonate sandstone aquifer, TDC activity recording enables dating through application of eqs. 2 and 6 : t (yr.) = 8033 In 100/A,
(11)
Meaning of the radiometric age We may consider that at present the flow system is stable in the whole system, i.e. the aquifer receives and supplies the same flow. The water volume corresponding t o the recent lowering of piezometric levels remains quite small as compared to the flow rates of intake and output.
177
TABLE I1 Water ages according to each flow model Activity (PmC)
90 80 70 60 50 40 30 20 10 5 1
Age of groundwater piston flow
exponential
846 1,792 2,865 4,103 5,568 = t 7,360 9,671 12,928 18,496 24,064 36,993
893 2,008 3,443 5,355 8,033 = h-' 12,049 18,743 32,132 72,296 152,626 795,260
According t o the characteristics of the flow 'within the aquifer, two models may be contemplated. In the former case, the age of water is proportional to the covered distance (the piston flow model). The radiometric age corresponds t o the actual mean age of water. In the latter where the various water streamlets are mixed as a result of different permeabilities [limit values recorded from test pumping: 0.3 10-5--8.1 m/s (S.C.G.A.L., 1972)], the flow is said to be exponential. Then the radiometric age is an apparent age (t,) related to the actual mean age (t,) by the following relation (Olive, 1970):
-
-
It can be seen from Table I1 that the actual ages vary to a large extent, depending on the adopted flow model. STUDY THROUGH THE STABLE ISOTOPES '*O AND 'H
The l80and 2H contents in all the analyzed waters have been plotted in a 6180--62H diagram (Fig. 7). All of them occur on the straight line of meteoric waters, with a slope of 8 and an initial value on the ordinate equal t o 12.5%0; that line will characterize the relative moisture of the oceanic areas from which the rain water which has recharged these water tables originates. Neither an effect of evaporation nor even any geothermal exchange with the reservoir could be detected. Two water families can be found. The larger group corresponds t o waters in which the 6180-values are lower than -9.5Yo0, whereas the 2H contents do not exceed -63yoO,and merely includes the waters of the saturated zone, confined in the Triassic lowermost sandstones. The other group, where the 6 l 8 0 - and 62H-valuesare respectively
+
178
Fig. 7. 6180-62H diagram of the saturated-zone waters. Age of piston -8.0
groundwater
5568
-44100 I
7400 I
I
12000
18700
I
exponent~ol-+5300
8033
(years) 9700
13000
18500 24000
32100
7 2 3 0 0 153000l
I
I
0.
0
Actual 80
I
I
I
I
I
I
I
I
70
60
50
40
30
20
10
0
'C
( pmc)
Fig. 8. Evolution of the 6180-values as a function of water ages.
higher than -8.7 and -57Yo0, includes some waters of the Vosges sandstone aquifer plus a number of recent waters from upper aquifers (occurrence of tritium). Fig. 8, where the F180-values are plotted vs. the age of water, shows that the oldest waters are those with the greatest depletion of stable isotopes. We
179
may therefore assume that they have infiltrated during a period with a colder climate than at present. Whatever the flow model being considered (piston or exponential), the infiltration took place during the Wurm glacial stage. The 6180-values in most recent waters are higher by lyooas compared to the previous group. A significant discontinuity separates these two groups. It occurs to be -lo4 yr. in the case of piston flow, or -2 *lo4yr. for the exponential-type flow. Taking into account the improvement of weather conditions, which has effected an enrichment of the heavy-isotope content in rainfall water and began during the Holocene, -lo4 yr. ago, we believe that the radiometric ages should rather be interpreted according to a piston model, without water mixing. Knowing the l 8 0 enrichment in rainfall water to be -0.4%0 per "C (Forstel et al., 1974), the climatic improvement may amount to 2.5OC.
HYDRODYNAMIC INTERPRETATION
The radiometric ages (piston model) have been used in the map of Fig. 9. Isochrons of 0.5 .lo4, lo4, 1.5 .lo4, 2 *lo4and 3 .lo4 yr. have been drawn. The confined aquifer may be divided into four sections bounded by faults where the rates of circulation are quite different. The southernmost area south of the Vittel-Rpinal fault contains recent waters aged less than lo4 yr. Very old waters appear just north of this fault, and we may assume that the latter acts as an impervious boundary. A well-distincted separation, in which the very old waters seem t o be blocked already in the outcropping areas, appears to be between this event and the River Meurthe fault which is traversing Nancy. A preferential circulation area, in which the waters are still fresh far below the outcrop, occurs NE of the River Meurthe. The existence of higher permeabilities may account for this anomaly. In addition, several SSE-NNW faults, reaching the basement and serving as a subsurface drain system, bring water up from the deep-seated horizons. Finally, we see a body of older waters still existing between the eastern sandstone outcrops and the Saar coal field. We have plotted these radiometric ages (piston flow hypothesis) in Fig. 10 as a function of the distance between the borehole and the outcrop according to the flow lines defined in Fig. 3. Using the field-recorded mean parameters (Dague, 1969), i.e. a permeability of lo-' m/s, a slope of 1/1000 and an effective porosity of 16/100 in the outcrop, a mean actual rate of 1.9m/yr. can be calculated. This rate, deduced from the groundwater hydraulic conditions, has been entered into the age-distance diagram. Most of the data points indicate that the radiometric rate is lower than the hydraulic rate. This fact may first be accounted for by the slight variations in permeability which, according t o test pumpings, range from 0.3 *lo-' t o 8.1*10-' m/s. A second explanation may con-
180
Radiometric age in 103 y e a r s
>30 lo3 y e a r s 30
>20 lo3 y e a r s
lo3>
20103>=
>15.103years
-Fault
0 ,
10 20
30 40 50km
I
l
l
"
Fig. 9. Isochrons of the aquifer confined in Triassic sandstones.
sist in the variation of the hydraulic gradient, not only in space, but also in time. The 100-m uprise of the general basic level from 18,00Oyr., which has very much reduced the hydraulic gradients, has also gradually blocked the confined system flow. The currently measured hydraulic gradients which have been taken into account in the calculation of the hydraulic head rates (1.9 m/yr.) result from
181
7.LOO
\
\ \ \
\ \ 1
1
,
2
,
4
*
6
*
8
1
1
1
I
37000
1
0
20
30
LO
50
Distance from recharge a r e a ( k m )
Fig. 10. Comparison between the 14C isotopic exchange rates and the hydraulic head rates.
the recent development of this aquifer. They are consequently much higher than those of the Holocene during which the confined aquifer was saturated. We may then question the hypothesis of a permanent flow regime on a time scale where the climate (effective rainfall) and hydraulic (basic level) changes should not be ignored. Conversely, during glacier invasions, with lower basic levels, i.e. higher gradients, the water replacement in the aquifer is favoured. Such is, perhaps, the explanation of the rather homogeneous 6"O-values in these old waters, as the aquifer water is preferentially replaced during the cold epochs.
SUMMARY
The following analyses have been carried out on samples from about 40 boreholes in the Vosges (eastern France) Triassic sandstones: the chemistry of the major ions dissolved and the stable isotopes l8O, 2H, 3H, 13C and I4C, in order t o study the water flow in the confined saturated zone of the aquifer. In this sandstone aquifer, the estimation of the I4C activity in the total dissolved carbon (TDC) t o 100% of modern carbon is justified by the absence of carbonates in the reservoir and by the 13C contents in the TDC. A significant relation could be determined between the TDC l4 C and l 3 C: AI4 = -7.7613 - 59.6%
This gradual enrichment of the TDC with I3C results from the escape of the dissolved C 0 2 , a process similar t o a Rayleigh distillation.
182
The meaning of the radiometric age is discussed as a function of piston and exponential-type flows. The study of the stable isotopes ( l8 0 and 2H) enables us to distinguish two water families. The most recent waters have 6180-valueshigher by lyoocompared with the more than lo4 yr. old waters. That l % o discontinuity corresponds to the improvement of weather conditions during the Holocene, which amounts to 2.5”C. The isochrons of the aquifer reveal the existence of hydrogeological sections bounded by faults where the rates of circulation are different. As the radiometric rates are lower than the rates calculated by means of the current hydraulic parameters, the hypothesis of the permanent flow regime becomes questionable. We can distinguish an epoch with an intensive replacement of water (very low basic level on a glacial maximum, resulting in high gradients and rates) from an epoch during which the confined aquifer absorbs almost no water (high basic level due t o the rise of temperature, and very low hydraulic gradients). Such a paleohydrological reconstruction is quite necessary before studying the mechanisms of the general water dynamics in the major confined aquifers.
ACKNOWLEDGEMENTS
The preparation of this study was undertaken with the technical and financial assistance of the Agence de Bassin Rhin-Meuse and the Service Regional d’Am6nagement des Eaux de Lorraine. Special acknowledgement is made by the authors t o Mrs. Ramon (Agence) and Salado (S.R.A.E.).
REFERENCES Dague, Ph., 1969. Etude gdologique et hydrogeologique de la nappe des gr&sdu Trias inferieur dans 1’Est de la France. University of Nancy, Nancy, 222 pp. Dassibat, C., Ramon, S. and Zumstein, J.F., 1975. Carte hydrogeologique du bassin Rhin -Meuse a 1/500.000e. Mission Dkleguee du Bassin Rhin-Meuse, Metz. Fontes, J.C. and Garnier, J.M., 1979. Determination of the initial 14C activity of the total dissolved carbon: a review of existing models and a new approach. Water Resour. Res., 15(2): 399-413. Forstel, H., Putral, A., Schleser, G. and Lieth, H., 1974. The world pattern of oxygen-18 in rainwater and its importance in understanding the biogeochemical oxygen cycle. Isotope ratios as pollutant source and behaviour indicators. Proc. Symp. Int. At. Energy Agency, Vienna, pp. 3-20. Mook, W.M., 1976. The dissolutionexchange model for dating groundwater with ‘4C. Interpretation of environmental isotope and hydrochemical data in groundwater hydrology. Int. At. Energy Agency,Vienna, pp. 213-225. Olive, Ph., 1970. Contribution a I’ktude geodynamique du cycle de I’eau dans I’hemisphbre nord par la methode du tritium. Thesis, University of Paris VI, Paris, 138 pp. Reardon, E.J. and Fritz, P., 1978. Computer modelling of groundwater 14C. J. Hydrol., 36: 201-224.
183 Saliige, J.F., 1979. Determination experimentale du fractionnement isotopique l4 C/13C au cours de processus naturels. Mkm. Conserv. Natl. Arts Mktiers, Paris, 76 pp. S.C.G.A.L. (Service de la Carte Gkologique d’Alsace Lorraine), 1972. fitude hydrogkologique de la nappe aquifhre des grhs infratriasiques dans le Nord-Est de la France. Serv. Carte Gkol. Alsace Lorraine, Strasbourg, 63 pp. Troughton, J.H., 1972. Carbon isotope fractionation by plants. Proc. 8th Int. Conf. on Radiocarbon Dating, Lower Hutt, pp. 421-438. Wigley, T.M.L., Plummer, L.N. and Pearson, F.J., 1978. Mass transfer and carbon isotope evolution in natural water systems. Geochim. Cosmochim. Acta, 42: 1117-1139.
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185
URANIUM ISOTOPES AND 226RaCONTENT IN THE DEEP GROUNDWATERS OF THE TRI-STATE REGION, U.S.A. J.B. COWART Department o f Geology, Florida State University, Tallahassee, FL 32306 (U.S.A.) (Accepted for publication February 27, 1981)
ABSTRACT Cowart, J.B., 1981. Uranium isotopes and 226Racontent in the deep groundwaters of the Tri-State region, U.S.A. In: W. Back and R. Lhtolle (Guest-Editors), Symposium on Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54: 185-193. Uranium isotopes have been analyzed in a number of water samples from the groundwater flow system of the Cambro-Ordovician aquifers in the Tri-State region of Missouri, Kansas and Oklahoma, U.S.A. The system consists of sodium chloride water having low uranium content (“0.04 pgfl) which meets westward flowing calcium bicarbonate water having uranium content about an order of magnitude greater in a transitional zone located just to the west of the Missouri-Kansas border. In the area where the waters mix and then flow southwestward, H2S is commonly found in the water, 226Rais relatively high and uranium concentration is lower than to the east. It appears that a t least part of the dissolved uranium in the westward flowing water is precipitated in the zone of mixing. Because the dissolved uranium has 2wU/238Uactivity ratio of 7-10, any precipitated uranium would have an enhanced capability of generating the daughter 226Ra.However, the distribution of uranium isotopes within the system suggests that the source of radium in the water of the transitional zone is not the inferred present-day zone of uranium accumulation but rather the brines themselves.
INTRODUCTION
In the Tri-State region of Missouri, Kansas and Oklahoma, measurements of public supply wells have shown that the U S . Environmental Protection Agency standard for radium concentration ( 5 pCi/l) is exceeded in a number of locations. A part of the investigation into the cause of this health hazard is the study of the uranium isotopes, 238Uand 234U,precursors of 226Ra. The longest lived nuclides of the uranium decay series, which are located at the early part of the series, are shown in Fig. 1.In a closed system, given sufficient time, the decay series will reach a state of secular equilibrium wherein the radioactivity (or “activity”) of each member will be the same, although the actual number of atoms present may vary greatly. Activity for a given nuclide is defined by:
186 2 3 8 ~
2 3 4 ~
t l / z =4.49x109yr
419Mev
234~a
4.72
t '12 = I,l8rn
230Th f
'/2=24.1 d
t
j/z = 7.5 x 104yr
226Ra t Y2 = 1622 yr
Fig. 1.The uranium decay series from 238Uto 226Ra.
where A = activity; X = the decay constant; N = number of atoms; and t I l 2 = half-life. At equilibrium, the activity ratio between any two members of a decay series is 1.00. However, at and near the Earth's surface, disequilibrium of the various nuclides of the uranium series is found to occur. The disequilibrium is especially pronounced in groundwaters (Cherdyntsev, 1971; Osmond and Cowart, 1976). The fractionation of the nuclides can occur as a result of chemical differences between elements, the fractionation of isotopes of a given element may occur because of preferential leaching of one (because of its radiogenic origin), or by the direct action of recoil during radioactive decay (Osmond and Cowart, 1976). Uranium, thorium and radium tend to fractionate because of chemical differences. Uranium tends to be mobile in oxidizing waters containing complexing bicarbonate, sulfate or phosphate anions; in reducing environments the solubility of uranium is sharply decreased and precipitation occurs. Thorium is virtually immobile under almost all surface conditions. Radium is quite mobile in high-Cl waters but in the presence of sulfate, radium is very insoluble. In addition to these elemental fractionations, separation between the two long lived uranium isotopes, 234U and 238U, occurs commonly in groundwaters (Cherdyntsev, 1971). Thus, although the ultimate parent of radium is 238U, the local source for a radium anomaly may be comprised mostly of 230Th with little uranium, or 234U and in-grown 230Th with little 238U activity, or 230Th, 234U and 238U in approximate equilibrium. The mix of these nuclides is a function of the history of the geochemical barrier which causes uranium precipitation.
GEOLOGIC AND HYDROLOGIC SETTING
In the Tri-State region, the stratigraphic section above the Precambrian basement, consists of Cambrian and Ordovician dolomites and sands, and generally less permeable Devonian, Mississippian, Pennsylvanian and Pleistocene sedimentary rocks. The Cambro-Ordovician section is the source of the
187 TABLE I Stratigraphic section for southwestern Missouri and southeastern Kansas* System
Series
Quaternary
Recen t-Pleistocene
Thickness (m)
Water-bearing character
0-1 0
small yields
0-1 00
small yields
UNCONFORMITY Pennsylvanian
Desmoinesian UNCONFORMITY
Mississippian
Chesteran Meramecian Osagean Kinderhookian
0-3 0 25-50 30-100 0-1 5
aquitard ( ? ) small yields shallow aquifer aquitard (?)
UNCONFORMITY Mississippian- Devonian
0-25
aquitard ( ? )
0-275 0-1 65
deep aquifer deep aquifer
UNCONFORMITY Ordovician Cambrian
Lower Upper UNCONFORMITY
Precambrian
aquiclude
*From Macfarlane (1980).
water produced by most public supply wells in the area, with the Cotter Dolomite and the sands of the Roubidoux and Gasconade formations being the most prolific sources (Macfarlane, 1980). Mississippian formations produce small amounts of water for domestic wells. Table I gives some general information on the aquifers and aquitards of the region. The major source of recharge of the Cambro-Ordovician aquifer is the area around the Ozark Mountains of southwestern Missouri. The generalized potentiometric surface for the Cambro-Ordovician aquifer is shown in Fig. 2. Since the productive zones are both sandstones and carbonates, the permeability is a function of both the primary permeability of the sandstones and the secondary permeability developed in the carbonates during times of uplift and subaerial erosion (Chenowith, 1968; Macfarlane, 1980). Freshwater moving westward in the Cambro-Ordovician section encounters saline, H2 S-bearing waters near the Kansas-Missouri boundary. The transitional zone between the two water masses serves as the source for various public supply wells in eastern Kansas. West of this zone, water too saline for general public use is encountered. Because of the paucity of wells in the more saline area, the configuration of the potentiometric surface presently is not well defined, although there is some evidence for a very gentle southward to southwestward plunging
188
POTENTIOMETRIC
SURFACE
CONTOUR INTERVAL = 25 m
Fig. 2. The potentiometric surface of the Cambro-Ordovician aquifer in the Tri-State region with sample locations. (Potentiometric data provided by the Kansas Geological Survey.)
potentiometric trough, west of the transitional zone (P.A. Macfarlane, pers. commun., 1980). METHODS
Samples of -17.51 in volume were collected, stabilized with nitric acid and shipped to the laboratory for further processing. The samples were volumetrically measured, acidified t o a pH of > 1.0, had an Fe carrier added and were spiked with a known quantity of 232U,an artificial isotope useful as a yield tracer. Each sample was then heated to boiling and treated with NH,OH until iron hydroxide precipitation occurred. The supernatant liquid was decanted and the precipitate was dissolved in HC1 and extracted with isopropyl ether to separate most of the Fe from the scavenged elements. Ion-exchange procedures were utilized t o separate uranium from interfering elements. The uranium was electroplated on a stainless steel planchet which was then used as a thin source for alpha spectrometry.
189
Details of procedures may be found in Osmond and Cowart (1976) and in references cited therein. RESULTS
The fresh waters of the eastern side of the area of investigation have the highest uranium concentrations and the lowest radium concentrations (Table 11). Along the eastern boundary of the transitional zone, where the fresh TABLE I1 Uranium isotope and radium analyses Site No.
1 2 3 4 5 5A 6 7
8 9 10 11
12 13 14 15 16 17 18 19 20 21 22
*' *2
Location
Uranium concentration*'
234~/238~
activity ratio*'
Marshfield, Mo. Republic, Mo. Mt. Vernon, M o Sarcoxie, Mo. Duenweg, Mo. Duenweg, Mo. Joplin, Mo., Farmers' Coop. Baxter Spgs., Kans., No. 1 Baxter Spgs., Kans., No. 5 Baxter Spgs., Kans., No. 6 Webb City, Mo. Carl Junction, Mo. Cherokee, Kans. Asbury, Mo., Scammon, Kans. West Mineral, Kans. McCune, Kans. Pittsburg, Kans. Crawford Co., Kans., No. 4 Frontenac, Kans. Crawford Co., Kans., No. 7 Girard, Kans. Walnut, Kans.
226Ra concentration*2 (pCi/l)
(Yg/l)
(pCi/l)
0.214 f 0.019 0.417 f 0.035 0.384 f 0.030 0.138 f 0.011 0.416 f 0.027 0.443 f 0.027
0.07 0.14 0.13 0.05 0.14 0.15
6.54 f 0.51 7.76 f 0.51 10.55 f 0.66 10.00 f 0.77 10.37 f 0.54 10.51 f 0.58
0.5 0.6 0.3 1.2 -
0.266 f 0.018
0.09
9.47 f 0.53
1.3
0.441
0.030
0.15
7.56 f 0.40
1.8
0.277 f 0.023
0.09
8.32
0.57
2.3
0.349 f 0.028 0.325 f 0.041
0.12 0.11
7.07 f 0.46 10.60 f 1.04
1.2 1.o
0.335 f 0.026
0.11
8.72
0.55
1.5
0.542 f 0.042 0.584 f 0.042 0.078 f 0.009
0.18 0.20 0.03
6.76 f 0.39 7.24 f 0.34 9.86 f 1.35
1.5 1.5 1.9
0.014 f 0.002 0.022 0.003 0.037 f 0.005
*
0.005 0.007 0.012
5.42 f 0.89 8.73 f 1.29 9.02 f 1.12
6.4 7.5 1.9
0.041 f 0.005 0.033 f 0.005
0.014 0.011
8.84 f 1.15 9.44 f 1.32
1.8 3.1
0.053 f 0.007 0.033 f 0.003 0.016 f 0.002
0.018 0.011 0.005
9.10 f 1.13 9.65 f 0.89 8.80 f 1.42
7.5 8.7 7.6
f
f
f
Confidence limits are 2a-values based solely on counting statistics. 226Radata provided by the Kansas Geological Survey.
-
190
waters first encounter the saline H, S-bearing waters, the uranium concentration decreases abruptly. In the same general location, the radium concentration begins a progressive increase in a westward direction as does the specific electrical conductivity (Fig. 3). In the transitional zone the uranium concentration is at least an order magnitude less than in the freshwater zone whereas the radium is about an order of magnitude greater. The 234U/238Ualpha activity ratio (A.R.) varies relatively little across the area; both fresh and saline waters have a high value of -10 and a low value in the range of 5.5-6.5.
- 1200
PROFILE A Hz /
- 1000
--
I
-800 z o Y
I
@
SPECIFIC CONDUCTANCE ____e-
~
- 600
-400
/”-
-200
lo
,
9,8,7
6
5
3
4
2
I
\ H
21
I 19
17
89 2-
6 W
71
I 1
E2 E
s
”
u v) N
W
y
13
II
10
5
%
191
g r
41 3
7
.
16
15
14
I2
I II
Fig. 3. Three sections showing uranium, radium, specific conductivity and H2S data across the region of study. The locations of samples is shown in Fig. 2. Profile A does not extend into the transitional zone. The values shown for samples 7, 8 and 9 are somewhat anomalous for the area and seem t o be from isolated zones of possibly relict water. Profiles B and C show a large decrease in uranium concentration associated with the edge of the transition zone and the sympathetic variation of radium with specific electrical conductance (in pS/cm).
DISCUSSION
The decrease of uranium concentration, shown in Fig. 3b between locations 13 and 1 7 and in Fig. 3c between locations 12 and 14, is similar to that observed at other locations where oxidizing waters encounter a decrease in redox potential but the lack of an associated well-defined increase in A.R. is unusual. An increase in A.R. at a redox front has been observed in both sandstones (Cowart and Osmond, 1980) and in carbonates (Cowart, 1980). The decrease in dissolved uranium serves to locate uranium presently accumulating as a coating on the aquifer rocks. By recoil or selective leaching modes, the A.R. in water can be increased. Direct recoil input is a function of the 238Upresent whereas selective leaching is strongly influenced by the relative amount of 234Upresent. The amount of 226Ragenerated by such a coating is a function of the amount of 234U(and 230Th)present but since its geochemistry differs from uranium, the radium is not constrained to remain with the accumulation. Thus, the high-A.R. uranium precipitating at the eastern edge of the transition zone should be an efficient 226Ragenerator, assuming enough time for 230Thin-growth has lapsed. Fig. 3b and c shows that: (1)only a relatively small, poorly defined increase in A.R. occurs downgradient of the inferred zone of uranium accumulation;
192
and (2) 226Radoes not increase markedly at the inferred accumulation but rather at some distance downgradient. This information seems to preclude the present-day precipitation site as the major source of high radium values, so we are left with several alternatives: (1)that the boundary between the freshwater zone and the transitional zone (hence the site of uranium accumulation) was once further west and that it has migrated eastward in the recent geologic past, or (2) that the source of radium is the high-Cl brine. The first alternative is attractive in that it can account for both the location of the high radium values and the lack of 234Uanomaly in the water. However, there is neither independent evidence that such an upgradient migration has occurred nor any apparent reason why it should have moved such a long distance so rapidly. This leaves us with the brine as the major radium source. This source must remain speculative for now as there are no appropriate samples presently available from this area although high radioactivity (presumably radium) has been reported from oil-field wells further west in Kansas (Gott and Hill, 1953) and in Oklahoma (Bloch, 1980). U-isotope analyses of brines from the Gulf of Mexico coast region indicate that the A.R. range is usually 1-2 (J.B. Cowart, unpublished data, 1976; Kraemer, 1981). Brine from one well in Israel is reported to have A.R.’s in the range 10-12 (Kronfeld et al., 1975). If the lower A.R.-values are valid for Kansas brines, then the U isotopes measured as far west as locations 16 and 22 must have originated in the freshwater region and not with the high-Ra waters. Thus, in a mixing zone, various nuclides may have different sources even though the nuclides are members of the same decay series; the deciphering of such multisource systems may be best approached by isotopic rather than elemental analyses.
ACKNOWLEDGEMENTS
The author wishes to express his appreciation t o the following people for their contributions t o this study: Allen Macfarlane of the Kansas Geological Survey for providing not only the samples but also the original impetus for this study, Lucy San Juan for technical assistance in the laboratory, Rosemarie Raymond, Tom Fellers and Dennis Cassidy for drafting and photographic expertise, Lynne Gaskin for typing the manuscript, and J.K. Osmond for useful criticism of the manuscript.
APPENDIX
Samples of water from three oil fields in Kansas, located from 32 to 40 km west of Baxter Springs, Kansas (sites 7-9) have recently been analyzed by the author for uranium isotopes. The waters were produced from
193
the Cambro-Ordovician formations in the section. The presently available data for these wells are as follows:
Location
Chetopa oil field Schreppel oil field Wackerle oil field
u/238u
U concentration
234
(CLg/l)
activity ratio
0.010 f 0.002 0.010 f 0.002 0.013 f 0.003
1.68 0.39 2.08 f 0.58 1.46 0.42
* *
Specific electrical conductance (CLS/cm) 2,150 3,550 12,400
These U-isotope values are similar to those found in oil-field brines from other parts of the U.S.A. but which previously could only be assumed present in the Kansas area.
REFERENCES Bloch, S., 1980. Origin of oil field brines - a hypothesis. Okla. Geol. Notes, 39: 177-182. Chenowith, P.A., 1968. Early Paleozoic (Arbuckle) overlap, southern mid-continent, United States. Am. Assoc. Pet. Geol. Bull., 52: 1670-1688. Cherdyntsev, V.V., 1971. Uranium-234. Israel Program for Scientific Translations, Jerusalem, 234 pp. (translation of: Uran-234, Atomizdat, Moscow, 1969). Cowart, J.B., 1980. The relationship of uranium isotopes t o oxidation/reduction in the Edwards carbonate aquifer of Texas. Earth Planet. Sci. Lett., 48: 277-283. Cowart, J.B. and Osmond, J.K., 1980. Uranium isotopes in groundwater as a prospecting technique. U.S. Dep. Energy, Rep. GJBX-119,112 pp. Gott, G.B. and Hill, J.W., 1953. Radioactivity in some oil fields in southeastern Kansas. U.S. Geol. Surv., Bull., 988-E: 69- 122. Kraemer, T.F., 1981. z34Uand 238Ucontent of brine from geopressured aquifers of the northern Gulf of Mexico basin. Earth Planet. Sci. Lett., 56 (in press). Kronfeld, J., Gradsztajn, Muller, H.W., Radin, J., Yaniv, A. and Zach, R., 1975. Excess 234 U: an aging effect in confined waters. Earth Planet. Sci. Lett., 27: 342- 345. Macfarlane, P.A., 1980. Distribution of radium-226 in the Cambro-Ordovician groundwater system, Tri-State region, Kansas, Missouri, and Oklahoma. Meet. South-Cent. Sect., Geol. SOC.Am., Rolla, Mo. (abstract). Osmond, J.K. and Cowart, J.B., 1976. The theory and uses of natural uranium isotopic variations in hydrology. At. Energy Rev., 14: 6 2 1 4 7 9 .
This Page Intentionally Left Blank
195
CARBONATE GEOCHEMISTRY OF VADOSE WATER RECHARGING LIMESTONE AQUIFERS
JOHN THRAILKILL and THOMAS L. ROBL
Department o f Geology, University o f Kentucky, Lexington, K Y 40506 (U .S.A.) Institute for Mining and Minerals Research, University of Kentucky, Lexington, K Y 40506 (U.S.A.) (Accepted for publication February 26,1981) ABSTRACT Thrailkill, J. and Robl, T.L., 1981. Carbonate geochemistry of vadose water recharging limestone aquifers. In: W. Back and R. Letolle (Guest-Editors), Symposium on Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54: 1 95-208. Vadose water from Lower Carboniferous rocks in Kentucky, U.S.A., was analyzed monthly in Mammoth Cave and in the caves and overlying soil at two other caves. Calcite saturation and equilibrium COz pressures of water were calculated after correcting for ion pairing and activity coefficients. Soil gas COz concentrations were determined directly. Both small flows of water which were usually depositing speleothems (vadose seepages) and larger flows (vadose flows) were studied. Some of the vadose seeps were often undersaturated with respect to calcite and showed little variation in Ca content (CV,, 7%), while others were supersaturated and showed positive correlations between Ca and soil COz or between Ca and discharge. At least some of the vadose seepages were believed to have been closed to COz introduction after leaving the soil zone. In the soil, COz varied with soil type, moisture and season. Ca increased with depth but saturation with respect t o calcite was not reached. Three types of vadose water may thus reach the aquifer undersaturated with respect to calcite: (1)vadose flows which transit the vadose zone rapidly; ( 2 ) low-Ca vadose seepage which is closed to COz as it transits the vadose zone, and which may remain undersaturated with respect t o calcite even after ventilation; and (3) high-Ca vadose seepage which leaves the soil zone slightly undersaturated, probably because of insufficient residence time.
<
INTRODUCTION
Although the number, size and distribution of voids in some relatively young limestones is primarily due to original depositional fabric, in most limestone aquifers it is mainly due to solution by waters with at least a major meteoric water component. It is apparent, therefore, that knowledge of the chemistry of such meteoric water is fundamental t o an understanding of the development of limestone aquifers. Meteoric water which descends through the vadose zone to the limestone aquifer is undoubtedly the major source of recharge, and its state of satuation with respect to calcite is important to the solutional development of
196
the aquifer. Specifically, an understanding of the mechanisms by which water transits the vadose zone without becoming saturated with respect to calcite or those which would produce undersaturation where the vadose water joins the water in the aquifer is of interest. Investigations of the state of carbonate equilibrium of water which is transiting the vadose zone have not been numerous. The equilibrium Pco2 (carbon dioxide partial pressure) and state of saturation with respect to calcite was studied by Holland et al. (1964) in Indian Echo Cave, Pennsylvania, and Luray Caverns, Virginia, and by Thrailkill (1971) in Carlsbad Caverns, New Mexico. These investigations showed that seepage water entering the cave was in equilibrium with a Pco2 much higher than that of the normal atmosphere and was often supersaturated with respect t o calcite. The present studies included four aspects not reported by Holland et al. (1964) or Thrailkill (1971): (1) repeated sampling of the same sites to elucidate seasonal variations; (2) measurement of discharge of the sources; (3) determination of the pH as soon as possible after the water entered the cave environment; and (4) investigation of the carbonate chemistry of the soil waters overlying at least some of the sites. S t u d y areas and methods Water samples were collected in Mammoth, Cascade and X caves, and water and gas samples were collected from the soil overlying Cascade and X caves. Mammoth Cave is in west-central Kentucky and Cascade and X caves are -3OOkm from Mammoth in northeastern Kentucky. Cascade and X caves are -3 km apart. All three caves are developed in limestones of Late Mississippian age and are within a few hundred meters of the outcrop of overlying Upper Mississippian and/or Pennsylvanian sandstones and shales. The temperatures of both the Mammoth Cave and the Cascade and X caves areas are similar, with a January mean minimum of -2OC and a July mean maximum of 32OC. The mean annual precipitation is 0.10m in the Mammoth Cave area and 0.13 m in the Cascade and X caves area and is fairly evenly distributed over the year (Karen and Mather, 1977). Approximately monthly samples were collected at Mammoth Cave during the period November 1967-June 1968 and during the period July 1975August 1976 at Cascade and X caves. Temperature, discharge, pH and alkalinity were determined at the time of sampling the cave water, and analyses for major ions were performed on a filtered (0.45pm) and acidified sample returned to the laboratory. Water samples collected from lysometers and open tubes emplaced at various depths in the soil at Cascade and X caves were treated similarly, and soil gas from open tubes was analyzed for C 0 2 , using a chromatographic gas analyzer. Details of the sampling and analytic methods for the Mammoth Cave work may be found in Thrailkill (1970) and for the Cascade and X caves investigations in Rob1 (1977).
197
Saturation with respect to calcite The degree of saturation of water samples with respect to calcite will be expressed in this paper by a saturation index Sc , defined as the ratio of the ion-activity product for calcite ( [CaZ+][COZ-], with the brackets indicating thermodynamic activities) and the thermodynamic solubility product of calcite. The calculations required t o determine this saturation parameter are extensively discussed in the literature (e.g., Garrels and Christ, 1965) and will not be repeated here. Two points, however, should be noted. First, the concentration of Ca2+ from which [ Ca2+] is obtained by multiplying it by an activity coefficient is somewhat less than the total calcium determined analytically (hereafter denoted as Ca) because of the existence of Cacontaining ion pairs. Because the concentration of such ion pairs and hence the concentration of Ca2+is a function of several variables, including the pH and the concentration of other ions in solution, Ca is a better indicator of the calcite solution and precipitation history of the water. Second, in the pH range of the waters examined, [COi-] is best determined from the second dissociation constant of carbonic acid ( K , ), [HCO;] (derived from the alkalinity titration), and [H'] . Similarly, the equilibrium partial pressure of CO, is determined from the first dissociation constant of carbonic acid ( K , ), K c o , , [HCO;] and [H'] . Both of the parameters are thus related t o [H'] , the antilog of the negative pH. Because the escape of CO, when samples are removed or flow from a high-PCo2 environment causes a rapid rise in pH, and because an increase in pH of only 0.3 will result in a doubling of the value of Sc calculated, any delay in determining pH will cause the waters t o appear t o be more nearly saturated or more greatly supersaturated than they were in the environment under study. In Mammoth Cave, it was possible in most cases to measure the pH of single drops as they emerged into the cave, and the Pco2- and Sc -values are believed representative of the water as it enters the cave. In Cascade and X caves, however, the seepage was from large stalactites on which a t least some of the water was flowing as a thin film on the outside and had therefore been in contact with the cave atmosphere for some time. All of the seepage samples from these two caves showed low values of Pco2 and values of Sc >1, indicating supersaturation with respect to calcite. The various calculations were conducted using the FO R TR A N program IO NPAI R , which yields results comparable to programs used by other workers (Nordstrom et al., 1979, p.879).
RESULTS AND DISCUSSION
Table I summarizes some data obtained for the fifteen cave sites sampled. It is apparent that there are substantial differences between the sites. The
TABLE I
+
Means (G(*') and Q ( * 2 ) ) and coefficients of variation (CVc, and CVQ) for calcium and discharge, respectively; equilibrium Pco,(*3); saturation index (SC);and classification of cave sites Site no. Numberof Ca CVCa Q CVQ PCO, ( ~ O - ~ a t m . ) S c Classification * 4 samples (mg/l) (a) (ml/s) (%) max. min. max. min.
O3
to
Mammoth Cave: 1 2 3 4 5 6
8 4 8 7 7 7
14 29 36 22 82 59
35 21 3 7 7 5
170 180*' 0.0069 0.20 0.15 0.0046
48 0 25 43 17
0.13 0.098 0.95 0.20 1.9 1.2
0.076 0.072 0.12 0.054 0.52 0.14
0.22 0.87 0.72 0.95 1.4 2.6
0.013 0.29 0.11 0.21 0.38 0.25
VF VF LS LS HS? HS?
10 * 6 16 *7 1 4 *8 1 5 *9 4 11 10 *lo
68 88 58 70 45 49 50
24 13 5 9 13 12 15
0.0080 0.50 0.076 0.10 0.13 0.25 1.o
62 95 11 17 53 300 200
0.20 0.64 0.43 0.37 0.075 0.091 0.093
0.085 0.095 0.12 0.13 0.050 0.057 0.052
6.1 9.7 3.1 4.4 4.2 4.3 5.1
2.1 1.4 1.3 2.0 2.1 1.3 1.1
HS HS HS HS HS HS HS
12 11
53 80
12 6
4.5 0.10
120 43
0.12 0.28
0.051 0.12
5.7 5.7
1.6 2.6
HS? HS
Cascade Cave: 7 8 9 10 11 12 13
X Cave: 14 15
*' 1mg/l= kg/m3 ; *2 1ml/s = m3 / s ; *3 atm. (= % atm.) N l o 3 Pa; *4 VF = vadose flow, LS = low-Ca vadose seepage, HS = high-Ca vadose seepage; *' representative discharge; * 6 11for Ca; *7 17 forPCO2 and S c , 1 8 for Ca; *8 1 5 for Q; *9 16 for Q and Ca; *lo 11for Q.
199
,001
pco,
.01
.1
1
Fig. 1. Log P,o,-log Ca (total Ca) diagram showing values of pH and SC in region where - 2mca2+, assuming all activity coefficients unity, n o species other than H+, Ca2+, mHC03 HzC0; (including COz, a s ) , HCOi, OH- and Cog-, and T = 298.15 K. Shaded areas where 2mco:10%mHCO; (upper left) or mH+> 10%2mCa2+(lower right). See text for discussion of labeled points and paths.
>
The most striking of these is the large between-site variation in mean Ca (5) together with the small within-site variations (as indicated by the coefficient of variation, CVca) for most of the sites which had a mean dism3/s (1 ml/s) or less. charge ( Q ) of In an earlier paper (Thrailkill, 1968, pp. 29-31), the terms vadose seepage and vadose flows were proposed to describe smaller and larger flows of water, respectively, which descend through the vadose zone and recharge the carbonate aquifer. Evidence available at that time suggested that these two types of vadose water may be chemically distinct. If a mean discharge of m3/s is taken as the upper limit of discharge for vadose seepage (such flows are usually as successive drops rather than a continuous flow), twelve of the fifteen sites would be considered vadose seepage and three vadose flows. Before further discussion of the chemistry of the sampled water, it will be useful to consider Fig. 1, which is a plot of logPco, vs. log total calcium (Ca) on which various relationships and reaction paths may conveniently be examined. On such a diagram, if there are no ions (or ion pairs) other than Ca2+, H+, OH-, H2CO: (including unhydrated COz), HCO, and Cog- in solution and all activities are assumed equal to molalities, isopleths of pH and Sc plot as straight lines in the region shown where mHCO3may be considered equal to'2rncaz+ (see Thrailkill, 1968, pp. 29-30 for a more
200
complete discussion). In addition, a number of reaction paths can be plotted and various environments of interest can be identified, as shown on Fig. 1. For example, path A represents water relatively pure water (Ca = 2mg/l) equilibrating in the soil zone from the Pco, of the normal atmosphere (point w ) to a Pcoz of 0.05 atm. typical of a soil atmosphere (point x). If this water then dissolves calcite to saturation at this soil Pco2 it will follow path B t o point y. Re-equilibration to the normal atmosphere Pco2 while the water remains saturated with respect to calcite (point z ) is shown by path C. Other paths of interest in the discussions which follow are: D and F , which represent CO, escape without solution or precipitation of calcite; E , representing calcite precipitation at constant CO, , and G, which illustrates the compositional changes which take place when water at point x dissolves calcite in an environment closed t o CO, . Vadose seepage
Six of the sites (3-6, 9 and 1 5 ) tentatively classified as vadose seepage had a coefficient of variation for Ca of 7% or less, which approaches the analytic uncertainty for Ca of -2% ((7). The individual determinations for these six waters are plotted in Fig. 2, which is an enlarged portion of Fig. 1.
Fig. 2. Log PCO, l o g Ca (total Ca) diagram showing analyses of vadose flows (crosses) where Ca = f(discharge) and vadose seeps with CVca < 7% (filled and open circles). Ca = f(soi1 PCo, ) for site 9 (open circles). Underlined values are maximum and minimum Sc for each site. Values adjacent to arrows (closed-system reaction paths) are: “calculated Ca (Pco, = 0.05 atm.); (Ca = 2 1 mg/l) calculated Pcol ”.
20 1
Although the line from Fig. 1indicating calcite saturation is shown, it should be noted that its position is only approximate, since the simplifying assumptions discussed (e.g., absence of other ions, activitives equal t o molalities) do not hold for these samples. The maximum and minimum Sc for each site is shown. The considerable range of Pco2 for samples from an individual site is interpreted to represent varying degrees of outgassing due t o ventilation rather than to variations of Pco2 in the soil water feeding the drip, since significant variations in soil water Pco2 would be reflected in variations of Ca as the higher-Pco, soil waters would dissolve more calcite. Although an effort was made to measure pH as soon as possible after the water entered the cave (see Introduction), the outgassing of C 0 2 is such a rapid process that some variation would be expected even if ventilation had not already occurred in higher cavities in the rock. The variations in C 0 2 shown by the seepages could not be correlated with discharge, season, or temperature. Diagrammatic cross-sections of the sampled sites are shown in Fig. 3 . Comparison of Fig. 2 with Fig. 3 shows that Ca of the vadose seepages in Mammoth Cave (sites 3-6) decrease with depth below the surface, and that site 4 , which had the lowest Ca, is beneath a sandstone caprock. Although no study of the soils was undertaken at Mammoth Cave, it is likely that the low Ca of site 4 is due to its being fed from soil water in a non-carbonate soil, and that the progressive increase in Ca in drips underlying the downslope soils reflects a progressive increase in carbonate soil fractions with distance X Cave ........
-
I - I
cascad. and x c
1~ - ~ r
8 7 .
~
11-13
1-
Cascade Cave
Mammoth C a v e
I
10 9
15
14
0 0
KENTUCKY 0
200
km
Fig. 3. Index map ,and cross-sections showing location of sites. Dotted pattern shows capping sandstone and shale at Mammoth Cave and soil at Cascade and X caves.
202
below the sandstone caprock. It is of interest to note that the vadose seepage water has apparently acquired its Ca in the soil, and this Ca is not related to the vertical distance of percolation through limestone, which is greatest for the lowest-Ca vadose seepage (site 4 ) . This apparent lack of calcite solution by the water as it percolates through the limestone is most easily explained by assuming that the water is closed to COz as it enters the limestone and thus follows a reaction path on Fig. 2 similar t o path G in Fig. 1.Note that path G initially has a slope of -1 but then flattens out as the initial COz is used up. Holland et al. (1964) discussed the possibility that vadose waters might equilibrate with limestone in a closed system. Deines et al. (1974) concluded that closed-system solution was the major process of carbonate solution in Pennsylvania, in a study that included stable-carbon-isotope data, and Pitman (1978) interpreted data from spring waters in the Chalk, England, as having dissolved calcite under closed-system conditions. The closed-system reaction paths for the highest-CO, analyses of sites 3-6 were calculated approximately (assuming [Ca"] = Ca, no ion pairing, and 25°C) and a short segment of the path is shown on Fig. 2 together with the calculated initial Ca assuming an initial Pcoz of 0.05 atm., and the calculated Pco2 assuming an initial Ca of 21 mg/l (that of the site-4 analysis). A range of soil Pco2 from 0.002 to 1atm. is required to explain the variations in Ca if the initial soil Ca was 21mg/l, which would seem t o confirm the earlier interpretation that there are variations in soil Ca between the sites. It should further be noted that, although the closed-system reaction paths have relatively small slopes, the reaction paths for the various analyses within a site are sufficiently different so that CO, escape remains the best explanation for the range of analyses within each site. Soil gas collecting tubes and lysometers to sample soil waters were installed at four locations above sampled drips in Cascade Cave (sites 16 and 1 7 ) and X Cave (sites 18 and 1 9 ) as shown on Fig. 3. Pcoz measured from the gas wells and equilibrium Pco, in the water withdrawn from the lysometers was higher during the summer when plant growth was active and soil temperatures were high, and at any one time increased with depth. The Pcoz of sandy soils with higher permeability tended to be lower than clayey soils, and deep drying of the soils caused a decline in P c o z . Some representative values are shown in Table 11. The relationship between the Ca content of the Cascade and X caves' vadose seepage and the same day Pco, of the soil sites was examined. Although no systematic data are available on the travel time between the soil and the sampling points for these vadose seepages, their shallow depth and observations of discharge variations related t o rainfall suggest that the travel time is short; probably on the order of one or two days a t most. The data used were from the 0.60-m gas collecting tubes, which were the deepest sampling points with a reasonably complete record. The data for soil site 16 was so incomplete that soil site 1 7 was used for all of the Cascade Cave sites. Three of the sites (8-10) showed a slope of the regression of log Ca on
203 TABLE I1 Representative values of PCO, and Ca from soils above Cascade and X caves Sampling date
Soil gas pCo2 (IO-'
atm.(*'))
Water PCOZ (10-2
Depth ( m ) 0.05
atrn.(*l))
Ca (mg/I (*')I
0.15
0.30
0.60
1.2
0.60
1.2
0.60
0.55 0.60
10.0 0.90
8.5 2.3
11.0 3.5
-
3.5
-
120
1.5 0.17
1.o 0.30
2.6 0.70
2.7 0.80
2.7
-
0.08 0.08 0.15
0.18 0.10 0.30
0.40 0.10 0.40
1.5 0.33 1.1
-
0.10 0.06 0.15 0.16
0.23 0.08 0.23 0.40
0.30 0.12 0.23 0.50
1.7 0.40 0.60 1.6
1.2
Site 1 6 : Jun. 5, 1975 Oct. 9, 1975
-
Site 17: Oct. 9, 1975 Mar. 3. 1975
-
0.99*3
-
36 *3
Site 1 8 : Oct. 6, 1975 Jan. 14, 1976 Jun. 25, 1976
--
-
-
Site 1 9 Oct. 6, 1975 Jan. 14,1976 May 19,1976 Aug. 24,1976
*' lo-'
atm. = (= % atm.)
l o 3 Pa; *'
1m g / l =
3.7 1.o 1.1 0.59 2.3 1.8*4 -
9.9 *4
110 81 53 -
kg/m3 ; * 3 1.61 m ; *4 0.64 m.
logPc0, (soil) which was significantly different than zero at an a of 0.05, and one (14) at an a of 0.1. The slope was positive in each case, indicating the expected relationship between high Pcoz in the soil and high Ca of the water. Site 9, which had a low coefficient of variation of Ca, is shown on the log Ca-log Pco2 plot of Fig. 2; the remaining three are plotted in Fig. 4. A similar examination of the possible relationship between Ca of the cave sites and their discharge was also undertaken. Two of the sites (7 and 1 0 ) had a positive slope ( a = 0.05) for the regression of log Ca on log discharge, of which one (10) also was related t o soil Pcoz (see above). Both of these are apparently precipitating calcite, with the amount of Ca loss being greatest at low discharges. Their position on the log Ca-log Pco, plot (Fig. 4)suggests a trend to saturation at the Pco, of the cave atmosphere. Site 7 showed no significant correlation ( a = 0.1) between (log Mg) and (log discharge), as would be expected if the above explanation is true, since Mg is generally not lost by precipitation on stalactites (Thrailkill, 1971, p.687). Site 1 0 did show a significant ( a = 0.05) positive slope for this regression, but this is probably
204
6X
8 X'IO-4
.001
pco,
.002
,003
Fig. 4 . LogPco,-log Ca (total Ca) diagram showing analyses for sites 8 and 14 (filled circles) where Ca = f(soi1 PCo2), site 7 (open circles) where Ca = f(discharge), and site 10 (crosses) where Ca = f(soi1 Pco,, discharge).
due to the correlation between Ca and Pco, discussed above (the sample size was judged insufficient t o attempt a multiple regression evaluation). To summarize, factors controlling the chemistry of all of the twelve sites tentatively classified as vadose seepages, except 11-1 3, have been discussed, as well as site 1 4 which was initially classified as a vadose flow. These latter four sites will be considered further in the next section.
Vadose flows The location of site 2 was such that its discharge could not be accurately determined, but it was approximately the same as site 1. The discharge of these two vadose flows was substantially higher than any other site. Unlike the vadose seepages, their variation in log Pco2 was less than their variation in log Ca (Fig. 2), and in contrast to the two vadose seepage sites which showed a positive correlation between discharge and Ca content (ascribed to calcite precipitation), site 1 showed a significant negative correlation between these two variables (a, = 0.05). Both of these vadose flows were emerging from solutionally enlarged vertical conduits, and their location in Mammoth Cave beneath the edge of the sandstone caprock (Fig. 3) suggests they represent an accumulation of low-Ca water which is dissolving calcite as it passes vertically through the limestone, with more solution occurring at lower discharges due to longer transit times and more intimate contact with the rock. The low and relatively constant Pco, only slightly greater than the Pcoz of the cave atmosphere indicates that these waters have nearly equilib-
205
rated with the cave atmosphere in the open conduits through which they are flowing. Site 1 4 in X Cave, which was tentatively classed as a vadose flow on the basis of its mean discharge greater than m3 / s , is probably more related to the vadose seepages despite its variable discharge (only slightly more variable than site 8, see Table I) and low Pco, (which probably represents nearly complete ventilation in the cave environment prior to sampling). Its correlation between Ca and soil Pco2 (as discussed earlier), the wide variation in C 0 2 relative to Ca (Fig. 4), and its supersaturation with respect t o calcite (Table I) makes it resemble the vadose seepages more than the flows. The same conclusion may be drawn from the remaining sites 11-13 (Table I). Sites 12 and 13 are separated by only a few centimeters and were originally sampled together as site 11. These waters have so completely equilibrated to the Cascade Cave atmosphere that they furnish little information on their original chemistry, except to illustrate the slowness of the process by which calcite-supersaturated water attains saturation by precipitation.
State of saturation with respect to calcite of soil water Although considerable effort was expended in obtaining soil samples from the four sites overlying sampled drips in Cascade and X caves, the numerous difficulties encountered precluded the collection of large amounts of data. These difficulties included dilution from higher water sources, absence of sample due to low soil moisture, small sample volumes, temperature disequilibria and possible chemical reactions in the lysometers prior t o sampling. The apparently most consistent set of data was obtained from soil site 19. These data suggest that essentially all solution of Ca is occurring at the soilrock interface. Ca content of samples taken from a depth of 0.64m ranged from 1 2 t o 22 mg/l while samples taken on the same days from a lysometer at the soil-rock interface 1.20m below the surface ranged from 76 t o 89 mg/l. N o calcite could be detected in soil samples from various depths, and it is believed that all of the Ca in the shallower samples was derived from upward diffusion. Values of Sc calculated for the six samples obtained from the lysometer installed at the soil-rock interface were 0.66, 0.44, 0.66, 1.47, 0.74 and 0.64, respectively. The one determination indicating supersaturation may have been contaminated by crystal growth in the lysometer following a decline in soil P c o z . Neglecting this value, the remaining values range from 0.44 t o 0.74 and have a mean of 0.63. To evaluate the hypothesis that the undersaturation of the soil water may be due to the kinetics of the calcite solution processes, calculations of the solution rate were performed using both the data of Morse (1978) and Plummer et al. (1979) and are shown in Table 111. It is apparent that there is little agreement, either in the overall average rates or in the calculated change
206 TABLE I11 Times t o attain various degrees of saturation (calculated using rate for lower SC of range) of soil water at site 1 6 , using Indian Ocean rate equation of Morse (1978, p.348) and equation 1 5 of Plummer et al. (1979, p.546) SC
Morse
Plummer et al.
0.5-0.6 0.6-0.7 0.7-0.8 0.8-0.9 0.9-1.0
2.4 min. 7.0 min. 28 min. 3.4 hr. 5.1 days
5.8 s 6.6 s 8.3 s 11.0 s 22.0 s
Soil is assumed to be composed entirely of 0.2-mm diameter spherical particle of calcite and have a porosity of 34.1%. Values of k,, [H'] , and [H'] (surface) needed for Plummer et al.'s equation 1 5 calculated from measured Pcoz and T but neglecting ions other than Mg2' in equilibrium calculations.
in rates as saturation is approached. The nature of the soil system is undoubtedly so complex that the absolute times shown, which are based on rather arbitrary assumptions of surface area and equilibration of the pore water, would not be expected t o be better than order-of-magnitude estimates. The lower rates of Morse (1978) may be due in large part to the nature of the material he used (calcareous ocean sediment), but the rapid decline in rates as saturation is approached contrast sharply with the relatively minor decline of Plummer et al. (1979), who found no such decline until equilibrium was nearly attained at Sc 2 0.9 (L.N. Plummer, pers. commun., 1980). The undersaturation of the soil samples observed at soil site 19 may alternatively be due t o the presence of inhibiting ions, as discussed by Plummer and Wigley (1976) and Plummer et al. (1976).
CONCLUSIONS
These studies have shown that there is apparently a chemical distinction between vadose flows and vadose seepage, as was earlier proposed, in that the vadose flows studied were originally more undersaturated with respect t o calcite and are actively dissolving calcite as they pass through the vadose zone. The source of the vadose flows in Mammoth Cave is not known, and they may represent concentrations of soil water draining from low-Ca soils and their chemical nature may be different than flows which represent the diversion of surface water underground in areas of high-Ca soils, which were not available for sampling in this study. In any event, flows of this type, which are probably common in areas where low-Ca soils are present at the surface, will introduce undersaturated water into the underlying aquifer, with the degree of undersaturation directly related t o the discharge.
201
It has also been possible t o identify two types of vadose seepage. The first, which may be termed low-Ca vadose seepage, is derived from low-Ca soils and its degree of calcite saturation and Ca content are controlled by the available soil Ca and soil Pco2. Such seepage does not appear to dissolve significant amounts of Ca as i t transits the vadose zone, probably because it is closed t o further additions of C 0 2 and may thus arrive at the aquifer substantially undersaturated if its Pco, remains at the high levels acquired in the soil zone. Further, some of this low-Ca seepage (such as site 4 ) will remain undersaturated even if ventilated t o the Pco2 of a cave atmosphere (Fig. 2). The other type of vadose seepage, termed high-Ca uadose seepage, leaves the base of the soil zone in equilibrium with a high Pcol but slightly undersaturated with respect to calcite. Such high-Ca vadose seepage, therefore, will also provide undersaturated recharge to the underlying aquifer if it remains unventilated during its passage through the vadose zone. High-Ca vadose seepage is probably the most abundant source of recharge t o limestone aquifers which are not overlain by other lithologies, and it is of interest to note that the solutional development of such aquifers would be inhibited by the presence of ventilated cavities in the overlying vadose zone. Several studies (Jacobson and Langmuir, 1970; Thrailkill, 1970, 1972; Shuster and White, 1972; Robl, 1974) have shown that the water being discharged from limestone aquifers has a Pco, intermediate between that of the soil zone and that of the normal atmosphere. The recharge of the aquifer by either type of vadose seepage which had not suffered ventilation would accomplish this, but it should be noted that vadose flows, which generally occupy ventilated conduits even if they are derived from high-Pco, soil drainage rather than low-Pco, surface water, would not. Also, it has been shown that the degree of saturation of a mixture of waters in equilibrium with different values of Pcoz will be less than the simple weighted average of the saturation of the two waters (Bogli, 1963; Thrailkill, 1968; Wigley and Plummer, 1976) which would produce further undersaturation.
ACKNOWLEDGMENTS
Appreciation is expressed t o M.T. Osolnik, R.H. Postley, and the superintendent and staff of Mammoth Cave National Park for their assistance in the Mammoth Cave study, which was funded by the Office of Water Resources Research (now the Office of Water Research and Technology), U.S. Department of the Interior; and to J. Tierney, park naturalist, and the staff of Carter Caves State Resort Park for their cooperation in the study of Cascade and X caves.
208 REFERENCES Bogli, A., 1963. Beitrag zur Enstehung von Karsthohlen. Hohle, 14: 63-68. Deines, P., Langmuir, D. and Harmon, R.S., 1974. Stable carbon isotope ratios and the existence of a gas phase in the evolution of carbonate ground waters. Geochim. Cosmochim. Acta, 38: 1147-1164. Garrels, R.M. and Christ, C.L., 1965. Solutions, Minerals and Equilibria. Harper and Row, New York, N.Y., 450 pp. Holland, H.D., Kirsipu, T.V., Huebner, J.S. and Oxburgh, U.M., 1964. On some aspects of the chemical evolution of cave waters. J. Geol., 72: 36-37. Jacobson, R.L. and Langmuir, D., 1970. The chemical history of some spring waters in carbonate rocks. Ground Water, 8(3): 5-9. Karen, P.P. and Mather, C. (Editors), 1977. Atlas of Kentucky. University Press of Kentucky, Lexington, Ky., 182 pp. Morse, J.W., 1978. Dissolution kinetics of calcium carbonate in sea water, VI. The nearequilibrium dissolution kinetics of calcium carbonate-rich deep sea sediments. Am. J. Sci., 278: 344-353. Nordstrom, D.K., Plummer, L.N., Wigley, T.M.L., Wolery, T.J., Ball, J.W., Jenne, E.A., Bassett, R.L., Crerar, D.A., Florence, T.M. Fritz, B., Hoffman, M., Holdren, Jr., G.R., Lafon, G.M., Mattigod, S.V., McDuff, R.E., Morel, F., Reddy, M.M., Sposito, G. and Thrailkill, J., 1979. A comparison of computerized chemical models for equilibrium calculations in aqueous systems. In: E.A. Jenne (Editor), Chemical Modeling in Aqueous Systems. Am. Chem. Soc., Symp. Ser., 93: 857-892. Pitman, J.I., 1978. Carbonate chemistry of groundwater from Chalk, Givendale, East Yorkshire. Geochim. Cosmochim. Acta, 42: 1885-1897. Plummer, L.N. and Wigley, T.M.L., 1976. The dissolution of calcite in CO2-saturated solutions at 25OC and 1 atmosphere total pressure. Geochim. Cosmochim. Acta, 40: 191-202. Plummer, L.N., Vacher, H.L., Mackenzie, F.T., Bricker, O.P. and Land, L.S., 1976. Hydrogeochemistry of Bermuda: a case history of ground-water diagenesis of biocalcarenites. Bull. Geol. SOC.Am., 87: 1301-1316. Plummer, L.N., Wigley, T.M.L. and Parkhurst, D.L., 1979. Critical review of the kinetics of calcite dissolution and precipitation. In: E.A. Jenne (Editor), Chemical Modeling in Aqueous Systems. Am. Chem. Soc., Symp. Ser., 93: 537-573. Robl, T.L., 1974. Factors affecting the state of saturation of the waters of Puckett Spring Creek with respect to calcite. Thesis, University of Kentucky, Lexington, Ky., 59 pp. Robl, T.L., 1977. Factors controlling the geochemistry of vadose and stream waters in a carbonate terrain. Dissertation, University of Kentucky, Lexington, Ky., 200 pp. Shuster, E.T. and White, W.B., 1972. Source areas and climatic effects in carbonate groundwaters determined by saturation indices and carbon dioxide pressures. Water Resour. Res., 8: 1067-73. Thrailkill, J., 1968. Chemical and hydrologic factors in the excavation of limestone caves. Bull. Geol. SOC.Am., 79: 19-46. Thrailkill, J., 1970. Solution geochemistry of the water of limestone terrains. Univ, Ky. Water Resour. Inst., Lexington, Ky., Res. Rep. 19, 125 pp. Thrailkill, J., 1971. Carbonate deposition in Carlsbad Caverns. J. Geol., 79: 683-695. Thrailkill, J., 1972. Carbonate chemistry of aquifer and stream water in Kentucky. J. Hydrol., 16: 93-104. Wigley, T.M.L. and Plummer, L.N., 1976. Mixing of carbonate waters. Geochim. Cosmochim. Acta, 40: 989-995.
20 9
A GEOCHEMICAL METHOD OF DETERMINING DISPERSIVITY IN REGIONAL GROUNDWATER SYSTEMS WARREN W. WOOD*
Geoscience Department, Texas Tech University, Lubbock, TX 79409 (U.S.A.) (Accepted for publication February 26, 1981)
ABSTRACT Wood, W. W., 1981 A geochemical method of determining dispersivity in regional groundwater systems. In: W. Back and R. Lbtolle (Guest-Editors), Symposium on Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54: 209224. A method for determining values of dispersivity for large-scale regional aquifer systems is developed using the concept of hydrochemical facies. In this approach the aquifer system is viewed as a large imaginary column. Analysis of hydrodynamic dispersion is made by observing the relative position of solute concentration contours in space. That is, because the natural aquifer is so large and the flow rate so slow, time has been interchanged for space, and solute ratios based on concentration contours are substituted for column effluents, which is commonly used in such analyses. The proposed method of evaluation of dispersivity yields a true value if conservative tracers are used, or homogeneous distribution of reacting materials can be assumed.
INTRODUCTION
Hydrologists and petroleum engineers have long been aware of hydraulic heterogeneity in aquifers and petroleum reservoirs; however, it is only in the last 10-15 years that we have recognized its practical significance. Management decisions concerning secondary and tertiary recovery of oil, burial of toxic chemical and radioactive waste and predicting groundwater quality all require quantitative knowledge of the heterogeneity of the flow systems. Consider a column experiment in which we have fluid flow in a saturated porous medium. If we introduce a small mass of tracer at a point in the flow field, we observe that, over time, the concentration is diluted in the direction of bulk liquid flow as a result of mixing with the interstitial fluid and is Mixing of the referred to as longitudinal hydrodynamic dispersion (DL). tracer perpendicular t o the direction of bulk flow is referred t o as transverse hydrodynamic dispersion (DT) and is frequently 20-50% smaller than longitudinal hydrodynamic dispersion. The dilution of the tracer occurs as a
210
result of both mechanical mixing and molecular diffusion. In most aquifers molecular diffusion (Dd), which results from a difference in the chemical activity between the induced fluid and that originally present within the porous medium, can be shown t o be negligible compared t o mechanical mixing and is therefore usually ignored. However, in the material with low hydraulic conductivity or low hydraulic gradients, molecular diffusion may dominate over mechanical mixing. Mechanical mixing results from at least three processes which include tortuosity of the flow path, difference in interstitial pore size in the direction of the flow path, and difference in flow velocity through an individual interstitial pore space. Mechanical mixing (D,) is generally characterized by two components: (1)dispersivity ( a ) which is a property of the medium and the fluids, and has the unit of length (L); and (2) average fluid velocity (V) having unit of length per time (L/T) such that:
D, = a v (L2/ T) The coefficient of longitudinal hydrodynamic dispersion can then be expressed to incorporate both mechanical mixing and molecular diffusion :
D,
D, + D d (L2/ T ) Chemical reactions that attain equilibrium within the confines of the experimental setup do not affect the degree of mixing as long as they do not change the physical properties of the medium or the fluid; rather they change only the velocity at which the tracer moves through the porous medium. However, reactions that are kinetically controlled, and do not reach equilibrium in the temporal or spatial confines of the experiment, do affect the concentration distribution and cannot be readily separated from mechanical mixing. Dispersivity is normally determined by laboratory or small-scale field experiments in which a small sample of the aquifer or reservoir is stressed and the results extrapolated to the regional system. This approach has two important limitations: (1)dispersivity is scaledependent (Bredehoeft et al., 1976), i.e. the greater the contrast in hydraulic conductivity the greater will be the values of dispersivity and at present there is no satisfactory way to scale laboratory-derived values t o regional sized systems; and (2) laboratory samples, by necessity, represent only a minute fraction of the aquifer system. Several groundwater hydrologists have investigated the problem of dispersivity on a regional scale. Winograd (1974) and Winograd and Pearson (1976) evaluated movement of the isotopes of water and carbon in a solutioncontrolled carbonate aquifer system in the southcentral Great Basin of Nevada. They found what was termed “megascale channeling”, that is a zone of large hydraulic conductivity, at least 11km wide, that moved the water 30-50 km from recharge t o discharge area without significant mixing. Their data indicated very small dispersivity for this zone of the aquifer as =
211
there was little mixing with the water or solutes in the surrounding aquifer but large dispersivities if the aquifer is viewed in its entirety. N o calculations on the magnitude of the observed dispersivity were made. In a theoretical investigation of large-scale dispersivity, Schwartz (1977) numerically simulated a hypothetical regional aquifer system and found that it was impossible to define a unique dispersivity value unless the hydraulic conductivity contrasts were homogeneously distributed within the aquifer. Schwartz and Muehlenbachs (1979) concluded that the distribution of 2 H and “0, in the Milk River sandstone aquifer in southeastern Alberta, Canada, was controlled by what they referred to as megascopic dispersion or mixing on a regional scale. They were, however, unable to identify the scale of dispersion responsible for the observed isotopic patterns. Smith and Schwartz (1980) demonstrate that dispersivity values controlled by diffusion are not applicable to systems controlled by mechanical mixing. That is, scaling-up from laboratory experiments that may be controlled by diffusion to regional flow systems is inappropriate because different mechanisms are controlling the concentration of ions in question. Clearly, more work is necessary in evaluating this parameter if it is t o be used in modeling solute transport in largescale regional aquifer systems. As analysis of hydrodynamic dispersion must be performed on a transient chemical system it is prohibitively expensive and would require years of real time to artificially stress a significant portion of an aquifer system, the author proposes for your considerations a method that uses naturally stressed regional systems t o evaluate dispersivity. In many natural aquifer systems, the chemical constituents and solute concentrations change in a regular manner in the direction of flow. The approach suggested requires that one visualizes these changes as occurring in a large imaginary column in which time and space have been interchanged. In a typical laboratory evaluation of dispersivity, a tracer of known concentration is introduced at the top of the column of material t o be tested and allowed t o flow through it. Its concentration is monitored along with flow volume at the column bottom. This process is usually completed in a few hours. In the natural system a millennium would be required for the introduced “tracer” to move through the system because water movement is slow and the mass of reacting material is large. Accordingly, the author suggests that the concentration of any solute or isotope observed in regional aquifer systems can be considered a “tracer”. The tracer may be something introduced in the classical sense, such as an unusual isotope or isotopic ratio resulting from climatic or tectonic events; it may be a residual solute, such as chloride, remaining after a marine origin or historical seawater intrusion, or it may be some solute or isotope generated in the recharge area of the aquifer from mineral dissolution due to changing equilibrium conditions. The chemical composition of natural systems appears to us mortals as though it is frozen in time, and equivalent to an instant in the typical laboratory column. Consequently, instead of evaluating the concentration of solute in the effluent as one
212
would normally do in a laboratory tracer test, it is necessary to evaluate the space between the solute concentration contours. Regional aquifer systems can then be viewed as large imaginary columns in which time and space have been interchanged. A concentration ratio is calculated between the concentration contours observed in the imaginary column and the solute leaving the bottom of the imaginary column, that is, the last contour. The pore volumes of water that have passed each concentration contour are calculated using Darcy's law and a graph of concentration ratio vs. pore volumes is made for use in determining dispersivity. The overall approach is an attempt to utilize some rather simple, easilyobtained chemical observations of natural aquifer systems to determine numerical values for dispersivity that are valid on a regional scale. The Aquia Formation in southern Maryland was selected t o illustrate this approach because significant solute changes occur in a regular manner in this aquifer (Back, 1966), because published chemical analyses were available (Woll, 1978) with which to contour the changes; and because the groundwater flow system in this aquifer is relatively simple and had recently been simulated in a digital model by Kapple and Hansen (1976).
GEOHYDROLOGY
The Aquia Formation of Paleocene age is a unit in the U.S.A. Atlantic Coastal Plain sequence and provides an important source of groundwater in Calvert, St. Mary's and Charles counties, Maryland (Fig. 1).Reports by Ferguson and Martin (1953), Otton (1955), Rasmussen and Slaughter (1957), Overbeck and Slaughter (1958), Mack (1962, 1966), Weigle et al. (1970) and Mack et al. (1971) discuss various details of the water resources of the Aquia Fm. Kapple and Hansen (1976) modeled the hydrology of the Aquia Fm., and Glaser (1971) and Hansen (1974) have given important geological descriptions of its aquifer. The geohydraulic descriptions given in this paper are summarized from the above reports. -
I-
A
V l RGlNI A 39' DC
I
L
'
-
. .. . -.
I
Fig. 1 . Map showing location of study area in the State of Maryland, U.S.A.
213
Fig. 2. Schematic cross-section from Prince Georges through St. Mary’s counties. N o scale intended.
The Coastal Plain formations in southern Maryland comprise a wedge of sediments whose thickness increases from 170 m in western Charles County to -1 km beneath the extreme southern end of St. Mary’s County. The formations generally slope between 2 and 20 m/km in an east-southeasterly direction although some local structures deviate from this general pattern (Jacobeen, 1972). Nonmarine Cretaceous age (Potomac Group) clastic sediments comprise the bulk of the Coastal Plain section in this area. However, the several marine formations of Late Cretaceous and Tertiary ages that act as aquifers in this area include the Magothy, Aquia, Nanjemoy and Piney Point Fms. A highly generalized cross-section (Fig. 2) shows the stratigraphic relationships. The Aquia Fm. is a marine unit deposited on a continental shelf and, according t o Hansen (1974), exhibits three facies: a nearshore of medium to coarse texture sand, an inner marine shelf of fine t o silty sand, and an outer marine shelf composed predominately of clay. The Aquia Fm. consists chiefly of quartz and glauconite. Although accounting for less than 50% of the sand-sized particles, the glauconite imparts a characteristic greenish cast to the unweathered formation, hence the name “Green Sand” is frequently applied to the formation. Figs. 3 and 4,modified from Kapple and Hansen (1976), illustrate the thickness and transmissivity of the aquifer in the area. The Aquia aquifer is artesian over the area with heads as much as 15m above sea level. Fig. 5, a potentiometric and outcrop map modified from Otton (1955), illustrates the head distribution prior t o development. Kapple and Hansen (1976) used a storage coefficient of 0.0003 in modeling the groundwater in this system. The Aquia Fm. is overlain conformably by the Nanjemoy Fm. which consists of silt and fine sand. In much of southern Maryland the basal 5-10m is composed of a pink kaolinite unit called the Marlboro Clay Member (Glaser, 1971). The Aquia Fm. unconformably overlies beds ranging in age from Early Cretaceous to Early Paleocene; however, the marine Brightseat Fm. of the Lower Paleocene lies beneath the Aquia forming a relatively homogeneous confining bed over most of the area of interest. Kapple and
214
Fig. 3 . Map showing the thickness of Aquia Formation in meters (modified from Kapple and Hansen, 1976).
Fig. 4 . Map showing transmissivity of the Aquia Formation in square meters per day (m2/day) (modified from Kapple and Hansen, 1976).
21 5
Fig. 5. Potentiometric head under natural, unpumped conditions (modified from Otton, 1955).
Hansen (1976) used a hydraulic conductivity of 4.3 m/day for the confining bed in their simulation of the groundwater system. A specific storage of 2.3 lo-’ m-l was used for all calculations in this simulation. Kapple and Hansen (1976) did not specifically include leakage t o and from the Nanjemoy and Brightseat Fms. but included it indirectly in their calibration of the model. Under natural conditions in the study area some leakage probably occurs downward across the formation from the Nanjemoy to the Aquia and from the Aquia t o the Brightseat, but because a very low hydraulic conductivity of the confining bed and small head differences between the aquifers, the amount of leakage is very small.
-
REGIONAL GEOCHEMISTRY
Foster (1950) and Back (1960, 1966) illustrated how different cation
216
Fig. 6. Sodium concentration contours at 0.5-meq.11 intervals.
concentrations develop in a typical coastal-plain aquifer system. Back (1960) coined the term hydrochemical facies for this development and, in general, illustrated how they are developed in a system containing ion-exchange sites. Ca and Mg is dissolved from calcareous material (primarily invertebrate shells) by the action of carbonic acid generated in the atmospheric and soil zone. As the Ca2+and Mg2+ions are released t o the solution by dissolution they are exchanged for Na+ ions, which in a marine deposit, are initially present on the ion-exchange sites. This reaction yields a sodium bicarbonate water because the Ca2+and Mg2+,being doublecharged, are preferred on the exchange site to singlecharged Na+ ion. The HCO; ion is derived from the dissolution of calcareous materials as well as the reaction of atmospheric C 0 2 in the recharge area. If the amount of calcareous material available for weathering exceeds the ion-exchange capacity of the system, the Na-bicarbonate facies proceeds down the groundwater flow path with time. The Nabicarbonate facies is followed by a Ca-Mg-bicarbonate facies as the ion-
217
exchange sites are then filled with Ca2+and Mg2+ ions from the previous pore volume. If sufficient weathering occurs at the recharge zone t o remove all of the carbonate, the Ca-Mg-bicarbonate facies is followed by a dilute CaMg-K-bicarbonate facies that reflects weathering of the silicate framework. Different solutes in the water from the Aquia aquifer (Woll, 1978) were plotted on a map to determine their variability along the flow line. It was initially hoped that the conservative C1- ion, remaining from the marine origin of the formation, might be used as a tracer but the data show a nearly constant concentration throughout the study area. Na+ concentration of the water showed the most uniforming distribution with the flow field and was contour on 0.5meqJl (Fig. 6). Ca and Mg are the “tracers” in the normal sense of adding a mass to an existing flow system, but because of the competition between Ca and Mg for the exchange site, a more consistent contour value is obtained by using the Na concentration. Analyses of the solutes showed that values of Ca and Mg were inversely proportional to the amount of Na present and were exceedingly small where the concentration of Na was greater than 5.0 meq./l. These analyses also indicated that chloride was less than 0.5 meq./l, suggesting that the original sodium chloride pore water had been completely displaced from the aquifer over the study area and that the Na concentration observed were from ion exchange and not residual seawater.
DISPERSIVITY
Using the concentration contour map (Fig. 6) and the potentiometric map (Fig. 5), four hypothetical flow lines, A , B , C and D ,were drawn indicating possible movement of water and solutes through the aquifer from a recharge point in western Prince George County. Imaginary columns, 1m x 1m, were conceptually created that followed the lines A , B , C and D. Lengths of the square columns were determined by the choice of starting and finishing contours. By creating these imaginary columns t o this dimension, calculations were simplified and boundary problems minimized. The top of the imaginary columns was taken t o be the 0.5-meq./l Na contour line on Fig. 6 as this appears t o be the highest point in the column at which some of the original Na remains on the exchange site. Concentration of Na less than 0.5meq./l probably reflects input from precipitation atmospheric in the coastal environment and not ion exchange. Any contour greater than 0.5 meq./l could have been chosen as the top of the column but it was considered desirable to make the imaginary column as long as possible t o integrate large sections of the aquifer, Measurement of the distances between contours along the lines A , B , C and D ,by scaling directly from Fig. 6, gives the length of the column segments between successive contours. By assigning a 1-m2 cross-section, a bulk volume to the imaginary columns can be calculated. Because of the
218
granular nature of the aquifer, effective porosity was estimated t o be 0.35, however, this value has no effect on calculations. Using the effective porosity and bulk volume, a fluid volume was calculated relative t o the top of the imaginary column and each concentration contour. For example, it is 7300 m from the 0.5- to the 1.0-meq./l contour on line A , therefore, 1m x 1m x 7300m x 0.35 equals 2555m3 of fluid volume results for this part of line A . It is an additional 1140m from the 1.0- t o the 1.5-meqJlNa contour yielding 1m x 1m x 8440 m x 0.35 equals 2954 m3 of fluid volume. This calculation is performed on all of the Na contours crossed by the column. Darcy’s law was applied t o calculate the average volume of water that has moved through the imaginary column over the period of time of interest. The distance from the recharge point to the highest Na concentration contour was used in this analysis. An average transmissivity for each imaginary column was determined using Fig. 4 by measuring the length of the column between successive transmissivity contours and multiplying this by the average of the two contours, summing the total and dividing by the number of contours crossed. An average thickness was determined in the same manner using Fig. 3. Using these two values, an average hydraulic conductivity in meters per day was calculated for each imaginary column: A , 1.44; B, 1.44; C, 1.53; and D ;1.58m/day. The hydraulic gradient for each column was determined from Fig. 5. The gradient, hydraulic conductivity and area of l m 2 of the imaginary column were used t o compute the volume of water per unit of time that has moved through the column. One million years was chosen as the time when fresh water was introduced into the aquifer. This choice of time, however, does not affect calculations of dispersivity. The number of pore values (PV) that have passed through each contour line was calculated using the formula: PV =
(total flow) - (interstitial column space) (interstitial column space)
For example, total flow for lo6 yr. along line A was 2.55 l o 5 m 3 , the interstitial column space from the top of the column 0.5 meq./l t o 1.O meq./l was 1m x 1m x 7300 m x 0.35 or 2555 m3. Therefore, 98.8 pore volumes of water have passed this point in the last million years. The pore volumes for each concentration contour line crossed were calculated and normalized, assuming that the pore volumes passing the largest Na contour value in each column was equal to unity. Normalizing was done for ease of plotting and clarity of presentation, and does nothing to alter the dispersivity value. The concentration of output, Co , divided by concentration of input, CI (Co /CI), in which the value of Co is represented by the individual Na contours and C, is determined from the largest Na contour value, was calculated for each contour crossed by the imaginary column. For example, the Co /C, ratio in column A for the first contour is = 0.5/4.0 = 0.125, for the second contour it is 1.0/4.0 = 0.25, etc. The Co/CI-values were plotted vs. their corresponding normalized pore
219
U l0 0 L
z 206PP
tt
_ _ _ Observed
-
*+
calculated
ELK
E04-
u u 0 z z 0
Data points for observed values
~
u u 02-
I
velocity = 0 2557 m/year column length = 20888 m porosity = 0 35
1
- 3
0
2
I
4
3 PORE
5
6
7
VOLUMES
_ _ - Observed -Calculated Doto points for observed values velocity = 0 0.1957 1957 m/year column length = 26032 m porosity = 0.35 035
i:I"
00 0
2 2
I
3
PORE
4
6 6
5
7
VOLUMES
-__
Observed
calculated Data points for observed values velocity = column length = porosity = 0 0
I
2
4
3
PORE
VOLUMES
Fig. 7A-C. For caption see p . 220.
5
0 1314 m/year
53149 m 0 35
I
6
7
220
Fig. 7 . Dispersivity values in meters, for imaginary columns: ( A ) column A ;( B ) column B ; (C) column C ; and (D) column D.
volumes for all four imaginary columns in Fig. 7. Trial-valuesof dispersivity were used with the calculated flux, effective porosity and column length, using Ogata and Banks’ (1961) one-dimension analytical solution of the dispersion equation:
c0/cI=
$[erfc@(X- u t ) ( ~ , t ) - ” ~+ ) exp(uX/D) erfc{$(X + ~ t ) ( D L t ) - ” ~ ) l
where Co = concentration of output (M/L3); CI = concentration of input (M/L3 ); erfc = complementary error function (1- erf); X = distance (L), t = time (T); u = velocity (L/T); and D, = coefficient of longitudinal hydrodynamic dispersion ( L2 /T). For the imaginary column A (Fig. 7A), a dispersivity value of 5800 m coincided almost exactly with the observed data. For column B (Fig. 7B),the calculated value of 25,000 m is very close to the observed value. For column C (Fig. 7C),which exhibits the poorest fit of calculated t o observed values of the four imaginary columns, a value of -40,000m is close t o that of the observed line. For column D (Fig. 7D), a value of 38,000 m gives a good fit to the observed data. In general, the observed curves have the expected theoretical shape and show excellent fit. However, one outlying point on both imaginary columns C and D has been ignored. In both cases it is the uppermost Na contour line and indicates that perhaps we should have selected the 1.O-meq./l rather than the 0.5-meq./l Na contour line as the point at which ion exchange dominates the chemical reactions. When this substitution is made, the outlying points are completely eliminated and 2 smooth curve results. The calculated values of hydrodynamic dispersion are certainly larger than the true value because of the effect of non-homogeneous distribution of exchange sites. That is, the values represent an effective dispersivity rather than an absolute value, which is discussed in the next section.
221 DISCUSSION
An attempt has been made t o develop a practical methodology for determining dispersivity on a regional scale. The approach is best used with a conservative ion such as Cl; or isotopes such as *H and l80,as it essentially eliminates the need to consider chemical reaction. However, the author has attempted to show that even a highly reactive ion like Na+ may be used to gain at least some information on the effective dispersivity of the system. This brings us to the point of discussing some of the assumptions made when Na’ or another nonconservative ion is used with the method. The first of these assumptions made in the above analyses was that ion exchange is the only significant chemical mechanism controlling the Na concentration in this system. If only ion exchange and not mineral precipitation or another kinetically slow reaction controls the concentration, then a potential problem of reaction kinetics can be essentially eliminated. That is, ionexchange reactions are very rapid, usually reaching equilibrium in a few minutes (see, e.g., Malcolm and Kennedy, 1970; Wood, 1973). Thus if the reactions are faster than the pore-fluid velocity, then kinetics of reaction are not important in analyses of dispersivity. If kinetic reactions were controlling the concentration, then the observed distribution could be explained by the assumption of plug flow (i.e. no dispersivity), and kineticallycontrolled reactions. The demonstration that ion exchange is controlling the Na concentration in this system is therefore critical, for the use of the method, for nonconservative ions. Several approaches were used t o evaluate ion exchange. The first and most qualitative was to look at the published analyses. This indicated that the parameters commonly associated with ion-exchange reactions in natural systems were responding as though ion exchange were occurring. That is the concentration of Ca, Mg and K were inversely proportional to the Na concentration, the HC03 concentration increased with increasing Na concentration and the pH was higher in the areas with the higher Na concentration. The large quantities of zeolites in the formation were qualitative evidence that ion exchange was likely t o occur, and the lack of C1 ion indicated that the Na was not derived from residual seawater. Having demonstrated that ion exchange was at least a possible reaction, it was necessary to demonstrate the other reactions were unlikely. Fifteen representative analyses from different positions in the flow field were evaluated for thermodynamic mineral equilibrium, using sOLMNEQ , a mineral equilibrium program developed by Kharaka and Barnes (1973). These analyses indicate that all common Na minerals were thermodynamically unsaturated. That is, mineral solution could, if the minerals were present, occur, but that mineral precipitation was unlikely. This demonstrates that kinetically-controlled mineral precipitation is not controlling the concentration but this says nothing about the mineral’s solution kinetics. However, it has been the author’s observation that minerals equilibrium by solution is
222
reached very rapidly in other clastic aquifers, for example, the Saginaw of Michigan, the Ogallala and Carrizo of Texas and the Fox Hills of North Dakota. Therefore, by analogy it is assumed that mineral solution is not controlling the Na concentration in this aquifer. Another factor that could be controlling the concentration distribution was that of solutes moving into the Aquia from adjacent aquifers. Without isotopic analyses it is impossible t o prove that leakage from the Namjemoy is not occurring. However, the distribution of Na in the Namjemoy is different from that of the Aquia and the pattern of Na concentrated (Fig. 6) in the Aquia suggests that it is the flow within the aquifer that is controlling the concentration. These observations, coupled with the low head difference, low chloride concentration and low hydraulic conductivity of the confining bed, suggest that mixing from adjacent aquifer is small, however, it does not unequivocally eliminate it as a possibility. There is no evidence of that ion filtration had any significant effect on this distribution on Na in this system. Based on the approach outlined above it was believed that the major chemical mechanism controlling Na concentration in the Aquia aquifer is ion exchange and, consequently, the distribution of Na concentration is not kinetically controlled. An implied assumption associated with ion exchange is that the ions on the exchange sites are at equilibrium with those in solution. This equilibrium is well documented and forms the basis for calculating the so-called “distribution coefficients”, “separation factors” and “selectivity” used in simulating ion-exchange system (see, e.g., Helfferich, 1962; Amphlett, 1964). The critical assumption that the exchange sites are uniformly distributed is almost certainly invalid. The values of dispersives range from 5600 t o 40,000 m, very much larger than those commonly reported in the literature. This may be a factor of the larger distance, of -100 km, used in this study, with a consequent opportunity for large variation in hydraulic conductivity, or a function of the distribution of exchange sites. We have assumed a homogeneous distribution of exchange sites in the aquifer and this is very unlikely in an aquifer that exhibits three stratigraphically identifiable sedimentary facies. That is, the distribution of concentration contours will be related t o the location of the ionexchange sites, as well as the dispersivity. If, for example, the ion-exchange sites are not homogeneously distributed, a “pileup” of concentration contours of the exchanging tracer will occur in response to the exchanger and not the flow field. How then does the failure to have homogeneous reaction sites affect the result of the analysis? The answer is that it fails to separate hydrodynamic dispersion from the nonhomogeneous distribution of exchange site, thus making these dispersivities effective and not absolute values. However, it has no effect on the methodology which is the main thrust of this paper. SUMMARY AND CONCLUSION
The proposed method of evaluating dispersivity views a regional aquifer
223
system as a large column in which concentration contours in space have been interchanged for concentration variation with time. Curves relating concentration contour and pore volume of fluids are compared to theoretically developed curves to determine dispersivity. The method was developed for use with a conservative solute or isotope and was extended t o include reacting solutes. However, the use of reacting solutes limits the separation of mechanical mixing from that caused by inhomogeneity of reaction sites. As a consequence effective dispersions rather than absolute values are obtained. The use of a conservative solute would yield a true value of dispersivity. ACKNOWLEDGEMENTS
The author wishes to thank Dr. William Back of the U.S. Geological Survey for his essential input into this investigation. Dr. Back not only contributed his time in discussion, review and suggestion of a suitable aquifer system to test these concepts, but he also gave his continuing encouragement and support. The author also thanks Ren Jen Sun for critical review of the manuscript, Walter F. White and Fred K. Mack for discussion on the geohydrology of Maryland Coastal Plain, and Ms. Leslie Adams and Robert Suddarth for help in preparing the illustrations. REFERENCES Back, W., 1960. Origin of hydrochemical facies of groundwater in the Atlantic Coastal Plain. Int. Geol. Congr., 21st Sess., Copenhagen, Part I, pp. 87-95. Back, W., 1966. Hydrochemical facies and groundwater flow patterns in northern part of Atlantic Coastal Plain. U S . Geol. Surv., Prof. Pap. 498-A, 42 pp. Bredehoeft, J.D., Counts, H.B., Robson, S.G. and Robertson, J.B., 1976. Solute transport in groundwater systems. In: J.C. Rodda (Editor), Facets of Hydrology. Wiley, New York, N.Y., pp. 229-256. Ferguson, H.F. and Martin, R.O.R., 1953. The water resources of St. Mary’s County. Md. Dep. Geol., Mines Water Resour. Bull. No. 11,195 pp. Foster, M.D., 1950. The origin of high sodium bicarbonate waters in the Atlantic and Gulf Coastal Plains. Geochim. Cosmochim. Acta, 1:33-48. Glaser, J.D., 1971. Geology and mineral resources of southern Maryland. Md. Geol. Surv., Rep. Invest. No. 15, 8 5 pp. Hansen, J.J., 1974. Sedimentary facies of the Aquia Formation occurring in the subsurface of the Maryland Coastal Plain. Md. Geol. Surv., Rep. Invest. No. 21, 47 pp. Helfferich, F., 1962. Ion Exchange. McGraw-Hill, New York, N.Y., 624 pp. Jacobeen, F.H., 1972. Seismic evidence for high angle reverse faulting in the Coastal Plain of Prince Georges and Charles County, Maryland. Md. Geol. Surv., Info. Circ. No. 13, 21 PP. Kapple, G.W. and Hansen, H.J., 1976. A digital simulation model of the Aquia aquifer in southern Maryland. Md. Geol. Surv., Info. Circ. No. 20, 33 pp. Kharaka, Y.K. and Barnes, I., 1973. SOLMNEQ: solution-mineral equilibrium computation. Natl. Tech. Info. Serv., Springfield, Va., Tech. Rep. PB 214-899, 82 pp. Mack, F.K., 1962. Ground-water supplies for industrial and urban development in Anne Arundel County, Maryland. Md. Geol. Surv., Bull. 2 6 , 9 0 pp.
224 Mack, F.K., 1966. Ground-water in Prince Georges County, Maryland. Md. Geol. Surv., Bull. 29, 1 0 1 pp. Mack, F.K., Webb, W.E. and Gardner, R.A., 1 9 7 1 . Water resources of Dorchester and Talbot Counties, Maryland - with special emphasis on the ground-water potential of the Cambridge and Easton areas. Md. Geol. Surv., Rep. Invest. No. 1 7 , 1 0 7 pp. Malcolm, R.L. and Kennedy, V.C., 1 9 6 9 . Rate of cation exchange on clay minerals as determined by specific-ion electrode techniques. Soil Sci. SOC. Am. Proc. 33( 2 ) : 2 4 7-2 5 3. McRae, S.G., 1972. Glauconite. Earth Sci. Rev., 8 ( 4 ) : 397-440. Ogata, A. and Banks, R.B., 1 9 6 1 . A solution of the differential equation of longitudinal dispersion in porous media. U.S. Geol. Surv., Prof. Pap. 411-A, 7 pp. Otton, E.G., 1955. Ground-water resources of the southern Maryland Coastal Plain. Md. Dep. Geol., Mines Water Resour. Bull. No. 1 5 , 247 pp. Overbeck, R.M., and Slaughter, T.H., 1 9 5 8 . The water resources of Cecil, Kent and Queen Anne's Counties, Maryland. Md. Dep. Geol., Mines Water Resour. Bull. No. 2 1 , 4 7 8 pp. Rasmussen, W.C. and Slaughter, T.H., 1 9 5 7 . The water resources of Caroline, Dorchester and Talbot Counties, Maryland. Md. Dep. Geol., Mines Water Resour. Bull. No. 1 8 , 4 6 5 pp. Schwartz, F.W., 1 9 7 7 . Macroscopic dispersion in porous media: the controlling factors. Water Resour. Res., 1 3 ( 4 ) : 743-752. Schwartz, F.W. and Muehlenbachs, K., 1 9 7 9 . Isotope and ion geochemistry of groundwaters in the Milk River aquifer, Alberta. Water Resour. Res., 1 5 ( 2 ) : 259-268. Smith, L. and Schwartz, F.W., 1 9 8 0 . Mass transport, 1. A stochastic analysis of macroscopic dispersion. Water Resour. Res., 1 6 ( 2 ) : 303-313. Weigle, J.M., Webb, W.E. and Gardner, R.A., 1 9 7 0 . Water resources of southern Maryland. U.S. Geol. Surv., Hydrol. Invest. Atlas, No. HA-365, 3 sheets. Winograd, I.J. and Pearson, Jr., F.J., 1 9 7 6 . Major carbon-14 anomaly in a regional carbonate aquifer. Geol. SOC.Am., Annu. Meet., Abstr. Progr., pp. 1008-1009. Winograd, I.J., and Pearson, Jr., F.J., 1 9 7 6 . Major carbon-14 anomaly in a regional carbonate aquifer: possible evidence for megascale channeling, south central Great Basin. Water Resour. Res., 1 2 ( 6 ) : 1125-1143. Woll, R.S., 1 9 7 8 . Maryland ground-water information: chemical quality data. Md. Geol. Surv., Basic Data Rep. No. 1 0 , 1 2 5 pp. Wood, W.W., 1 9 7 3 . Rapid reaction rates between water and a calcareous clay as observed by specific-ion electrodes. J. Res. U.S. Geol. Surv., l ( 2 ) : 237-241.
225
FLOW-SYSTEM CONTROLS OF THE CHEMICAL EVOLUTION OF GROUNDWATER F.W. SCHWARTZ' , K. MUEHLENBACHS' and D.W. CHORLEY'
'
Department of Geology, University of Alberta, Edmonton, Alta. T6G 2E3 (Canada) 2Sirnco Groundwater Research L t d . , Vancouver, B.C. V 6 R 2 K 3 (Canada) (Accepted for publication February 26,1981)
ABSTRACT Schwartz, F.W., Muehlenbachs, K. and Chorley, D.W., 1981. Flowsystem controls on the chemical evolution of groundwater. In: W. Back and R. Lktolle (Guest-Editors), Symposium on Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54: 225-243. Isotopic and major-ion analyses of 130 fresh and brackish groundwater samples reveal a strikingly consistent pattern of variation over 28,000 km2 of the western Canada sedimentary basin. Hydrodynamic interpretations based on drill-stem tests and piezometric data reveal a pattern of broad, regional flow from south to north. However, there is some evidence t o suggest that patterns of groundwater flow have been influenced in the past by Wisconsin glaciation. All the permeable units appear t o be recharged by meteoric water in the south where they outcrop or come close t o the surface. Samples of groundwater collected down dip in each of six major sandstone units are progressively enriched in D, l8O, Na+ and C1-. For example, the deepest waters of the artesian Milk River aquifer are enriched by up to 70YoO and 15Yw with respect t o the 6D and 6 ' * 0 of the modern recharge waters. The pattern of chemical evolution is strongly related t o the hydraulic characteristics of individual units. The isotopic composition is determined by the extent t o which hydraulic conductivity has facilitated meteoric water flushing of connate water. Thus, more permeable units and permeable zones within a unit tend to be isotopically lighter because of a more complete and more rapid invasion of meteoric water. Interestingly, the largest conductivity values are found in some of the deepest units (1500 m). Consequently, water in these deeper units have 6D- and 6"O-values which approach that of the recharge. The major-ion chemistry is controlled both by this process and a complex set of rock-water interactions. In addition t o providing information about the chemical evolution of groundwater, this study can begin t o quantify the complex pattern of flushing in a large sedimentary basin.
INTRODUCTION
The patterns of flow and chemical evolution of groundwater in the western Canada sedimentary basin have been studied intensively since the early 1960's. The pioneering work in establishing the character of groundwater flow was that of Hitchon (196Oa'b). The existence of basin-wide flow systems and the profound influence of topography and geology on the patterns
226
of flow were firmly established in those two papers. At the same time, Hitchon and Friedman (1969) completed a study on the origin of variability in the stable-isotope ratios of formation water in the basin. That work indicated the importance of mixing, rock-water interactions, and ultra-filtration in controlling the isotopic character of formation water. Our work in this basin has been concentrated in a much smaller area, north of the Alberta-Montana border. The initial research was designed to investigate the oxygen and deuterium content and major-ion chemistry of groundwater from the Milk River aquifer system (Schwartz and Muehlenbachs, 1979). This very extensive unit is an important artesian aquifer. It is recharged in outcrop areas in the southern part af the study area. Flow is mainly northward down the regional dip. The northern portion of the unit is found at depths greater than 350 m and contains very substantial quantities of natural gas. The very striking variability in 6D and 6 l 8 0 of groundwater in the Milk River aquifer is surprising in such an active, near surface flow system. “0 and D contents of water from the recharge end of the system fall close t o the meteoric water line with 6D = -155yoO and 6 l 8 0 = -20°/00 SMOW. Waters from the northern end of the aquifer are enriched by -12yo0 and 70%0 in “0 and D, respectively, and exhibit a systematic deviation from the meteoric water line. The existence of such a marked and regular variation in water chemistry and some of the unresolved problems in explaining the origin of these patterns naturally lead us to an examination of the water chemistry of other permeable units deeper in the stratigraphic section. Therefore, the first objective of this paper is t o describe, in detail, features of the hydrogeologic setting and geochemistry of units in the upper 1500m of the sedimentary basin. The second objective is t o explain how features of the flow systems within the basin can exert a control on the isotope and major-ion chemistry of the groundwater system.
GEOLOGY
The study area comprises -28,000 km2 in southeastern Alberta (Fig. 1). Rocks of interest to this study range in age from Tertiary to Carboniferous. The stratigraphic sequence, however, does contain older Paleozoic ,and Precambrian rocks at depth. With the exception of certain carbonate rocks in the Upper Paleozoic, the dominant lithology is shale with occasional sandstone units and minor coal. Some of the sandstone units contain substantial quantities of oil and gas. As a result of continued exploration for and development of these resources, a tremendous quantity of conventional lithologic log, geophysical log, core and drill-stem test data is available for this area. The essential features of the regional geology are depicted in a series of cross-sections (Fig. 2). Each of the sections is prepared from structure con-
227
Fig. 1 . Location of the study area.
tour maps which we constructed for the top of each of the main geologic units. Fig. 2 shows that beds are dipping northward in what is essentially a broad plunging anticlinorium. The purpose of the idealized stratigraphic section added to the figure is not only to provide additional detail in the lithology but also to define several of the thin sandstone units that could not be shown on the cross-sections. These sandstone units have been tested and sampled extensively mainly by conventional oil-field methods. The oldest unit on the sections is the Livingstone Formation, which is Mississippian in age. It is a medium-to-coarse crystalline limestone partly altered to dolomite (McCrossan and Glaister, 1964). It is unconformably overlain by the Ellis Group of Jurassic age, which is comprised of three formations. They are, in ascending order, the Sawtooth, Rierdon and Swift formations. The Sawtooth Formation is a fine- t o medium-grained sandstone with interbeds of shale and limestone (McCrossan and Glaister, 1964). The Rierdon Formation consists of thin interbeds of calcareous shale and argillaceous limestone. The Swift Formation, which consists of shale grading upwards t o fine quartzose sands, is limited to the southeast corner of the study area. Unconformably overlying the Ellis Group is the Mannville Group. I t consists of an interbedded sequence of sandstone, shale and siltstone with a few coal beds. The Sunburst, Cutbank and Glauconitic sandstones are the most important shaly sandstone formations of the Lower Mannville (Fig. 2). The overyling Colorado Group, which is divided into an upper and lower part, is predominantly a sequence of shale. However, four important sandstone units
Scale so0
,
0 - 0
100 I
229
are defined. The Bow Island Formation consists of interbedded sandstone and shale. Because this unit contains large quantities of oil and/or gas, there is a large quantity of available geological and geophysical data. Other sandstone units are the Fish Scale Sandstone, the Second White Specks and the Medicine Hat Sandstone. The Milk River Formation conformably overlies the Colorado Group and consists of up to 100 m of very fine- t o medium-grained sandstone, siltstone and shale. More detailed geological descriptions of the aquifer are presented by Meyboom (1960). The Milk River aquifer, which comprises the lower part of the formation, crops out in the southernmost part of Alberta and northern Montana. Overlying the Milk River Formation is the Pakowki Shale. This unit consists entirely of marine shale with some very thin sandstone and bentonite beds (McCrossan and Glaister, 1964). Other bedrock and glacial drift units overlying the Pakowki have not been further subdivided (Fig. 2).
HYDROGEOLOGY
Sufficient information is available t o develop a reasonably accurate picture of hydraulic-conductivity and hydraulic-head distributions in the non-shale parts of the stratigraphic sequence. Fig. 3 illustrates the spatial variation in hydraulic-conductivity values determined from 590 drill-stem tests and 35 aquifer tests from the southern portion of the Milk River aquifer. Basic drillstem test data were obtained from the Energy Resources Conservation Board of Alberta and were interpreted using standard methods (Earlougher, 1977). Permeability values are calculated from a knowledge of flow rate, formation volume factor, viscosity, packer spacing and the change in shut-in pressure per log-cycle. For comparative purposes, The permeability values are converted to hydraulic-conductivity values assuming that the pore fluid is water at 15.6"C. In each unit, hydraulic-conductivity values typically vary over a broad range. Some of the highest values are obtained for the Mannville Group, and the Rierdon and Sawtooth formations of the Ellis Group, In this area, these more permeable units are found relatively deep in the stratigraphic sequence. Unfortunately, no data are available for the intervening shale units. However, by analogy with conductivity values reported elsewhere for shale and clay sediments (C.D.W.R., 1971), they should be lower than 5 - 1 0 - 6 cm/s. Some preliminary work has been completed to explain the origin of the permeability of the sandstone units. In the Bow Island Formation, an attempt was made to correlate conductivity values from drill-stem tests with the estimated true formation resistivity determined from electric logs. The lack of correlation, however, suggests that fracture permeability may be prevalent. Comparison of the hydraulic-conductivity patterns for the Bow Island Formation, the Sunburst and Cutbank sandstones and the Sawtooth For-
-
Fig. 2. Three-dimensional representation of the near-surface geology and a stratigraphic column.
230 Milk River
-I
u u 20 _ Range
Range I
I
Range
Range I
Glauconitic
/
Sunburst and Cutbank
/ ,
10
20
5
10
Range
Range
I
Sawtooth
Mississippian
115
Range
Range
Fig. 3. Maps depicting the spatial variation in hydraulic conductivity.
231
r-
1
5
0
Range
1
L
Medicine Hat
Range
7
@ Second White Speckled Shale -4
a 10
s c
b
Range
Range
Sawtooth, Rierdon and Swift
Range
Range
Fig. 4. Maps depicting the spatial variation in hydraulic head.
232
mation (Fig. 3) indicates a coincident east-west trending zone of relatively high hydraulic conductivity through the middle of the study area. The presence of a similar zone in such a diverse set of units may suggest a broad fracture zone. The hydraulic-head distributions are determined mainly from extrapolated drill-stem pressure data (Earlougher, 1977). In all cases, test results have been rejected if the initial and final maximum reservoir pressures differed by more than 5%. Pressure determinations are expressed as equivalent freshwater hydraulic head in meters above sea level at 15.6'C. Maps of the potentiometric surfaces for eight major geological units have been prepared from this data set (Fig. 4). Because potentials may have been influenced by gas and oil production (or groundwater in the case of the Milk River aquifer), the location of major fields is indicated on the figure. A reasonably consistent set of flow conditions can be noted by comparing the flow directions indicated by the arrows on Fig. 4. Groundwater moves generally from south to north across the study area. These results are in agreement with the larger-scale model of Hitchon (1964) and confirm the existence of large groundwater flow systems operating in the study area. Detailed examination of data from Bow Island Formation and Sunburst and Cutbank sandstones shows how the patterns of flow within a unit are controlled in part by the local variation in hydraulic conductivity (cf. Figs. 3 and 4). For example, in the case of the Bow Island Formation, the tendency for groundwater to move in a northeasterly direction is probably related to the northeasterly-trending conductivity high. Evaluation of the vertical flow components is difficult because the head calculations from drill-stem test data probably have errors associated with them of the same order of magnitude as the actual differences in head. It is probably possible, however, to make certain generalizations. The flowingwells completed in the Milk River aquifer provide good evidence of an upward flow component from the aquifer toward the ground surface. I t is also possible that water is also leaving the aquifer from the bottom and moving downward. The Bow Island Formation appears to be a potential low with flow directed upward towards it from the Mannville and downward from the Second White Specks, Medicine Hat Sandstone and Milk River Formation.
GEOCHEMISTRY OF FORMATION WATERS
In this description of the geochemistry of formation water, we will consider first the distribution of Na+, C1- and SO$- ions which are some of the most dominant major-ion species and second the stable-isotope ratios of D/H and 180/'60. The basic data for the ion chemistry come from a set of 240 good-quality unpublished analyses of samples collected mainly from drill-stem tests but some from separators. These data are supplemented by -1500 analyses of formation water gathered over the last 20 years by the
233
Energy Resources Conservation Board of Alberta. An additional 93 goodquality analyses were available for well water collected from the Milk River aquifer. The distributions of Na+, C1- and SO:- ions are shown in Figs. 5-7 for the Milk River Formation, Bow Island Formation, Mannville Group and Sawtooth Formation. Because each of the maps is based on a relatively large number of data points, because the data are reasonably consistent in given areas of each map and because the trend of increasing concentrations is completely compatible with the hydrodynamic data, we believe that the maps accurately reflect the major variations in Na', C1- and SO:- ion concentrations. However, we had no control over the collection or analyses of these samples and as a result are not able t o evaluate more specifically the possibility for errors in the results. Na' ion concentrations in all formations show a systematic increase from south to north across the study area (Fig. 5). Values in the Milk River aquifer range from -500 to 2000 mg/l. In the Bow Island Formation and the Mannville Group, Na' concentrations are higher, ranging from -2000 t o more
-
- 15 - r
Range
Range
Fig. 5. Maps depicting Na+ ion concentrations of groundwaters.
l
234
r
I
Milk River
Il5 20
a
20
Range
15
10
5
Range -
20
I5
Sawtooth
.a c
-*O
I15
.a
than 8000 mg/l. Na' ion concentrations in the Sawtooth Formation may be somewhat higher (Fig. 5). Concentrations of C1- in the Bow Island Formation reach a maximum of 14,00Omg/l; while, those from the Mannville reach a maximum of 20,000 mg/l (Fig. 6). For SO:- ion, patterns are more complex. Typically, values for the Milk River aquifer are less than 100 mg/l except in areas along the edge of the aquifer (Fig. 7). Concentrations in the Bow Island Formation are typically less than 50 mg/l except togthe west. A similar pattern is evident within the Mannville Group with values t o the west greater than 100 mg/l (in a range from 100 to 800 mg/l) and those t o the east less than 50 mg/l (with the exceptions indicated). In the Sawtooth Formation SO:- ion concentrations are variable making it very difficult to subdivide regions where the concentration is greater than 50 mg/l. The Sl8O- and SD-values were determined for 118 formation water samples from various units in the stratigraphic section. Most of the 39 samples from the Milk River were collected by the authors from water wells. The
-
235 I
u
1
Milk River
20
15
10
1
L/
Bow Island
2o
5
20
Range
Range
25
20
15
.a
r
u)
10
f
0 I-
5
20
Range
15
10
5
Range
Fig. 7 . Maps depicting SO$- ion concentrations of groundwaters.
remainder are samples from drill-stem tests, which were collected in a detailed program of studies sponsored by the Alberta Research Council and the Energy Resources Conservation Board of Alberta. Major-ion analyses available for these samples made it possible t o compare results with existing major-ion data to assure that samples were not contaminated by drilling mud or other fluids. Evidence that the major-ion chemistry of nearly all the samples is appropriate for the formation from which they were collected, and that patterns of spatial variability in isotope chemistry are consistent within formations leads us to believe that the samples are representative of the formation water. Oxygen-isotopic ratios were determined by the C 0 2 equilibrium method of Epstein and Mayeda (1953). The hydrogen-isotopic ratios of water were determined by the method of Friedman and Woodcock (1957). The data are reported in the usual delta notation with SMOW as the standard (Craig, 1961). Our analyses of I.A.E.A. SLAP was -424.6yoO for 6D and -55.95%0
236
for 6 l 8 0 . The standard deviation calculated from the pooled residual variance of duplicate analyses was +2.3YO0for 6D and rtr0.13%, for also. For those geological units for which sufficient samples are available (i.e. Milk River Fm., Bow Island Fm., Sunburst Sandstone and Sawtooth Fm.), it is possible to describe systematic variations in 6D and 6l8O (Figs. 8 and 9). The most marked variations are observed for the Milk River aquifer where 6D-values range from -150?00 to greater than - 9 0 ~ o o ,and 6180-values range from less than -20?00 to greater than -7.5%, . Formation water from
110
5
s
0 I-
5
20
15
10
Range
5
.-Q c 0 I-
I 10
Range
5
c 5
237 I
Milk River
Sunburst ]I5
\
Sawtooth
Range
1
l5
Range
Range Fig. 9. Maps depicting variations in 6l80 (ym).
the Bow Island Formation generally falls within a much narrower range. D contents have a maximum variation of "25yoO and "0 contents have a maximum variation of -5yo0. The range in variation for the Sunburst Sandstone is almost identical t o that for the Bow Island (Figs. 8 and 9). In the Sawtooth Formation, GD-values vary from -125 t o -77%0 with many of the values less than -lOOyoo. "0-values range from ----16 to -5.5YoO. Determinations for the Mississippian and Devonian units represent some of the lowest GD- and G1'O-values for water from units deeper in the section than the Milk River aquifer.
-
238
68t
2
Milk River
Mississippian -160
+ Devonian
I:.i,rmAH
-140
-120
-100
I
-80
I
,
-60
6 D (%o)
Fig. 10. Summary diagram comparing 6 D (Too) content of groundwater.
Fig. 10 presents a more simplified description of the variability in 6D among the various geological units. Note how water from the Bow Island Formation and Sunburst Sandstone is enriched in D relative to the other units. In addition, it is evident that the range in values is quite variable from formation to formation.
DISCUSSIONS AND CONCLUSIONS
The hydrodynamic data indicate that the generalized direction of flow in almost all units is from south to north. Recharge to these systems is probably occurring in Montana where units in the study area or their stratigraphic equivalents outcrop (subcrop). The most important structural feature is the Little Belt Uplift which brings Precambrian rocks to the surface. The character of this uplift is such that hydraulic continuity is maintained along individual stratigraphic units dipping northward into the western Canada sedimentary basin. In the uplift, which formed the Rocky Mountains to the west, hydraulic continuity has been disrupted by a complex sequence of overthrust faulting. This difference in the character of the uplifts probably explains why the n o r t h s o u t h flow systems have developed in southern Alberta instead of east-west ones controlled by nearby mountain ranges t o the west. Because of the lateral continuity of individual permeable units and
239
the relatively low permeability of intervening units, it is probable that the quantity of cross-formational flow occurring is relatively small. Evidence presented by Hitchon (1969a, b) suggests that many of these large basin-wide flow systems probably discharge in northeastern Alberta, -550 km north of the study area. One important question t o be considered is whether flow conditions that exist now were the same over the last l o 5 yr. Toth (1978) has suggested that temporal changes in the character of bedrock flow systems in Alberta should be expected. Much of the early Middle and Late Wisconsin from -0.12 .lo5 t o l o 5 yr. B.P. was a time of active glaciation particularly in the northern half of the province where ice was present much of the time. There actually has been only a relatively short time period (-0.12 -10’ yr.) since glacial ice has been totally absent from Alberta. Over the last -0.7 *lo5yr. (Middle and Late Wisconsin), glacial ice advanced only as far as the northernmost half t o two-thirds of the study area (Rutter, 1980). Thus, the presence of ice and the location of the ice-front may have been a significant factor in controlling groundwater flow in the southernmost portions of Alberta. With glacial ice covering the discharge end of the flow system during much of the last l o 5 yr. and hydraulic potential available south of the ice-front t o cause flow, it may be expected that groundwater would be forced t o discharge in front of the ice in southern Alberta. These flow conditions would result in more pronounced cross-formational flow with deeper formation water moving into shallower units. This long-term change in flow conditions could explain why the Milk River aquifer in downdip locations contains water which is not meteoric and which is isotopically similar t o water from the Bow Island Formation (Fig. 11).If flow conditions as they exist today were maintained for lo5 yr. or more, it is difficult t o conceive why the Milk River aquifer had not been completely flushed by meteoric water. As the various tests of hydraulic conductivity show, the Milk River aquifer system is a relatively permeable and active flow system. In contrast t o the striking patterns in the Milk River aquifer, only meteoric water is found in the Madison Limestone over parts of Wyoming, Montana, South Dakota and North Dakota (Hanshaw et al., 1978). This flow system is analagous t o the Milk River in terms of its wide extent, relatively high hydraulic conductivity and depth below ground surface. The variation in isotopic values for water from the Milk River aquifer has been attributed t o macroscopic dispersion (Schwartz and Muehlenbachs, 1979). In this process, meteoric water with an isotopic composition similar t o present-day recharge water moves along the formation t o displace preexisting formation water. This simple mixing process is complicated by the continued leakage of water from both the top and bottom of the Milk River Formation. Because this process probably operated for less than 0.15 -10’ yr. the flushing is incomplete and a zone of mixing still remains between isotopically very different waters. N o significant number of water samples from any formation except the
240
';j -80
-120
-140!- 2 5 -160
Fig. 11. Plots of 6 D
("/oo ) vs. 6 l80("/oo ) in groundwater from the various units.
24 1
Milk River are similar isotopically t o present-day meteoric water. Water from deeper units is generally brackish and have 6D-values ranging from -70 to -120700 which suggests that they are the product of long-term mixing of meteoric and connate water. Hitchon and Friedman (1969) used this argument to explain the isotopic character of shallow formation water in the basin. Deviation from an idealized mixing line between meteoric water and connate water can probably be related t o the exchange of oxygen with carbonates of the reservoir rocks (Clayton et al., 1966). On a regional scale, there is an apparent correlation between the extent of mixing and the gross permeability of individual units. For example, Mississippian units and the Sawtooth Formation have a relatively high permeability as compared to the Bow Island Formation. Indications from the 6D data (Fig. 10) are that these deeper units also have a higher proportion of meteoric water. Thus, it appears that the pattern of basin flushing depends upon the broad-scale features of hydraulic conductivity. The higher conductivity elements of the flow system represent pathways for initial and continued introduction of meteoric water to the basin. This kind of mixing provides a good example of the process of megascopic dispersion. On a broad scale, the character of the groundwater patterns, which is related to the hydraulic conductivity, controls water chemistry through mixing. In addition, there is evidence that variability in hydraulic conductivity within a formation also provides an important control on the variation of 6l8O and 6D. The best example is that of the Milk River aquifer. Here, a northward protrusion of isotopically depleted and more dilute water coincides with the thickest and most transmissive part of the aquifer. There are indications that a similar control on chemistry might be occurring in other units, particularly the Bow Island Formation. However, because the 6D-values are very similar in this unit, these dispersion processes are not nearly as obvious. Look at Figs. 5 and 6 and notice how the distribution of Na+ and C1- ions appear to be related to the conductivity (Fig. 3). The band of relatively dilute water from southwest t o northeast across the area coincides with a permeability high. Examination of the Na', C1- and SO:- ion concentrations in relation to the deuterium data shows that their concentration is controlled by the geochemical evolution of water in the direction of flow. For example, in the Bow Island Formation the concentration of Na' and C1- increases markedly in the direction of flow whereas D concentrations remain within a narrow range. Similar relationships are evident in the Mannville units. Only in the Milk River Formation is there an apparent relationship between Na' and C1- ion concentration and 6D. However, our previous work has shown that the salinity of the water increases independently of D concentration once 6Dvalues are greater than - l O O ~ o o . In summary, both the potentiometric and geochemical data provide evidence of the existence of an extensive groundwater flow system operating
242
over many hundreds of kilometers. This flow system is recharged in Montana and flows down the formational dip northward into the western Canada sedimentary basin. We conclude that the character of this flow system probably changed in response t o continental glaciation over the past l o 5 yr. This process provides the only feasible way t o explain the presence, in an active groundwater flow system, of water isotopically enriched relative t o presentday meteoric water. The l80/l6O and D/H ratios of water is controlled by mixing processes related t o the flow conditions. We would conclude from this work that the isotopic characteristics of water provide the best way t o evaluate basin mixing processes. The Na', C1- and SO;- ion concentrations are controlled to a limited extent by mixing processes but more importantly by ongoing geochemical processes. ACKNOWLEDGMENTS
This work was supported by the Natural Sciences and Engineering Research Council of Canada. The authors thank Brian Hitchon for providing samples and unpublished chemical analyses. Mrs E. Toth ably assisted in the laboratory. REFERENCES C.D.W.R. (California Department of Water Resources), 1971. Sea water intrusion: Aquitards in the coastal groundwater basin of Oxnard Plain, Ventura County. Calif. Dep. Water Resour. Bull. 63-4, 569 pp. 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, 1. Isotopic composition. J. Geophys. Res., 71: 3869-3882. Craig, H., 1961. Standards for reporting concentrations of deuterium and oxygen-18 in natural waters. Science, 133: 1833-1834. Earlougher, R.C., Jr., 1977. Advances in well test analysis. Soc. Pet. Eng., Am. Inst. Metall. Pet. Eng. (A.I.M.E.), Monogr. Vol. 5, 264 pp. Epstein, S. and Mayeda, T., 1953. Variation of 0-18 content of waters from natural sources. Geochim. Cosmochim. Acta, 4: 89-103. Friedman, I. and Woodcock, A.H., 1957. Determination of deuterium hydrogen ratios in Hawaiian waters. Tellus, 9: 553-556. Hanshaw, B.B., Busby, J. and Lee, R., 1978. Geochemical Aspects of the Madison Aquifer System. In: Williston Basin Symposium. Mont. Geol. SOC.,385-389. Hitchon, B., 1964. Formation fluids. In: Geological History of Western Canada, Ch. 15. Alta. SOC.Pet. Geol., Calgary, Alta. Hitchon, B., 1969a. Fluid flow in the western Canada sedimentary basin, 1. Effect of topography. Water Resour. Res., 5( 1):186-195. Hitchon, B., 1969b. Fluid flow in the western Canada sedimentary basin, 2. Effect of geology. Water Resour. Res., 5(2): 460-469. Hitchon, B. and Friedman, I., 1969. Geochemistry and origin of formation waters in the western Canada sedimentary basin, 1. Stable isotopes of hydrogen and oxygen. Geochim. Cosmochim. Acta, 33: 1321-1349.
24 3 McCrossan, R.G. and Glaister, R.P. (Editors), 1964. Geological History of Western Canada. Alberta Society of Petroleum Geologists, Calgary, Alta., 230 pp. Meyboom, P., 1960. Geolgoy and groundwater resources of the Milk River sandstone in southern Albrta. Alta. Res. Counc., Mem. 2, 89 pp. Rutter, N.W., 1980. Late Pleistocene history of the western Canadian ice-free corridor. Can. J. Anthropol., 1: 1-8. Schwartz, F.W. and Muehlenbachs, K., 1979. Isotope and ion geochemistry of groundwaters in the Milk River aquifer, Alberta. Water Resour. Res., 15(2): 259-268. Toth, J., 1978. Gravity-induced cross-formational flow of formation fluids, Red Earth region, Alberta Canada: Analysis, patterns, evolution. Water Resour. Res., 14(5): 805-843.
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245
CHEMICAL EVOLUTION OF GROUNDWATER IN A DRAINAGE BASIN OF HOLOCENE AGE, EAST-CENTRAL ALBERTA, CANADA
E.I. WALLICK
Groundwater Department, Alberta Research Council, Edmonton, Alta. T6H 5R7 (Canada) (Accepted for publication April 16, 1981)
ABSTRACT Wallick, E.I., 1981. Chemical evolution of groundwater in a drainage basin of Holocene age, east-central Alberta, Canada. In: W. Back and R. L6tolle (Guest-Editors), Symposium of Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54: 245-283. Chemical evolution of groundwater in a small drainage basin of glacial origin (10,250 yr. B.P., based on radiocarbon age dating of gyttja from a closed saline lake in the basin) was studied in order t o understand the generation of salts in surface-mined areas on the interior plains of Alberta. The basin was considered to be a natural analogue of a surfacedisturbed area because of the large volumes of rock that had been redistributed by glaciers with the resulting change in topography and drainage. The distributions of hydraulic head, total dissolved solids (TDS), and environmental isotopes essentially reflect the superimposition of groundwater flow systems associated with the post-glacial topography upon a regional bedrock flow system of older but undetermined age. In the glacial drift aquifers and aquitards (sands and till), the groundwater composition was typically Ca-Mg-bicarbonate type at depths less than 30 m, but at depths of 30-100 m, the composition was Na-bicarbonatesulfate type. In the deeper bedrock aquifers ( > 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 systems
256
Primary and secondary mineralogy of the Belly River Formation In view of the above considerations, the main primary and secondary minerals of the Belly River Formation may be listed. The primary minerals, that is, those that were probably present upon deposition, are: (a) volcanic glass and plagioclase crystals of oligoclaseandesine composition - Nao.62 Cao.3sA11.38Si2.620s; (b) quartz - SiO,; (c) K-feldspar - KA1Si30s; (d) biotite - K(Fe,Mg), AlSi3010(OH); and (e) organic compounds, containing elements H, C, S and N, which are associated with coal and clayshale. Secondary minerals include: (a) calcite - CaCO, ; (b) siderite - FeCO, , (c) pyrite - FeSz (trace); (d) montmorillonite; (e) illite; (f) kaolinite; (g) chlorite; and (h) cristobalite. Mineralogy of the glacial drift Core samples of the glacial drift were retrieved during the installation of wells and piezometers in the study area. In order to'avoid contamination of the core samples, a hollow-stem auger and split-spoon sampler were used. An approximation of the relative abundance of minerals in the samples was obtained by X-ray diffraction on whole crushed (<2-mm sieve) and dried sediment. The data were interpreted in a general manner by calculating the ratio of the peak height of a given mineral with respect to the 3.33 (101) quartz peak. Although the technique is subject t o error because of the fact that peak height does not only depend upon the amount of mineral present, it does give some idea of the relative abundances of these minerals. The ratios indicated that quartz, dolomite, plagioclase, calcite, kaolinite, illite, montmorillonite, gypsum, mirabillite and tremolite were present in the glacial till (in order of decreasing ratio). While dolomite was the most abundant carbonate mineral in the glacial till, it was almost absent in the drift sands. Plagioclase and quartz were the major minerals of the drift sands and presumably represent the residue of the glaciofluvial erosion of the Belly River Formation. Coal fragments were preserved in the deepest drift deposits, most likely because of inclusion of large fragments of the bedrock during glacial override.
a
Chemical composition of groundwater and saturation indices Standard methods were used for chemical analyses of groundwater samples in the Geology Division Chemistry Laboratory of the Alberta Research Council. Concentrations of Ca2+,Mg2+, Na+ and K+ ions were determined by atomic absorption spectrometry. Turbidometric and volumetric titration methods were used to determine SO:- and C1- concentrations respectively. Concentrations of HCO, and C0;- were determined by potentiometric titration. Silica was determined colorimetrically. Results of analyses are summarized in Table I.
257 TABLE I Some results of groundwater chemical analyses Welt
Lab
Ho.
NO.
HC03
Sa rn p l ~n g
CI
SO4
NO1
S102,
F1:Hd
Cond
pH
Date
IY mho
~ a b
II/O/YR
Crn
215 204 218
7 I 6 7 6 7
19 2 580 53.0
195 278 266
I8 20 22
0 5 2.2 0 8
0 4
9 I 9 6 9.7
8 8 9 4 9.1
I050
213
1100
9 0 1 0 8.5
5.9 2 6 2 6
I0 9
275 283 290
4 6 5 0
14.4 77.0
227 220 216
2 8 0.0 7.1
I 1 I 4 0 5
7 7 9.4 9 5
8 6
24 18
9 4
50 0
476 142 361
20
5 0
1245 1290 1250
8 0 2 8 9 5
107.9 107.9
7.9 4.8
2.6 1.0
155
2.5 1.3
0.0
614 654
207 227
12
6.8 4.1
8.0
-
8.1
8 6
8 3
1400 I500
6 0
10/15/76 9 1 5/76
67.7 67 7
14.7 9.9
4.7 2 9
2.9 4.2
0.0
515 592
118
-
1500
5.0
5.8
8 0 5 0
8 1
126
8 I4
4 1
400
8.2
1550
-
5/22/75 12/19/75 4/1/76
68.6 68.6 68.6
2.6 1.5
9.1 2 I 2.1
276 149 112
108
10 28 26
1.4 I 1 0 6
0 8 1.6 I 0
9.5
210 212
10.1
9.6 9.8 9.7
I050 I015 830
6.0
751060 760019 760249
8/14/76 1/13/76 6/8/76
29.0 29.0 29.0
8.6
I-Z(P1
751059 760020 760250
8/14/75 1/11/76 6 1 8/76
68.6 68.6 68 6
RMI-4
761058 761055
9/7/76 10/15/76
RIw-LIP1 RMI-6
761057 761059
0BS-2lPl 00s-2 OBI-2
750261 751336 760094
RAH-4lQl
X
I
1.0
1.0
1.3
9.1 9 I
180
150
I
126f107
8/14/75 4/30/76 6/ 4/76
10.7 10.7 10.7
51.0 61.0 61.0
15.9 19.5 19.6
AH-5lYl
AH-5
751062 760139 760241
8/14/75 4130176 6 1 4/76
27.3 271 27.1
76.0 83.0 83.0
AH-7lWl AH-7 MI-7
750721 760017 760141
8/ 7/75 1/13/76 4/30/76
21.8 21.8 21.8
80.0 106.0 117.0
21.2
M-8lWl AH-8 M-8
751058 760142 760259
8/14/75 4130176 6/4/76
29.1 29.3 29.1
111.0
AH-9lWl M-9
760141 760219
4110176 6/4/76
15.8 15.8
470 50.0
0.0
0.0
72.0 122.0
1270 55+7
. 4.89.7
9.8i19.6 751061 760116 760256
l01.0 99.0
6.7 6 2 7.1
215
971i15 .. .
AH-2IWl AH-? Iw-2
AH-5
--
I2
I75
I11 ~~
10.6 19.0 19.9
7.1 7.4 7 7
7.4 7.6 7.9
675 650 670
18.0 8.2
4.0 0 0
9 5 14.8
0.0
14.0
7.4 7.6 7.2
7.8 7.4 7.2
575 780 850
180
10
1.5 5.5 2.9
11.2 12.3 12.1
7.2 7.2 7.4
7.3 7.4 7.6
900 900
8.0 18.0 12.0
10
1.0
8
1.6
16.8 181
7.5 7.7
675 725
11.0
750 1500
I5 0 9.0
71 82
183
52
6 1 5.8 6.7
0.0 0.0 0.0
127 456 476
71 77 87
6.7 5 4 5 6
0 0
419 417 429
I10 112
6 8
0.0 0.0 0.0
159 600
78 66
24.0
41 . 3
29.0 29 0
25 20
6 7 5.4 7 9
14.0
20 24 23.8
40.0 42.0 44.0
15 325
14.8
91.0
17.5
95.0
6.7 14.6
0 0
155
X f I
760144 760240
4/10/76 6 1 4/76
26.8 26.8
15.0 21 0
11.9 7 7
68.1*27
291
143
6.7 5.4
73.8.78 23.I!9.5
0.0 0.0
447 669
0.5 1 0
4
4.2
4
0.8
4 8
0.0
2
4 10
1.8
5.7t2 6
365 256
28 24
400-91
0.0
7.4
0.8
9.8
1.6t1.7 7.4r6 2
7.7
8.1
~
7.7 7.8
7.5t.J 12.8r4
855
7.5 7.7 ~
l07i71
0.0
5.0
..
9 I t . .~ W 1479f490 . 7 222 8.9r.6
8.0 16.0 8 0
8
10.9
8.0
600 700 750
181
6.5
-
7.4 7.8 7.5
390
98
0 0 0 0
9.0
8.0 7.5 7.8
0 0 0.0 0.0
5.8
__
1110
10 2 11.0
6
80
-
8 1.8.4
71.76 67 I50 96
0.0
I 9 1.1
2 01t2 4
197.7?94 137 176 400
5.4
12.5
188
451t161
71
31.0
195
~~~~~~~~~~~~~~~
AH-IO(YI M-I0
i’c)
19 I 14.6 13.7
I-IIPI 1-1 1-1
1-2 1-2
Temp
-11
9.0
1 8
12.0
~
740t190 7.5’.2
11.815
p = piezometer; W = well.
The digital computer program SOLMNEQ (Kharaka and Barnes, 1973) was employed t o calculate individual ion activities which were corrected for ionic strength and ion pairing, saturation indices for various minerals present in the sediments and the equilibrium partial pressure of C 0 2 . As shown in
258
Fig. 5, groundwater from both drift and bedrock aquifers is generally oversaturated with respect to calcite. However, a somewhat greater tendency toward oversaturation was exhibited by the groundwater in bedrock aquifers. Although groundwater from both drift and bedrock aquifers showed considerable variability in the saturation index with respect to dolomite (Fig. 6), the general tendency was toward oversaturation. Groundwater from both drift and bedrock aquifers was generally undersaturated with respect to gypsum, with groundwater in the bedrock aquifers being somewhat more undersaturated (Fig. 7). Inasmuch as cristobalite is one of the secondary minerals, the saturation index with respect to this mineral is of interest. Fig. 8 shows that groundwater in drift aquifers is saturated with respect to cristobalite and has a characteristic log-normal distribution, whereas groundwater in the bedrock is distinctly undersaturated. Finally, the equilibrium partial pressure of C 0 2 of groundwater in drift aquifers has an average value of
SATURATION I N D E X ( C A L C I T E ) 12
8
D R I F T AQUIFERS
f
4 0
Ii BEDROCK AQUl FERS
8 f
4 0 Log A P / K ( T )
SATURATION I N D E X (DOLOMITE)
4 f
0 log A P / K ( T )
f
4 0
BEDROCK A Q U I F E R S
-2.0
-1.0
0
1.0
2.0
log A P / K ( T )
Fig. 6. Saturation index with respect to dolomite.
3 0
259
lo-' atm. in contrast to atm. for groundwater in bedrock aquifers (Fig. 9). The former value for groundwater from drift aquifers is practically equal to the value 10-2.29atm. calculated from the empirical relation of Brooks et al. (1977), which relates the partial pressure of CO, in groundwater to soil temperature in the recharge area:
+ 0.06T("C)
logPco, = -3.00
This correspondence indicates that the source of CO, in groundwater from the drift aquifers is likely the soil atmosphere. The mean temperature of the groundwater in the recharge area is 11.S0C, and this value was substituted in the above equation.
SATURATION INDEX (GVPSUH)
12
f
8
4 0
-1.5
-2.5
-0.5
log AP/K(T)
I
log AP/K(T)
Fig. 7 . Saturation index with respect to gypsum.
SATURATION INDEX (CRISTOBALITE) 1
1
I
I
1
I
1
I
8
1
I
,
, I
f
-2.0
-1.0
0
Fig. 8. Saturation index with respect to cristobalite.
260 (Am)
EQUILIBRIUM P
co2 20
16 f
IZ 8 4 0.5
1.5
3.5
2.5
4.5
12
f
8 4 0
0 -LOG P
COZ
Fig. 9. Equilibrium partial pressure of COz ( P c o , ) .
Sources of major anions Carbon. Two sources of carbon are generally recognized as contributing t o the carbon load of groundwaters: (1)the CO, present in the soil atmosphere that is derived from plant root respiration and decay of organic matter; and (2) the COz resulting from oxidation of organic matter in the unsaturated zone and in the zone of fluctuating water table. Equilibrium partial pressures of C 0 2 for groundwater from the drift aquifers reflecting expected soil Pco, levels indicate that soil C 0 2 is the principal source. Gas bubbles were frequently observed in spring discharge pools in the study area, and methane was produced during aquifer tests. In addition, values of 6I3C determined on dissolved total inorganic carbon (DIC) from groundwaters fell into two groups: a light group that fell about half-way between a theoretical value of -25yoO for soil CO, and O ~ o ofor marine limestone and an abnormally heavy group (Table 11). These data tend t o suggest that methanogenesis must be taken into account in the manner of Barker et al. (1978) in order t o determine the proportion of the DIC that was derived from the production of methane by a reaction of the type:
2(CH20)
+H20
+
CH,
+ HCO, + H+
Sulfur. There are two possible sources of sulfate ion in the groundwater: (1)Sulfate derived from the dissolution of marine evaporite gypsum admixed with the till; and (2) Sulfate derived from oxidation of sulfide compounds in the drift that were derived from the Belly River Formation. One method to determine which of these two possibilities is correct is t o measure 634S of sulfate in aqueous extracts of drift and bedrock sediments and compare with values for marine evaporite gypsum and for Alberta coal.
TABLE I1 Results of analyses of environmental isotopes in drift and bedrock groundwater samples Well No.
Sampling date
OBS-1 OBS-2 1-1. 1-2 1-3 2-1 2-2 2 -3 3-1 3-2
212117 5 4/01/76 6/08/76 6/08/76 6/08/76 6/09/76 6/09/76 6/09/76 6110176 6110176
60
6IS0
('//oo, SMOW)
(%,
-143 f 4 (a) -148 f 4 -150 f 4 -153 f 3 -138+ 1 -151 f 3 -153 f 4 -147 + 3 -150 f 3 -160 ? 3
-18.3 -19.5 -20.4 -19.9 -19 -20.2 -20.6 -18.5 -20.3 -20.3
X = -149 AH-2 AH-3 AH-4 AH-5 AH-6 AH-7 AH-8 AH-9 AH-10
6/04/76 6/04/76 6/04/76 6/04/76 6/04/76 6/04/76 6/04/76 6/04/76 4/30/16
-140 -152 -132 -125 -133 -141 -138 -147 -161
+6
f 0.1
(b)
f 0.2
f 0.2 f 0.1
f0.2 f 0.1 f 0.1 f 0.4
+ 0.3 f 0.3
X = -19.7
f3 f3
+3 f4 +3 f4
f4
+ 11
X = -20.5
3H (TU)
7 f 10 (a) 3 Of 4 Of 3 o+ 3 o+ 3 Of 3 Of 3 15 3 Of 3
14cage (yr. B.P., cf. 1950)
6i3c ("&a
3
PDB)
o+
12,830 ? 285 (c) 24,160 f 1,070 (c) >30,320 (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 (<10,000 yr. B.P.) modifications t o the hydrogeologic environment, and given the generally low permeabilities, it is reasonable t o suspect that the total drift-bedrock system is not yet completely adjusted to these new boundary conditions. Rather what has occurred is that two distinct hydrogeochemical end-member systems exist: one in the shallow glacial drift and the other in the deep bedrock aquifers. At this point it is believed that these two systems did not develop simultaneously, and that the boundary conditions that govern the flow and chemistry of groundwater in each system are different. The conceptual model is presented in Fig. 18. Recharge to the drift
274
Fig. 18. Conceptual mixing model.
aquifer (V,, ) is denoted by Q 1 , recharge through the drift to the bedrock aquifer ( Vpz) is denoted by Q 3 , and water that recharged the bedrock aquifer outside of the spatial or temporal boundary conditions of the basin is indicated by Q 2 . The zone of active mixing of groundwater from drift and bedrock aquifers is given by V, and the discharge from the total system is Q1 Q 2 . (The terminology used is V for volume, p for plug flow, and m for mixing.) In terms of groundwater chemistry, the water from the drift aquifer differs chemically from that in the bedrock aquifer. (as indicated previously in Fig. 4),and mixing of these waters can result in compositional changes.
+
Isotopic evidence f o r groundwater mixing Evidence for mixing of groundwater from the drift aquifers with groundwater from the bedrock aquifers is found in the distributions of stable isotopes: deuterium and ''0 in water samples obtained along section B-B' (see Table I1 for data and Fig. 19A for meteoric-water line plot). Inasmuch as the stable isotopes are conservative tracers of water mass distribution and movement, it is possible to relate an heterogeneous isotopic distribution (large scatter) to plug flow in an aquifer where flow lines are essentially parallel and little mixing occurs. This would typify the type of flow observed in the drift aquifers where the flow of groundwater is downward through poorly permeable glacial till. On the other hand, homogeneity of the isotope distribution can be considered diagnostic of open flow and mixing such as in a groundwater discharge area where flow lines strongly converge. With respect to the distribution of the stable isotopes of hydrogen and oxygen along section B-B' as shown in Fig. 19B, a qualitative assessment of mixing was made on the basis of the scatter in 6-values. For example, in zones I , 11, and III,, there is much greater spread in the deuterium values than in the so-called well-mixed zones IIb and IIIb. Mixing appears t o be poor in the west half of the basin where the movement of groundwater is primarily downward through the glacial drift. This is in contrast to the condition that prevails at the groundwater discharge area in the valley bottom where, according to the latter criteria, groundwater is well mixed. The boundary between the local flow system that discharges in the basin and the deep regional system was marked by a rapid decrease in head with depth across the boundary and a change in the deuterium content.
275
-lO.O
-n.5 -11.0
-11.5 -10.0
3 42.5
-v,8
-13.0
5
n -13.5
4
-14.0 -14.5
-15.0 -15.5 -16.0 -16.5
-17.0
-23
-22
-21
-20
;
-17
-18
-19
6.0
(
- 549
I
549.
-16
x 1 SMOW
518-
-518
(km) LEGEND AND RANGES OF 6 VALUES -WATER
I na mb,Ob JJa
Range 60-18 -18.5 to - 1 9 . 0
Range 6D
- 1 3 8 . 4 t o -147.1
_-
TABLE
-Boundaries
I
of dynamic zones
P o o r l y mixed, T = >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
} T123 YBP
YEP
-15
276
Radiocarbon and tritium can be used t o determine the approximate age of water masses. Three age-depth zones were deduced from the distribution of tritium and I4C in the groundwater. As shown in Fig. 1 9 B, there is an upper zone containing tritium that extends from the water table down t o a depth of -6-21 m. The levels of tritium indicate that the source is probably fallout from atmospheric nuclear-fission bomb tests. Inasmuch as the testing of fission devices began in 1954, the age of water in this depth zone is probably less than 25yr. An intermediate zone in which tritium is absent but radiocarbon is detected extends to a depth of -75m below the upper zone. If the possibility of loss of 14C by mechanisms other than radioactive decay is neglected, a residence time of lo2-lo5 yr. is indicated. The deepest zone contains no detectable radiocarbon. It is in this zone that the regional groundwater flows system moves to the northeast, as previously indicated on section A-A’. MIX2 model computations The digital computer model MIX2 (Plummer et al., 1976) was employed to simulate the composition of groundwater that would result from a mixture of various proportions of groundwater from shallow drift aquifers with groundwater from the deep bedrock. The model MIX2 utilizes an aqueous model similar t o WATEQ (Truesdell and Jones, 1974) and the constraints of mass balance and electrical balance t o compute the pH and equilibrium distribution of inorganic species as a result of net reaction progress in the closed system : CaO-MgO-Na20-K,
0 x 0 -H2 SO4-HCl-H,O
In applying the model, the following assumptions were made: (1)that the calcite phase boundary is followed, and calcite is allowed to precipitate or t o dissolve in order to maintain chemical equilibrium; (2) that the temperature of the mixture is equal t o the weighted average of the end-member groundwater temperatures; and (3) that the volume of the mixing cell remains constant. Of these assumptions, evidence for the first is perhaps the most tenuous inasmuch as the saturation indices with respect t o calcite deviated considerably from equilibrium for the groundwater samples. Nevertheless, because calcite is disseminated through the aquifer, the assumption of equilibrium is a useful descriptor of the field case under consideration in the paper. Results of the model computations are presented in Table V. Three trial runs simulated the chemical equilibrium of various mixing proportions of deep bedrock (solution 1) and shallow drift (solution 2) groundwater. When groundwater from the shallow drift aquifers mixes with the alkaline bedrock groundwater, little change in the content of bicarbonate plus carbonate occurs because of the precipitation of calcite. The alkaline groundwater from the bedrock aquifers has a large buffer capacity, as shown by the gradual
TABLE V Results of computations with MIX2 solution-mineral equilibrium model Per cent solution 1
Per cent solution 2
pH
Total concentration in solution (meq./l) Ca
Mg
Na
K
C1
SO4
(HC03+C03)
CaC03 precipitated mol)
Temper
Run A : 0 100 5 10 20 40 60 80
9.7 7.8 9.66 9.63 9.53 9.18 8.27 7.83
1.3 58.0 0.32 0.35 0.44 1.04 9.08 28.8
0.3 16.8 0.75 1.34 2.63 5.8 9.4 12.5
350 48.8 331.0 317.0 288.0 229.0 168.0 108.0
1.3 1.8 1.3 1.3 1.4 1.5 1.6 1.70
174 2.0 167 159 143 112 80.5 48.9
7.4 73 10.5 13.7 20.0 32.7 44.4 54.4
517 281 400 394 382 357.9 326.0 288.7
0 100 5 10 20 40 60
9.5 7.5 9.45 9.40 9.37 8.82 8.09
1.4 47 .o 0.44 0.50 0.60 28.8 10.9
0.1 14.0 0.56 1.1 2.2 8.93 7.7
384 91.0 365.0 351.0 322.0 208.0 207.0
8.3 6.7 8.2 8.1 7.9 7.4 7.3
10 10 17.0 17.0 17.8 19.6 19.6
242 78 229 221 205 144 136
616 359 520 513 498 450 422
80 70 60 50 40
100 20 30 40 50 60
9.7 7.7 9.43 9.19 8.74 8.17 7.87
1.3 83 0.54 0.92 2.58 9.79 20.3
0.3 29 4.59 7.37 10.5 13.6 16.3
350 20 282 249 217 184 151
1.3 7.9 2.62 3.27 3.93 4.59 5.25
174 8 149 133 116 100 84
7.4 52 15.9 20.1 24.1 27.9 31.2
507 383 397 396 393 383 370
Run
Solution 1
Solution 2
A B
RH-1-3(760251) RH-3-2(751 064) RH-1-3(760251)
AH-4 (760257) AH-9 (760143) AH-5( 760241)
100 0 95 90 80 60 40 20
0.0901 0.160 0.299 0.566 0.636 0.393
11.5 7.5 11.3 11.1 10.7 9.9 8.3 7 .Y
0.0740 0.129 0.238 0.43016 0.403
9 .oo 12.00 9.15 9.30 9.60 10.20 10.80
0.422 0.615 0.775 0.791 0.717
11.50 8.20 10.84 10.51 10.18 9.85 9.52
RunB: 100 0 95 90 80 60 40
Run C : 100 0
~~
c
278
10.00-
I
G
0
0
o
I
0
\
0 0 0
,MIX2
0
TREND
0 0 0-2 0
BEDROCK A Q U I F E R S
0
D R I F T AQUIFERS
0
M I X 2 MODEL
00 0
0 0 0
7.00
I
LEGEND 0
BEDROCK A Q u I FERS
0
DRIFT AQUIFERS
0
MODEL RESULTS
9
8. PH
8. RUN C 0
0
7. 0
0 0
RUN B
0
0 Coo 00 0
7.8
50 “a+]
IOD
500
mg/l
Fig. 21. pH vs. Ca2+empirical and model results.
decrease in pH with increasing proportion of essentially neutral-pH shallow drift groundwater. The chemical type of groundwater mixture remains NaHCO,-C1/SO4 over most of the mixing range. The results of the mixing simulations are plotted together with the actual
279
analytical data from Table I in Figs. 20 and 21 for the cases of pH vs. Na' and pH vs. Ca2+ concentrations, respectively. The plot of pH vs. Na+ shows that a direct relationship exists. However, Ca concentration is inversely related to pH. As mixing of groundwaters from shallow drift and deep bedrock aquifers occurs, calcite precipitates in the aquifer. The model curves generally reflect the trend of the experimental data.
Discussion In order to clarify how mixing likely occurs in the groundwater system described in this paper it is important to consider the effect of hydrodynamic dispersion. Dispersion is a spreading phenomenon that causes attenuation of a physical or chemical property of groundwater. The dispersive properties of aquifers cause an inhomogeneity in the groundwater to gradually spread and t o occupy and ever-increasing portion of the flow domain beyond the region that it is expected t o occupy according t o average velocity and flow direction. Dispersion occurs due to: (1)mechanical mixing due t o fluid convection; and (2) molecular diffusion (Grisak et al., 1976). The spreading diameter cone ( u T ) in a groundwater flow fluid may be calculated from the relation (Harleman et al., 1963):
where p = mean groundwater flow velocity; d S 0 = mean grain size; v = kinematic viscosity of water; and L = the length of the flow path. For flow through the upper drift aquifer, the input values are p = 6.7 * 10' cm/s (Wallick, 1981); dSo = 4 l o p 4cm (very fine silt); v = 0.01 cm2/s;L = lOOOcm (arbitrary). The value of uT is calculated as 4.7 * lo2 cm. Similarly, in the case of the lower drift-upper bedrock aquifer, the values of p and dS0 are, respectively 2.77 * cm/s and 1.25 * cm (fine sand), and with the same values for the other two parameters, uT is equal t o 3.45cm. The lower drift bedrock aquifer is therefore almost 100 times as dispersive as the upper drift aquifer. Of interest is the extent to which dispersion is the result of mechanical mixing, or of molecular diffusion, in this particular groundwater system. One way t o tackle this problem is t o compare the flux of a relatively conservative chemical constituent such as C1-ion as dependent entirely on the rate of fluid movement with the flux by diffusion, assuming no fluid movement. For the case of the drift, the downward velocity of groundwater in the vertical direction is 6.7 * cm/s. The flux of C1- ion is the concentration (2.08 * mol/cm3) divided by the velocity, equal to 0.311 molcm-2 s-'. If it is true that the movement of chloride ion is not impeded by interaction with sediments, then C1- ions will move at the same velocity of water. Fick's first law may be used t o calculate the flux, F , of chloride in a direction normal t o the concentration gradient. In the case of chloride, the
280
concentration gradient is inverse with respect t o the head gradient, because the groundwaters from the bedrock aquifers have higher chloride content than those from the drift aquifers. Fick's law is written simply as:
F, = -o,(ac/az) where F, = flux in the z-direction = net mass of chloride transferred across a unit area of section normal to the flux in unit time (M L-2 T-' ); c = concentration (M/L3) ; and z = distance travelled (L). If the maximum difference in chloride concentration between groundwaters from bedrock and drift aquifers is equal to 200 mg/l = 5.63 mol/l (Table I), and the distance travelled is the mean saturated thickness of the drift, or 1500cm, and the appropriate diffusion coefficient of NaCl at 25OC is 1.576 lo-' cm2 s-' (Robinson and Stokes, 1955), F, is calculated molcm-2 s - l . This calculation, indicates that molecular to be 5.91 * diffusion is insignificant in magnitude when compared with the velocity of groundwater. It therefore appears that mechanical mixing in the lower driftupper bedrock aquifer is the most important cause of dispersion.
-
SUMMARY AND CONCLUSIONS
The chemical evolution of groundwater within the study area was explained by considering: the geologic history of the region since the Pleistocene, the geochemical sources of dissolved constituents, and the processes of groundwater flow and mixing. The author believes that the drastic changes in topography, geology and climate (components of the hydrogeological environment) brought by Wisconsin glaciation must be taken into account when interpreting the present distribution of chemical constituents in groundwater. For example, the distributions of hydraulic head, TDS, and radioactive and stable isotopes essentially reflect the super imposition of groundwater flow systems associated with the post-glacial topography (<10,000 yr.) upon a regional bedrock flow system of older but undetermined age. The glacial drift consists of sediments derived from local carbonaceous Late Cretaceous sediments plus a carbonate-rich fraction that was transported by glacier from the periphery of the Precambrian Shield. The weathering and leaching of these sediments during post-glacial time constitute the principle mechanisms that control the shallow groundwater chemistry. Mineralogic sources of major cations in groundwater from drift aquifers were calcite and dolomite (Ca, Mg), plagioclase-composition volcanic glass, smectite clays and illite (Ca, Na, K, Mg). Major anions in the drift aquifers consisted of sulfate that originates through oxidation and leaching of organo-sulfur compounds and pyrite, and bicarbonate derived from dissolved soil C 0 2 and carbonate minerals, and oxidized organic matter during bacterial processes such as sulfate and nitrate reduction, and possibly methanogenesis.
281
Groundwater compositions from the bedrock aquifers plotted in the gibbsite field on a Na2O-Al2CE-SiO2-H2O equilibrium phase diagram. This was interpreted as evidence of a very sluggish diagenetic chemical system in which the products of decomposition of the silicate minerals were very poorly leached. The principle products of the breakdown of the least stable mineralogic component in the bedrock sediments, plagioclase-composition volcanic glass, would be Na+, Ca2+and HCO; . Because the solubility of Ca2+ is limited by the solubility product of calcite, pore waters can only be enriched in Na+ and HCO;. The observation that groundwaters from bedrock aquifers are generally oversaturated with respect to calcite, that equilibrium Pco, is often much less than 10-3.5atm., the normal atmospheric value, and that calcium cementation is associated with the bedrock aquifers, tend to suggest that Ca2+ is mainly lost from solution by precipitation of calcite, although the relative importance of precipitation of Ca-zeolite and exchange on montmorillonite cannot be evaluated at this time. A mixing-model approach using the program MIX2 was employed to explain the range in chemical composition of the groundwater. The major assumptions inherent in the model were: that the drift groundwater flow systems were superimposed on the regional groundwater flow systems in the bedrock as a result of glaciation, that mixing of groundwater from these two systems gives rise to the variety of observed compositions, and that the equilibrium with respect to calcite is maintained. The model calculations showed that calcite precipitated from the mixtures, and that the observed range of groundwater compositions were in agreement with theoretical expectations. Mixing within the groundwater system was shown to result from mechanical dispersion rather than through chemical diffusion. ACKNOWLEDGEMENTS
The author would like to thank Drs. J. Toth and S. Moran of the Alberta Research Council and Dr. W. Back, U.S. Geological Survey, Reston, Virginia for their helpful comments. Mr. T.S. Balakrishna aided in completion of the field work. REFERENCES Barker, J.F., Fritz, P. and Brown, R.M., 1978. 14C measurements in aquifers with methane. In: International Symposium o n Isotope Hydrology, June 19-23, 1978. Int. At. Energy Agency, Vienna, pp. 661-678. Bayrock, L.A., 1967. Surficial geology of the Wainwright area (east half), Alberta. Res. Counc. Alta., Rep. 67-4. Brooks, G.A., Cowell, D.W. and Ford, D.C., 1977. Comment o n Regional hydrochemistry of North American carbonate terrains, b y R.S. Harmon, W.B. White, J.J. Drake and J.W.Hess; and The effect of climate o n the chemistry o f carbonate groundwater; by J.J. Drake and T.M.L. Wigley. Water Resour. Res., 13(5): 856-858.
282 Carrigy, M.A. and Mellon, G.B., 1964. Authigenic clay mineral cements in Cretaceous and Tertiary sandstones of Alberta. J. Sediment. Pertrol., 34( 3) : 461-472. Charron, J.E., 1969. Hydrochemical interpretation of groundwater movement in the Red River Valley, Manitoba. Can. Dep. Energy, Mines Resour., Inland Waters Branch, Sci. Ser. No. 2, 3 1 pp. Cherry, J.A., 1972. Geochemical processes in shallow groundwater flow systems in five areas in southern Manitoba, Canada. 24th Int. Geol. Congr. Proc. Sect. 11, pp. 208221. Christiansen, E.A. and Whitaker, S.H., 1976. Glacial thrusting of drift and bedrock. In: R.F. Legget (Editor), Glacial Till. R. SOC.Can., Spec. Publ., No. 12, pp. 121-132. Curtis, C.D., 1977. Sedimentary geochemistry: environments and processes dominated by involvement of and aqueous phase. Philos. Trans. R. SOC.London, Ser. A., 286: 3 5 3-37 2. Curtis, C.D., 1981. Chemical diagenesis of clastic sediments. Can. SOC.Pet. Geol. Semin. Text (available from Dep. Geol. Geophys., Univ. Calgary, Calgary, Alta.). Davison, C.C. and Vonhof, J.A., 1978. Spatial and temporary hydrochemical variations in a semi-confined buried channel aquifer: Esterhazy, Saskatchewan, Canada. Ground Water, 16(5): 341-351. Freeze, R.A., 1969. Regional groundwater flow - Old Wives Lake drainage basin, Saskatchewan. Can. Dep. Energy, Mines Resour., Inland Waters Branch, Sci. Ser. No. 5, 245 pp. Garrels, R.M., 1976. A survey of low temperature water mineral relations. In: Interpretation of Environmental Isotope and Hydrogeochemical Data in Groundwater Hydrology. Int. At. Energy Agency, Vienna, pp. 65-84, Garrels, R.M. and Christ, C.L., 1965. Solutions, Minerals, and Equilibria. Harper and Row, New York, N.Y., 450 pp. Green, R., 1972. Geological map of Alberta. Res. Counc. Alta., Map No. 35. Grisak, G.E., Cherry, J.A., Vonhof, J.A. and Blumele, J.P., 1976. Hydrogeologic and hydrochemical properties of fractured till in the interior plains region. In : Glacial Till, An Interdisciplinary Study. R. SOC.Can., Spec. Publ., No. 1 2 , pp. 304-335. Hackbarth, D.A., 1975. Hydrogeology of the Wainwright area, Alberta, Alta. Res. Counc., Rep. 75-1,16 pp. Harleman, D.R.F., Mehlhorn, P.F. and Ruiner, R.R., 1963. Dispersion-permeability correlation in porous media. J. Hydraul. Div., Proc. Am. SOC.Civ. Eng., 89(HY2): 67-85. Helgeson, H.C., Brown, T.H. and Leeper, R.H., 1969a. Handbook of Theoretical Activity Diagrams Depicting Chemical Equilibrium in Geologic Systems Involving an Aqueous Phase a t One Atmosphere and 0-300°C. Freeman and Cooper, San Francisco, Calif., 253 pp. Helgeson, H.C., Garrels, R.M. and Mackenzie, F.T., 196913. Evaluation of irreversible reactions in geochemical processes involving minerals and aqueous solutions, 11. Applications. Geochim. Cosmochim. Acta, 33: 455-481. Hitchon, B., Billings, G. and Klovan, J., 1971. Geochemistry and origin of formation water in the western Canada sedimentary basin, 111. Factors controlling chemical composition. Geochim. Cosmochim. Acta, 35 : 567-598. Kharaka, Y.K. and Barnes, I., 1973. SOLMNEQ: Solution-mineral equilibrium computations. U.S. Geol. Surv. (available from U.S. Dep. Commer., Natl. Tech. Info. Serv., Springsfield, Va., Rep. PB-215 899). Kovda, V.A. and Samoilova, V.A., 1969. Some problems of soda salinity. Proc. Symp. on Reclamation of Sodic and Soda-Saline Soils. Res. Inst. Soil Sci. Agric. Chem., Hung. Acad. Sci., Budapest, pp. 21-36. Le Breton, G. and Jones, J.F., 1962. A regional picture of the groundwater chemistry in particular aquifers in the western plains. Proc. 3rd Symp. o n Hydrology of Ground Water, pp. 207-245. Locker, J.G., 1973. Petrographic and engineering properties of fine-grained rocks of
283 central Alberta. Res. Counc. Alta. Bull., 30, 1 4 4 pp. Miller, B.E., 1972. A study of the authigenic minerals in the Blairmore Group, southern Alberta foothills. M.Sc. Thesis, Department of Geology, University of Calgary, Calgary, Alta. Moran, D.R., Groenwold, G.H. and Cherry, J.A., 1978. Geologic, hydrologic, and geochemical concepts and techniques in overburden characterization for mineral-land reclamation. N. Dakota Geol. Surv., Rep. Invest. No. 63, 1 5 2 pp. Oki, Y., Suzuki, T. and Hirano, T., 1977. High pH groundwaters of Tanzawa Mountains, Japan. 2nd Int. Symp. o n Water-Rock Interaction, Strasbourg, Aug. 17-25, 1977. Plummer, L.N., Parkhurst, D.L. and Kosiur, D.R., 1975 MIXB: a computer program for modelling chemical reactions in natural water. U.S. Geol. Suw., Water-Resour. Invest. 75-61, 6 8 pp. Robinson, R.A. and Stokes, R.H., 1955. Electrolyte Solutions. Butterworths, London. Rozkowski, A., 1967. The origin of hydrochemical patterns in hummocky moraine. Can. J. Earth Sci., 4 : 1065-1092. Rutherford, A.A., 1966. Water quality survey of Saskatchewan groundwaters. Chem. Div. Sask. Res. Counc., A.R.D.A. Proj. C-66-1. Sakai, H., 1957. Fractionation of sulphur isotopes in nature. Geochim. Cosmochim. Acta, 1 2 : 150-169. Scafe, D.W., 1973. Bentonite characteristics from deposits near Rosalind, Alberta. Clays Clay Miner., 21: 437-449. Toth, J., 1966. Groundwater geology, movement, chemistry, and resources near Olds, Alberta, Canada. Alta. Res. Counc., Bull. 1 7 , 1 2 6 pp. Toth, J., 1968. A hydrogeological study of the Three Hills area, Alberta. Alta. Res. Counc., Bull. 24, 117 pp. Truesdell, A.H. and Jones B.F., 1974. WATEQ: a computer program for calculating chemical equilibrium in natural waters. J. Res. U.S. Geol. Surv., 2: 233-248. Vanden Berg, A. and Lennox, D.H., 1969. Groundwater chemistry and hydrology of the Handhills Lake area, Alberta. Alta. Res. Counc., Earth Sci. Rep. 69-1, 49 pp. van Everdingen, R.O., 1968. Mobility of main ion species in reverse osmosis and the modification of subsurface brines. Can. J. Earth Sci., 5 : 1253-1260. van Everdingen, R.O. and Krouse, H.R., 1977. Stratigraphic differentiation by sulfur isotopes between upper Cambrian and lower Devonian gypsum-bearing units, District of Mackenzie, N.W.T. Can. J. Earth Sci., 14(12): 2790-2796. Wallick, E.I., 1981. Origin of the Horseshoe Lake sodium sulfate/carbonate deposit, Metiskow, east-central Alberta. Alta. Res. Counc. Bull. (in press).
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285
THE RATE OF FLUSHING AS A MAJOR FACTOR IN DETERMINING THE CHEMISTRY OF WATER IN FOSSIL AQUIFERS IN SOUTHERN ISRAEL
ARIE ISSAR
Institute f o r Desert Research and Geological Department, Ben-Gurion University, Beer Sheva 841 20 (Israel) (Accepted for publication November 24, 1980)
ABSTRACT Issar, A., 1981. The rate of flushing as a major factor in determining the chemistry of water in fossil aquifers in southern Israel. In: W. Back and R. LGtolle (Guest-Editors), Symposium on Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54: 285-296. A parallel trend has been found to exist between the isosalinity lines and the isopiezometric lines in the fossil aquifers of the Nubian sandstones in southern Israel and the central Sinai. The flow pattern in these aquifers determines the rate of mixing between paleo-seawater and paleo-meteoric water. This flowpattern is determined by the spatial relations between the Nubian sandstone outcrops in the central Sinai and the major fault lines of the Dead Sea and Suez rift valleys. A cumulative chemical composite diagram, in which " 0 stable-isotope data were plotted, makes it possible to classify the water and to trace the ratios in the mixture.
INTRODUCTION
The sequence of rocks of Early Cretaceous and pre-Cretaceous age in southern Israel and the Sinai includes in its stratigraphic units thick carbonate and clastic aquifers containing water of different salinities. The Lower Cretaceous sandstone aquifer (Kurnub Group) contains, for the most part, saline water. However, in the Sinai and the southern part of Israel it contains brackish water due t o the inflow of meteoric water which occurred during the humid Pleistocene periods. The recharge is believed to have taken place mainly in the sandstone outcrops in the central Sinai and to some extent in the erosion cirques of the Negev and Sinai (Issar et al., 1972) (Fig. 1). The Jurassic layers are mainly carbonates in the north and they become sandy toward the south. In the north, carbonate rocks of Jurassic age form the aquifer of Mount Hermon, which is the main source of the Jordan River through the Dan and Banias springs. Saline water has been found in deeply buried blocks which were reached by oil exploration wells (Bentor, 1969), in northern Israel.
Fig. 1. Generalized geological map and location map of the Negev and Sinai areas.
In the southern part of Israel, the Jurassic layer is partially composed of carbonates, and is partially sandy (Arad Group). Saline t o brackish water was found in these layers. The relatively low salt content of the water in the northeastern Negev was explained by Issar (1979) as a result of the same process, namely flushing, that changed the salinity of the Lower Cretaceous aquifers (Kurnub Group). Pre-Jurassic rocks were reached by wells only in the southern part of Israel and in the Sinai. As shown by Bentor (1969) the water is generally saline to highly saline. Along the major fault lines that border the downfaulted valleys of the
287
Syrian-African rift system, many thermo-mineral springs emerge. A large number of theories have been postulated t o explain their origin. Among them is the idea that trapped deep brines emerged to form springs (GoldSchmidt et al., 1967; Magaritz and Issar, 1973). It is agreed by all the authors who have dealt with the geochemistry of the saline water in the subsurface of Israel that most of the saline water found in deep wells is originally of ancient marine origin (Bentor, 1969; Mazor and Mero, 1969); Starinsky, 1974; Fleischer et al., 1977; Issar, 1979; Kafri and Arad, 1979). The difference in opinion is in regard t o the time and concentration processes of the water of these intrusions. Thus, while Bentor (1969) and Issar (1979) assume a pre-Cretaceous age and intra-formational diagenesis, Starinsky (1974) assumes surface evaporation processes in seawater of Neogene age; Mazor and Mero (1969) assume a Pliocene age (in the latter two references maintaining of a lagoonal process is assumed); and Kafri and Arad (1980) even discuss a recent penetration of the sea interface in addition t o the flushing of brines by meteoric water. This problem, however, is beyond the scope of the present study, which aims to explain the negative salinity found in the various aquifers, i.e. the sweetening of the saline water. In other words, in this work it is assumed that the deeply buried aquifers were saturated by formation water, and we will try to explain what the processes were which caused the water salinity to become less concentrated in time.
METHODS OF RESEARCH
The chemical character of the primary formation water groups It is true that the problem of the chemical character of the brines cannot be totally separated from the problem of their chemical genesis. However, it is suggested that the problem of genesis can be avoided by examining the deepest known brines and finding out their special chemical characteristics. One can assume that these are the compositions of the formation water of marine origin of this region, before being diluted by meteoric water. For this purpose, a characteristic graphical presentation of the chemical analysis is hereby suggested. This diagram presents the cumulative concentrations of the major anions and cations according t o their equivalence, the total concentration of salts, as well as the stable-isotope content of "0 as 6 l 8 0 (?oo) related to SMOW. The use of the l80analysis was preferred in this case to that using deuterium data, owing to the larger amount of 6"O analysis data available. The author agrees with Prof. J. Gat's comment that "0 content may have changed in the course of the subsurface flow as a result of the interaction with the aquifer rocks. By plotting the chemical composition of the water from the deepest wells, but from different aquifers and different regions, three characteristic graphs emerged (Fig. 2A-C). It is
288
suggested that the most saline water characterized in each of the graphs is to be regarded as the original ancient formation water. Thus, the original formation water can be divided into the following groups: deep water of the central Negev (Fig. 2A) deep water of the Dead Sea area (Fig. 2B) deep water of the northwestern Negev (Fig. 2C).
Group A: Group B: Group C:
As can be seen from this figure, groups A and C are enriched in salts and in the heavy isotope of l80in relation t o SMOW, which may point to their genesis through evaporation. Group B does not show a much higher concentration of l80 from seawater that could explain such an enrichment by evaporation. This may indicate that the genesis of this water is involving either a more highly concentrated brine mixed with fresh water or a genesis
meq
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289
which involved the dissolution of salts by meteoric water and evaporation. For a more detailed discussion of the possible origin of these brines the reader is referred t o Fleischer et al. (1977). Again, this paper does not aim to deal with the problem of the genesis of the brines, but with the processes which brought about a reduction in their salt content. Within this statement, however, a working hypothesis is hidden, namely that the saline and brackish water attained their salinity through dilution and not through dissolution of salts by fresh water. This working hypothesis can be shown t o be true if one surveys the isotopic content of the water groups (Fig. 2) which are less saline than the above discussed groups. This content clearly shows that the change which exists in the amount of isotopes in the less saline water is their depletion in relation t o the primary groups. This shows that the desalination of the brines .AkAVA
(N 1
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SUB-GROUP
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. HElMh(Tr) ............. ...............
I
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TAMAR-G(Lc)
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Fig. 2A and B. For caption see p. 290.
f!
8I8O
290
me( GROUP
c
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Fig. 2. Cumulative chemical composite diagrams, plotting the total concentrations of the major anions and cations and total concentrations of salts in milliequivalents against the stable-isotope content of ''0 as 6180 ("/oo) related to SMOW for: ( A ) deep water of the central Negev (group A); (B) deep water of the Dead Sea area (group B); and (C) deep water of the northwestern Negev (group C).
or the salination of the fresh water was mainly caused by dilution rather than solution procedures. This does not totally exclude the addition of a certain amount of salts by the dissolution of rocks by fresh water, but such an addition can be claimed to be negligible in comparison with that contributed by dilution.
The chemical character of the mixed saline water In order to understand how the dilution processes are operating, the data
291
from the aquifers below the Judean Group, i.e. the Nubian Sandstone aquifers of the Sinai and Negev, are shown as cumulative isotope curves. In Fig. 2A the cumulative semi-log curves of the chemical analysis of the water from the Negev and the "0 content are represented. Three main subgroups may be distinguished: Sub-group Al. This sub-group is composed of water from water wells in the Lower Cretaceous layers (Kurnub Group) in the southern Negev (Yotvata 9, Nakhel3, Makhtesh 3, Yorqeam 1) as well as from oil exploration wells which have penetrated into deeper layers: Avdat 1 (Jurassic), Ramon 1 (Paleozoic), Sherif 1 (Jurassic), Rekhme l a (Jurrasic). All waters from the Lower Cretaceous rocks exhibit rather parallel cumulative curves and the 6l'O-value is below -5.5%, related to SMOW. These waters represent a typical depleted type of paleowater found in this aquifer throughout the Negev and Sinai (Gat and Issar, 1974). However, the chemical analysis of the water from the deepest layers is different in some respects as it contains lower amounts of Ca but is richer in carbonate, sulfate and Na, and also slightly higher in its "0 content. This water may be regarded as sub-group Al-1.
Sub-group A2. This sub-group is distinctive in that it includes the water sampled from the deeper layers of wells in the northern Negev, Sherif 1 (Triassic), Dayya 1 (Jurassic). The cumulative curve as well as the "0 content being related to total milliequivalents show that sub-group A2 occupies an intermediate position between the original formation brine (group A) and the fresh paleowater (sub-group Al-1). In order to understand the relationship between the various groups, synthetic lines presenting the compositions (salinity and l80)of mixtures between water from group A and sub-groups A1 and Al-1 were plotted (Fig. 2A). (The numbers in circles on these lines present the ratio in volumes between the fresh and saline water.) As can be seen, sub-group A2 lies on a mixture line in the range of (7-5) : 1 of water of sub-groups Al-1 and group A, while sub-group Al-1 falls on a mixture line beween group A and sub-group A l . It is interesting to note that the waters of Yotvata 9 and Nakhel are outside the mixture line, namely that they did not play a role in the determination of the nature of the mixed saline water. This can be explained by the fact that the main bulk of water which caused the mixing was less depleted in its l80content (namely nearer in composition to Yorqeam 1 and Makhtesh 3 than t o Nakhel and Yotvata 9). An examination of the curves representing the water from the Dead Sea region (Fig. 2B) generally gives a similar picture; one can discern a sub-group ( B l ) which resembles sub-group A1 of the Negev in its chemical content and its isotopic character. In between, one can see sub-group B2 which resembles sub-group A2 although it exhibits a greater range in its chemical contents. The distribution of sub-groups B1 and B2 is representative of water from
292
me
I OO(
OZOHAR 7 ( J )
ID(
IC
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L Ca HC03 Mg SO4 Na
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Fig. 3. Cumulative chemical composite diagram with mixing of groups A, B and C (cf. Fig. 2), for water of the northeastern Negev.
different depths from two wells (Massad and Heimar) and the water wells Tamar 9 and 6, and shows that the salinity is becoming less the shallower the well. The water from the northeastern Negev (mainly from the Zohar gas field) includes a sub-group which is very similar t o sub-group B1 in its chemical and isotopic composition. It also includes another group which may represent a dilution between water similar to that of group C and sub-group B1, or another group of formation water having a composition similar to that which was found in well Zohar 8, in Paleozoic rocks (Fig. 3). A mixture of between group C originating in the northwestern Negev, and
293
water of sub-group B1 from the central Negev could be feasible in this region; it is situated at the junction of the flow lines of the water from these two regions (i.e. the central Negev and northwestern Negev) toward the Dead Sea which formed the outlet area for all of the Negev as it is the deepest downfaulted portion of the Syrian-African rift system (Fig. 1). Flushing and mixture model for the Negev A dynamic spatial conceptual model which illustrates the evolution of the various sub-groups of water as a function of the flushing and mixing processes is suggested in Fig. 4.
P R E RIFT DOWN FAULTING TO SUEZ RIFT ( AYUN MUSA)
STAGE
OUTFLOW TO DEAD SEA RIFT
POST RIFT DOWN-FAULTING
LEGEND Confining layer
nClosed border conditions Open border conditions
Fig. 4. Model of the evolution of the various sub-groups as a function of the flushing and mixing processes in the Negev.
I
NE
sw Tth escarpment
tl El
0 0
E!4
Carbonates of the Judeo Group (mainly limestone ond dolomite Cenomonion Turonion(aquiferous) Corbonotes of the Mt Scopus Avdot Groups ! M o i n l j cholks) Senonion to Eocene (oquicludes) Tarbonotes of the Arod ond KJrnJb Groups- Jurassic to L c (limestones m a r k ond shales)(semioquiferous) Clasttc of tile Ramon-Arad and Kurriub Groups Triossic I o L o w e r Cretoceous (aquiferous) Zlostic of the Ynm SJf ond Negev Groups Poleozoic !aquiferous) Crys'olline and mefomorphics - Precombrinn (imgerviods woter toble I1 x p r e z o m e t r i c heod -direction o f g w flow
Fig. 5. Hydrogeologic cross-section A-A' (see Fig. 1).
295
The basic hydrogeological assumptions which decide the regime of this model are the following: (1)The model is open in the south (the sandstone outcrops of the central Sinai) at altitudes ranging between 500 and 1000 m MSL. Along its western border (the Suez rift) at sea level, and along its eastern borders its outlets are at altitudes ranging from 100 m to -400 m MSL. It is closed along its northern and northwestern border. (2) The aquiferous layers are continuous from the southern end of the model t o its northern end. Further going north, the layers become subdivided and permeability is reduced (Fig. 5). (3) The upper aquifers are phreatic in the extreme south but are confined under most of the model area. The history of the inflow, flushing and mixture can thus be described as follows: (1)All aquifers were saturated by saline water or formation water. Each region was characterized by a certain group of water as illustrated in Fig. 4. (2) The inflow of meteoric water into the aquifers saturated by the formation water was made possible only when an outlet was formed for the formation water t o be pushed out. This may have started in the Neogene when transversal faulting occurred, but certainly took place during the Pleistocene after the down-faulting of the Syrian-African rift system. (3) The inflow occurred by a piston action which started mainly through the sandstone outcrops in the central Sinai (Fig. 1) but some inflow may have occurred in the sandstone outcrops in the northern Sinai and central Negev regions. (4)As the piston acted on all of the layers (Fig. 4) upwelling of saline water along major fault lines and unconformities could have occurred which caused mixture processes in spite of the piston action. (5) Another process of mixture occurred when the piston action in the south also caused formation water in the northern Negev t o flow toward the Dead Sea rift.
+
ACKNOWLEDGEMENT
The author would like to thank Alison Greengard for editorial assistance, and Prof. J. Gat and Dr. A. Arad for their useful comments.
REFERENCES Bentor, Y.K., 1969. On the evolution of subsurface brines in Israel. Chem. Geol., 4: 83110. Fleischer, E., Goldberg, M., Gat, J.R. and Magaritz, M., 1977. Isotopic composition of formation waters from deep drillings in southern Israel. Geochim. Cosmochim. Acta, 4 1 : 511-525.
296 Gat, J. and Issar, A., 1974. Desert isotope hydrology. Water resources of the Sinai Desert. Geochim. Cosmochim. Acta, 38: 1117-1131. Gilad, D. and Bachmat, Y., 1973. Israel’s groundwater basins. Water in Israel, Part A. Min. Agric., Tel Aviv, Water Comm., pp. 37-51. Goldschmidt, M.J., Arad, A. and Neev, D., 1967. The mechanism of the saline springs in the Lake Tiberias depression. Isr. Geol. Surv., Bull. No. 45. Issar, A., 1979. The paleohydrology of southern Israel and its influence on the flushing of the Kurnub and Arad Groups (Lower Cretaceous and Jurassic). J. Hydrol., 44: 289303. Issar, A , , 1981. A cumulative semi-log equivalence diagram, including stable isotope data to describe hydrochemical groups. Inst. Desert Res., Sde-Boker. Issar, A., Bein, A. and Michaeli, A., 1972. On the ancient water of the Upper Nubian Sandstone aquifer in central Sinai and southern Israel. J. Hydrol., 17: 353-379. Kafri, U. and Arad, A., 1979. Current subsurface intrusions of Mediterranean seawater. A possible source of groundwater salinity in the rift valley system, Israel. J. Hydrol., 44: 267-287. Magaritz, M. and Issar, A., 1973. Carbon and oxygen isotopes in epigenetic hydrothermal rocks for Hamam el Farun, Sinai. Chem. Geol., 12: 137-146. Mazor, E. and Mero, F., 1969. Geochemical tracing of mineral and fresh water sources in the Lake Tiberias basin, Israel. J. Hydrol., 7: 246-275. Starinsky, A., 1974. The relationship between Ca-chloride basins and sedimentary rocks in Israel. Ph.D. Thesis, Hebrew University, Jerusalem (in Hebrew).
297
GEOCHEMICAL INPUTS FOR HYDROLOGICAL MODELS OF DEEP-LYING SEDIMENTARY UNITS: LOSS OF MINERAL HYDRATION WATER D.L. GRAF and D.E. ANDERSON
Department of Geology, University of Illinois, Urbana, I L 61801 (U.S.A.) (Accepted for publication February 26, 1981)
ABSTRACT Graf, D.L. and Anderson, D.E., 1981. Geochemical inputs for hydrological models of deep-lying sedimentary units: loss of mineralogical hydration water. In: W. Back and R. Lbtolle (Guest-Editors), Symposium on Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54 : 297-314. Hydrological models that treat phenomena occurring deep in sedimentary piles, such as petroleum maturation and retention of chemical and radioactive waste, may require time spans of at least several million years. Many input quantities classically treated as constants will be variables on this time scale. Models sophisticated enough t o include transport contributions from such processes as chemical diffusion, mineral dehydration and shale membrane behavior require considerable knowledge about regional geological history as well as the pertinent mineralogical and geochemical relationships. Simple dehydrations such as those of gypsum and halloysite occur at sharply-defined temperatures but, as with all mineral dehydration reactions, the equilibrium temperature is strongly dependent on the pore-fluid salinity and degree of overpressuring encountered in the subsurface. The dehydrations of analcime and smectite proceed by reactions involving other sedimentary minerals. The smectite reaction is crystallographically complex, yielding a succession of mixed-layered illite/smectites, and on the U.S.A. Gulf of Mexico coast continues over several million years at a particular stratigraphic interval.
INTRODUCTION
Hydrological models that treat fluid behavior in sedimentary environments that lie under some kilometers of overburden are likely to require much greater time spans than is customary in studying near-surface processes. The longer-lived isotopes of transuranic elements and their daughters present in buried radioactive waste will persist for several million years (Bredehoeft et al., 1978). An example of a substance that needs, in principle, t o be retained indefinitely because of the strict limits set upon its allowable concentration in drinking water is nitrate, an anion that is unfortunately not adsorbed effectively by clays and other common subsurface minerals. Taking 29'C/km, the mean of the extreme values of geothermal gradient cited by Heard and Rubey (1966) for the U.S.A. Gulf of Mexico coast, and 40 cm/ka, a rate of accumulation of substantially compacted sediment on the Gulf of Mexico coast cited by Bredehoeft and Hanshaw (1968), the
298
0.5 km of sediment accumulated in a million years would raise the temperature of an underlying organic-rich unit by 14.5"C. Even if increased t o allow for greater thickness of the l-Ma deposit because of incomplete compaction, this temperature change is only a small fraction of the temperature range over which hydrocarbons are produced by maturation of organic matter. One of the sources of fluid for hydrocarbon transport is pore-water expulsion because of sediment compaction and cementation, and it is instructive t o note that the shallow marine sediments off the U.S.A. New Jersey coast appear t o have a reasonably constant porosity only below the Cretaceous-Jurassic boundary (Steckler and Watts, 1978), i.e. after -135 Ma (Kulp, 1961). Standard expressions for convective flow and for rate of chemical reaction, for example, can obviously treat the lengthened time span easily enough if the required input information can be considered constant with time. But a million years is more than adequate for sharp changes of climate and of relative rates of erosion and sedimentation and, therefore, of topography, all of which have obvious effects upon shallow aquifers. The thermal regime of deeper horizons may be altered both by changes in thickness of overlying column and by igneous intrusion. Deeper in sedimentary piles there are processes taking place that are rarely if ever obeserved at the surface, e.g., injection of hot water from igneous intrusives and the release of water of hydration from minerals such as gypsum, zeolites, halloysite and smectite. Other processes that are relatively unimportant at shallow depths, e.g., the lateral diffusion of dissolved salts on gradients of concentration or temperature and the cross-formational flow of water and dissolved solids through the tiny, electrically-charged pores of shales, become significant in this region where lateral convective flow is slight or nonexistent. Over short time spans, the deep environment could be modelled as static with respect t o these slow processes. However, in view of the arguments presented above for long-time-span models, this simplification is not permissible. This paper considers the amount and quality of information about one of these slow processes, loss of hydration water from sedimentary minerals, that is available for input into hydrological models.
LOSS OF HYDRATION WATER FROM SEDIMENTARY MINERALS
Gypsum The equilibrium between gypsum, CaS04 *2H20, and anhydrite, CaS04, is probably the best-understood dehydration of a rock-forming sedimentary mineral, now that procedures have been devised for extrapolating experimental measurements into the region of simple aqueous solution below 80°C where anhydrite dissolution and crystal growth are kinetically hindered. The gypsumanhydrite transition temperature, T,,at 0.101 MPa (1atm.) in
299
simple aqueous solution was fixed at 58 f 2°C by up-temperature extrapolation from values for Na, SO4- and H, SO4-loaded solutions of lower water activity (Hardie, 1967) and at 56 k 3°C by down-temperature extrapolation from runs in simple aqueous solutions at 80°C and above, with correspondingly increased fluid pressures (Blount and Dickson, 1973). The pressure coefficient of Tt computed by Zen (1965) from thermochemical data for systems in which fluid and solid are under the same pressure is 7.09 k 0.193 MPa/"C (70 f 1.9 atm./"C) whereas Blount and Dickson (1973) obtained 7.8 f 0.7 MPa/"C (78 k 7 bar/"C) from their experimental runs in the range 0.1-100 MPa (1-1000 bar). If mean sedimentary rock density is taken to be 2.4 g/cm3, and the solids are under lithostatic pressure but the pore fluid is under the pressure only of the fluid column above it, the pressure dependence of Tt is calculated to be -4.32 f 0.071 MPa/"C (-42.6 k 0.7 atm./"C) (Zen, 1965). Taking 22 and 36'C/km as extreme values for Gulf of Mexico coast geothermal gradients (Heard and Rubey, 1966) and using values of the preceding paragraph for simple aqueous solution, gypsum should vanish between the -1.07-km (3500-ft.) and 2.44-km (8000-ft.) depth in that region. In fact, considering a variety of geographical localities, gypsum is seldom observed in core samples from below 0.61-0.91 km (2000-3000 ft.) (MacDonald, 1953; Stewart, 1963); Murray (1964) reported an instance of gypsum replacing anhydrite at 1.07-km (3500-ft.) depth. The discrepancy results from the combined effect in nature of pore-fluid salinity and of unequal pressure on liquid and solid, both of which decrease Tt and therefore reduce the depth to which gypsum persists. The reality of the latter effect on the Gulf of Mexico coast is shown by fig. 5 of Kharakha et al. (1977), in which pore-fluid pressure is seen t o be hydrostatic t o 2.85 km depth in the Corpus Christi, Texas, area and to 3.15 km depth in Brazoria and Galveston counties, Texas, increasing rapidly t o -85% of lithostatic in the 3.7-4.5 km depth interval ("overpressured"). Gay (1965) concluded from crystallographic measurements and phasetransformation experiments that there are only two additional solid phases in the system CaS0,-H2 0 at 25°C and 0.101-MPa (l-atm.) pressure besides gypsum and anhydrite, i.e. soluble anhydrite and bassanite, CaS04*0.5 H,O. Soluble anhydrite has never been reported in nature, and forms in the laboratory by dehydration of bassanite, but not in experiments involving liquid water. Heard and Rubey (1966) cited preliminary experiments of G.C. Kennedy, suggesting that the quadruple point (gypsum-anhydritebassanite-H20) is near 600 MPa (6 kbar) and 130"C, with the bassanite field widening in temperature at higher pressures. Zen (1965) estimated from thermochemical calculations and chemographic analysis that the quadruple point should be near 900 MPa (9 kbar) and 170°C. Solubility and phasetransformation experiments (Zen, 1965; Kinsman, 1974) indicate that bassanite is metastable at 0.101 MPa (1atm.) and temperatures of 25-70°C. The conclusion of Goodman (1957) from petrographic examination that
300
bassanite replaces both anhydrite and gypsum in a core from evaporite deposits in Nova Scotia appears inadequate t o prevail against the contrary evidence. Bassanite may be presumed to appear in diagenetic environments only as a metastable intermediate between gypsum and anhydrite, with a limited lifetime, and can therefore be ignored in modeling. Gypsum dehydration is thus a relatively simple, short-time event occurring at temperatures close t o the lower end of the range observed in sedimentary piles. Selecting a pore-fluid salinity vs. depth relation for a particular geographical region involves assumptions about the spatial distributions through time of different sedimentary lithologies, in particular that of the highly soluble halite often found interbedded with gypsum in evaporite sequences and intruding overlying formations as domes. Inferring the distribution of overpressured zones with time is a much more difficult task, involving such elements of geologic history as sedimentation rate and the changing permeability of the resulting sediments as they compact. Overpressuring exists at shallow depths in some regions (Deju, 1973; Fertl, 1976), so that it cannot be assumed that gypsum dehydration always takes place in normallypressured environments. The reverse reaction, the hydration of anhydrite, has been observed in cores from regions where the most recent geologic event has been erosion (Murray, 1964).
Smectite The water releasing mineral of greatest geological interest is dioctahedral smectite, because it is relatively abundant and widespread in sediments, but the physicochemical parameters of this dehydration are much less well understood than those for gypsum. Interlayer water can in principle be driven from smectite by changing pressure, temperature, or the ionic strength of the pore fluid, or by carrying out a diagenetic reaction involving smectite plus components dissolved in pore fluid and/or present in other detrital minerals, t o yield one or more anhydrous solid phases. van Olphen (1963) computed the work of adsorption of successive interlayers from water vapor sorption isotherms and changes in basal interplanar spacings with water loss, and concluded that 255 MPa (37,000 p.s.i.*) was needed t o remove the next t o last water interlayer from a Ca-bentonite, 545 MPa (79,000p.s.i.) for the last. This calculation has been taken to indicate that smectite cannot be dehydrated by pressure alone in sedimentary piles of the thickness available. It applies, of course, to a sealed system at equilibrium, i.e. solid and fluid are under the same (lithostatic) pressure and the activity of water has adjusted to a common value in the two phases. For a calculation of the consequences of pore pressure being hydrostatic or intermediate between hydrostatic and lithostatic, analogous t o that of Zen (1965) for gypsum, one would like to
* 1 p.s.i. = 1 pound-force per square inch (lb./in.2) = 6 . 8 9 5 . 1 0 3 Pa = 0.006895MPa.
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have available concordant stability-field determinations for fhe solid phases determined by two methods, e.g., solubility and calorimetric measurements. The only clay mineral for which the latter, heat-of-solution and lowtemperature specific-heat measurements, are available appears to be kaolinite (Hemingway et al., 1978), but is may be possible t o extend this thermochemical information t o related structures and compositions by using estimation procedures such as those of Tardy and Garrels (1974), and Nriagu (1975). Most of the information about dehydration because of temperature increase has come from differential thermal analysis curves, which are deceptive with regard t o the subsurface environment because of the rapid rate at which temperature is raised (water loss therefore occurs at a higher temperature than it would otherwise), and because the sample is surrounded by air at 0.101 MPa (1atm.) pressure rather than a saline solution at elevated pressure. Nevertheless, these curves (e.g., Hendricks et al., 1940; Rowland et al., 1956) indicate a relative order of events. There have been some studies (e.g., Norrish, 1954; Norrish and Quirk, 1954) of the loss of interlayer water to a highly saline pore fluid t o equalize the fugacity of water in the two environments. These authors monitored the changes in do,, of montmorillonites placed in NaC1, CaCl, and MgC1, solutions of various ionic strengths, at room temperature. Studies in which interlayer water content adjusts to match the water fugacity of an atmosphere maintained at constant relative humidity have been more common and have repeatedly shown that interlayer water loss with decreasing relative humidity is stepwise and that each step can be interpreted as the loss of one or more molecular layers of water (Bradley et al., 1937; Hofmann and Hausdorf, 1942; Mering, 1946; Cornet, 1950; Mooney et al., 1952; Slabaugh, 1955). At relatively low humidities, water molecules coordinated around adsorbed cations such as Ca2+ and Mg2+ may yield incomplete molecular layers of water (Hendricks et al., 1940; Mering, 1946). Although there is no bulk liquid with which adsorbed cations can be exchanged in the experiments made in atmospheres of controlled relative humidity, successive experiments can be made with different adsorbed cation populations, and it is found that the succession of dehydration steps varies with size and hydration energy of the cation (Posner and Quirk, 1964). There is also a dependence upon the type and extent of substitution in the aluminosilicate framework; White (1958), for example, observed that two montmorillonites with the same cation-exchange capacity, loaded with the same cation, may have different water sorption and swelling properties. Experiments have not been carried out in which the combined effect upon smectite dehydration of increasing temperature, lithostatic pressure, hydrostatic pressure and ionic strength of pore fluid is tested in order t o simulate the diagenetic environment. Pairs of these variables have been studied in limited numbers of experiments, some of them difficult to interpret, by Stone and Rowland (1955), Weaver (1959), von Engelhardt and Gaida (1963), Khitarov and Pugin (1966), and Demirel et al. (1970). The accumu-
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lating evidence about how smectite dehydration proceeds in nature clearly indicates that future laboratory studies must also monitor alteration of the aluminosilicate framework t o form other anhydrous phases. The succession of changes that takes place when detrital montmorillonite is subjected t o progressively deeper burial is best displayed in regions of unusually high geothermal gradient, e.g., the Ohaki-Broadlands hydrothermal area, New Zealand (Steiner, 1968; Browne and Ellis, 1970; Eslinger and Savin, 1973); Steamboat Springs, Nevada (Schoen and White, 1965); and the Salton Sea geothermal region (Muffler and White, 1969). The smectite first reacts to form a randomly interlayered illite/smectite that increases steadily in percentage illite layers, and then a nearest-neighbor-ordered illite/smectite containing only 20--30% smectite layers (allevardite), and is finally converted to illite plus chlorite. In thick sedimentary sequences with more nearly average geothermal gradients, the sequence can be followed as far as the ordered interlayered illite/smectite, e.g., the Ventura Basin (Quaide, 1956); Japan (Iijima and Utada, 1971; Mitsui, 1975); Papua (van Moort, 1971); Cameroon (Dunoyer de Segonzac, 1964); Northwest Territories (Foscolos and Powell, 1978); the Rhine graben (Heling, 1978); and the most thoroughly studied of all, the Gulf of Mexico coast (Burst, 1959, 1969; Powers, 1959; Weaver, 1959; Perry and Hower, 1970, 1972; Reynolds and Hower, 1970; Weaver et al., 1971; Hower et al., 1976; Yeh and Savin, 1977). The smectite-illite structural sequence t o allevardite is from Reynolds and Hower (1970), and various of the other studies deviate from it in detail, but agreement is general about the insignificant amount of trioctahedral smectite in these sedimentary piles. Hower et al. (1976) concluded that the illitization reaction is: K+ + A13++ X-smectite
+ illite + Si4++ X +
where X denotes adsorbed cations other than K+, and that the K+ and A13+ are supplied by breakdown of K-mica and K-feldspar, which are observed t o decrease in amount with depth. Yeh and Savin (1977) found that authigenic SiO,, required by the above reaction unless fluid drainage is assumed to be improbably rapid, can be distinguished in Gulf of Mexico coast cores from detrital quartz by its stable oxygen-isotopic composition. In contrast, Boles and Franks (1979) suggested that A13+ is an immobile component in the smectite + illite reaction, which is then approximately:
K+ t X-smectite + illite iquartz
+ X + + H+
This reaction involves the loss of part of the clay phase, yielding only 0.8 mol of illite per mol of smectite, and presumably requires solution and reprecipitation. However, the Ar/K studies of Aronson and Hower (1976) indicate that inherited radiogenic argon from illite interlayers in original detrital smectite is retained, so that the original clay mineral structure must have persisted. The smectite + illite conversion is crystallographically and geochemically
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complex in a number of ways in addition to the formation .of random and regular illite/smectite mixed-layered assemblages already mentioned. The basal-spacing values cited for pure smectites in many instances result from either random or regular mixed-layering of two hydration states, and it must be assumed that the interlayer water remaining in the smectite component of illite/smectite mixed-layered assemblages is distributed in similar fashion. Jefferson (1938) and Hofmann and Hausdorf (1942) described the continuous but nonuniform variation of do,, for random-mixed layering of hydration states. Grim (1968, p. 263) reported that smectite held a long time under substantially uniform moisture conditions has a hydration of considerable stability, that is, it is slow t o change when placed in a sharply different hydration environment, such as liquid water. The implication is that with time the hydration layers develop a high degree of uniformity of thickness and distribution. A varying charge density on different smectite layers is inferred from experiments in which, variously, hydroxy-aluminum polymers, piperidine, dodecylamine, and n-alkylammonium ions of different chain lengths were intercalated (Byrne, 1954; Hsu, 1968; Stul and Mortier, 1974). In particular experiments, there was a failure t o develop a regular periodicity along c, or a less than sharp transition from monolayer to double layer of the intercalated substance, or a range in the difficulty of subsequent extraction of all of the intercalated material from a particular sample. Laboratory adsorption studies with mixed Na+--Ca2+solutions (Glaeser and Mering, 1954; McAtee, 1956; Mungan and Jessen, 1963; Levy and Francis, 1975) indicate that, for subsurface brines having certain ratios of alkalineearth to alkali cations, there will be demixing of adsorbed ions into separate interlayer regions or into particular particle size ranges rich in one or the other of the cation types. The occurrence in the Gulf of Mexico coast and elsewhere of overpressured zones, which typically have pore-fluid pressure greater and pore-fluid salinity less than expectable for their depth, has been attributed by some t o the release of smectite interlayer water (e.g., Burst, 1959; Powers, 1967; Dickey et al., 1968). This mechanism appears t o imply: (1) that the partial molar volume of water in the interlayer region of a clay mineral is less than that which it has in adjacent pore fluid after release; (2) that the rate of interlayer water release is greater than the rate at which it moves away from the stratigraphic region in which it is being released; and (3) that the lithologic framework remains fixed during dehydration. The first requirement is supported by the calculation of Martin (1962), who concluded that the last few interlayers of water in the crystal probably have a density of -1.4 g/cm3, but expressed a general lack of confidence in the reliability of the values available in the literature for making the comparison. Anderson and Low (1958), on the other hand, assigned lower values t o the density of interlayer water than t o adjacent free water. Hanshaw and Bredehoeft (1968), assuming an interlayer water density of 1.4g/cm3, calculated the change of pore pressure with time for different sets of assumed parameters and concluded
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that an abnormally high fluid pressure could be developed and maintained if the amount of montmorillonite were high and the permeability were low. The third assumption is debatable. Erosion or sedimentation at the top of the sedimentary column may proceed at uniform or varying rate. The dehydration of smectite with two interlayers of water represents a loss of -40% of mineral volume, so that the rocks of the dehydration zone can hardly retain strength unless the mineral hydrate is concentrated in scattered small regions, more likely for nodular gypsum than for smectite. Water released with no molar volume change might therefore be observed later t o be at unusually high fluid pressure because it is sustaining part of the weight of a rock column that has subsided and compacted. The principal additional processes that have been suggested for overpressuring are: (1) rapid sedimentation coupled with a decrease in sediment permeability during burial, such that fluids are unable t o escape rapidly enough to maintain hydrostatic equilibrium (“nonequilibrium compaction”, “compaction disequilibrium”) (Magara, 1975a); (2) a.quatherma1 pressuring (Barker, 1972; Magara, 1975b), the result if a volume of brine-saturated sediment that can be considered sealed is raised in temperature because of deeper burial; and (3) lateral tectonic compression (e.g., Gretener, 1976). A recent review of overpressuring (Plumley, 1980) concludes that shale porosities in Gulf of Mexico coast overpressured zones are too low t o have resulted solely from nonequilibrium compaction, and that aquathermal pressuring and/or mineral dehydration must also have been operative. Existing experimental apparatus permits several types of dehydration experiments for the simpler case in which smectite and brine are under the same uniform pressure. Using X-ray cells of the general type described by Graf (1974), basal diffraction maxima can be monitored to detect the loss of interlayer water as smectite in contact with brine is subjected t o increasingly higher temperature and brine fluid pressure. The fluid-pressure change at constant volume and the volume change at constant fluid pressure can in principle be measured in other types of pressure vessels, for a mixture of smectite and K-feldspar reacting in a chloride brine t o form interlayered smectite/illite. There is a limited temperature region around 250” or 260°C that is about the upper limit for long-term integrity of Teflon@ seals and paste-on strain gauges, and about the lower limit for obtaining reaction in reasonable times for illitization runs. K-saturated Wyoming bentonite, sealed with water in a gold capsule under 200 MPa ( 2 kbar) external fluid pressure and held for 167 days at 250”C, yielded an ordered illite/smectite containing -20% expandable layers (Eberl and Hower, 1977); a similar run for 74 days at 285°C gave the same product, but with -35% expandable layers. It seems unlikely that Na’ and Ca2+in the brine will enter the anhydrous interlayers formed in these experiments; Velde (1969) pointed out that the upper stability limit of K-montmorillonite relative t o K-mica, near 250”C, is markedly lower than that for Na-montmorillonite, which does not convert t o Na-mica until as high as 400°C. Velde (1977) also predicted from the results
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of Hemley et al. (1971) that calcic beidellite should have a relatively high thermal stability. Water release from smectite would appear t o be a good deal more difficult to incorporate into a hydrological model than that from gypsum. Obviously, geochemists must translate the complex structural changes during dehydration into an expression simple enough to be usable. But the size of the time and temperature ranges over which dehydration takes place in itself probably necessitates a more elaborate model. Hower et al. (1976) reported that in the Gulf of Mexico coast core that they studied the proportion of illite interlayers increased from 2.15 to 3.70 km depth, the allevardite structure with 80% illite layers then persisting t o at least 5.50km depth. The illitization interval corresponds to a AT of -5OoC, a At of -3 Ma. Mixed-layered illite/smectites from cores in geologic areas where there has been substantial erosion of the overlying stratigraphic section give no evidence of the conversion of illite interlayers back t o smectite.
Hall0 ysite
Parham (1969a) found that halloysite, (OH), Si4A14010*4H2 0, was the product formed in laboratory experiments that simulated tropical weathering of feldspar. Likewise, it is the abundant clay mineral in deeply weathered metamorphic, igneous and sedimentary terrains at Hong Kong (Parham, 1969b). In nature, initially-formed halloysite subsequently converts to kaolinite, (OH),Si4A140,0. Nagasawa and Miyazaki (1975) concluded from examination of a considerable number of natural deposits that halloysite will be present in greatly reduced amounts in deposits as old as Cretaceous. There is disagreement in the literature about whether kaolinite abundance decreases with depth in a particular region, or is random. The view that kaolinite is progressively destroyed with depth and that it would therefore contribute A13+t o the formation of both illite and chlorite is summarized by Muller (1967) and Weaver et al. (1971). If halloysite is air-dried, it retains 24-23 of the four formula waters, which corresponds roughly t o one water interlayer per four silicate layers (Brindley and Goodyear, 1948). During dehydration, halloysite maintains a statistical distribution of hydrated and nonhydrated layers as described theoretically by Jefferson (1938). Roy and Osborn (1954) concluded from sealed hydrothermal runs in simple aqueous solution that the 4-hydrate converts t o the 2-hydrate at -175"C, water pressures of as much as -192 MPa (30,000 p.s.i.) increasing this temperature only very little. In air, heating to -400°C is required to drive off the remaining interlayer water (Brindley et al., 1948). Dehydroxylation of kaolinite is seen in differential thermal analysis traces between 400" and 600°C (Grim, 1968), but with expanded heating times of 200 hr. this reaction takes place as low as 350°C (De Keyser, 1939). There are unusual crystallographic aspects t o the dehydration. The tubular
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crystals of the 4-hydrate frequently split, collapse, or unroll when they convert t o the dihydrate (Bates et al., 1950). And the X-ray diffraction measurements of Brindley et al. (1963) documented the existence of a continuous structural series from halloysite to kaolinite, made possible by layer-stacking disorder, as well as a corresponding morphological series from tubes through curved forms to laths to plates. It appears that there may be a significant release of water t o pore fluid from the halloysite -+ kaolinite conversion in stratigraphic sections containing reasonably feldspathic rocks that have undergone tropical weathering. Because of the relative rates involved, it should be feasible t o measure experimentally the effect, upon the dehydration, of pore-fluid salinity and of different pressures upon liquid and solid, without producing additional solid phases. In nature, after conversion of halloysite t o kaolinite, there would be an additional transfer of water between pore fluid and mineral if the kaolinite -+ illite or the kaolinite -+ chlorite reaction occurred, and it would not be the same for the two reactions. There seems little point in calculating the consequences of postulated kaolinite reactions until there is better agreement about their geological significance. Vermiculite Vermiculite is a widespread hydrated clay mineral of lesser abundance than smectite. Understanding of its diagenetic behavior is complicated by the fact that most of the laboratory measurements on vermiculite have been made on the hydrothermal alteration products of coarse-grained biotites, whereas most soil vermiculites that would be fed into a sedimentary pile like that of the Gulf of Mexico coast are weathered dioctahedral illites containing a lot of interlayer Al- and Fe-hydroxides. The maximum amount of interlayer water taken up by vermiculite, even when the sample is immersed in liquid water, is two layers for Mg2+or Ca2+ saturation (Barshad, 1948, 1950; Walker, 1956; van Olphen, 1963). There is one interlayer for Na' saturation, and there is some interlayer water present in K+-saturated vermiculite, because heating causes a basal spacing decrease but the decrease is less than that required for a water layer in to 10.1 vermiculites saturated with the other cations mentioned. Walker (1956,1957) and van Olphen (1963) discussed the conditions of temperature and humidity at which partly-hydrated Mg-vermiculites exist in air. Walker interpreted differential thermal analyses (DTA) charts as indicating that interlayer water not associated with the hydration of adsorbed Mg2+ ions is released at a lower temperature than is that around the Mg2+.Most of both of these types of interlayer water is released by 275°C. Vermiculite held under -63 MPa (10,000 p.s.i.) water vapor pressure shows its first dehydration at -550°C (Roy and Romo, 1957). Roy and Romo (1957) noted a change in relative X-ray reflection intensities, starting at 200°C and becoming marked at 3OO0C, which they attributed
a,
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to partial structural rearrangement to a chlorite-like phase. The end-product is thus the same as for smectite, and the overburden thicknesses required for dehydration are likely to be comparable t o those for smectite, but it is less clear what are the intermediate structural states in nature that laboratory experiments must reproduce.
Zeo 1ites Sedimentary formations that contain significant amounts of zeolites are deposited in a number of relatively restricted geologic environments. The most important sedimentary zeolite minerals are analcime, clinoptilolite, heulandite, laumontite and phillipsite, and the dominant reaction is the alteration of volcanic tuff by alkaline solutions, whether in deep-sea deposits, saline, alkaline lake deposits and soils, or thicker, volcanic-rich sedimentary piles (Hay, 1966, 1978; Ikjima and Utada, 1966). Meteoric water moving downward through the last-named units typically forms smectite as an alteration mineral until solution salinity and pH have increased to the point where alteration is principally to zeolites. Several of the zeolite dehydration equilibria have been delineated recently in simple aqueous solution, using classic hydrothermal techniques. Equilibbrium temperatures for the simple dehydration: analcime
* albite + nepheline + fluid
range from 492 ? 5°C at 50 MPa (0.5 kbar) to 598 f 5°C at 300 MPa (3kbar), whereas those for: analcime
+ quartz + albite + fluid
range from -200°C at 200MPa ( 2 kbar) t o 183 k 5°C at 500MPa (5 kbar) (Liou, 1971a). A zeolite dehydration reaction that yields laumontite: stilbite
* laumontite + 3 quartz + 3 H 2 0
has equilibrium temperatures ranging from -170'C at 200 MPa ( 2 kbar) to 183 5 10°C at 500 MPa ( 5 kbar) (Liou, 1 9 7 1 ~ )One . of the possible reactions that destroys laumontite on further heating is: laumontite
* wairakite + 2 H 2 0
for which equilibrium temperatures range from -230°C at 50 MPa (0.5 kbar) to 325 5°C at 600MPa (6 kbar) (Liou, 1971b). The typical association of authigenic albite with quartz in sedimentary and low-grade metamorphic terrains suggests that the reaction between analcime and quartz, rather than the simple dehydration at higher temperatures, is the geologically important one (Liou, 1971a). This author notes that in nature the reaction may take place at even lower temperatures because of several factors: amorphous silica rather than quartz may be present in sediments, the plagioclase produced may be well-ordered and multi-cation in contrast with the highly-disordered albite produced in experimental runs,
*
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pore fluids will have significant salinities, hydrostatic pressure may be less than lithostatic. Thus, in some natural environments albite may indeed be stable at 25OC (Campbell and Fyfe, 1965; Hay, 1966). The first authors had calculated that, if the pore fluid were saturated NaCl solution, the equilibrium dehydration temperature would be -lOO°C lower than in simple solution, and even lower in, e.g., chloride brines containing Ca2+and Mgz+as well as Na’. Coombs et al. (1959) had estimated that the equilibrium dehydration temperature would be lowered by -2OOOC if hydrostatic pressure were a third of lithostatic, rather than equal to it. The occurrence of zones along joints and minor faults in the Taveyanne sandstones of the western Alps, in which the laumontite of this unit has been dehydrated and formed other minerals, is taken by Coombs (1971) as a plausible indication of higher hydrostatic pressure within the sandstone than in the joints. Calculations that attempt to reproduce the observed vertical successions of zeolite minerals seen in various sedimentary pi1.es (e.g., Miyashiro and Shido, 1970) typically assume that hydrostatic and lithostatic pressures are equal and that the pore fluid is pure water. The broad features of the successions can be predicted, but there is a “bewildering overlap of zeolite subfacies” (Coombs, 1971). In addition t o the complexities already noted, it should be recalled that zeolites occur in an unusually rich variety of crystal structures, with several different major cations, extensive cation solid solution, significant Si/A1 ratio differences in different structures, and frequent growth and persistence of metastable phases. The bulk chemical composition of the parent-rock is important in determining which zeolite mineral forms, and the Pcoz of the diagenetic environment may determine whether a nonzeolitic clay-carbonate assemblage persists (Zen, 1961). Zeolite minerals undoubtedly release water of hydration when buried in thick sedimentary piles. Predicting the time and depth of particular releases would require an unusually large amount of information, both about mineralogy and geochemistry and about the geologic history of a particular locality.
CONCLUSIONS
The incorporation of appropriate mineral dehydration reactions into a particular hydrological model involves not only mineralogical and geochemical relationships, but also a good deal of regional geologic history. The onset of dehydration is sensitive to geothermal gradient, sedimentation rate, porefluid salinity, and degree of overpressuring, values for all of which must somehow be estimated throughout a geologic time interval that begins with the deposition of the stratigraphic interval of interest. Some of the hydrates such as the zeolites and halloysite are found only in particular geologic facies. Some, such as analcime and smectite, will follow different decomposition
309
reactions and therefore lose water at different temperatures, depending upon whether certain other minerals are present locally in adequate amount for chemical reaction. This situation should lead to stimulating research collaboration between hydrologists and geochemists. However, a process such as mineral dehydration that is strongly dependent upon geologic detail will not be as useful in discriminating among competing models as would a substantially independent processs such as radioactive decay. The discussions of individual hydrates in this paper relied principally upon experimentally measured solubilities and mineral interconversions and upon field observations as criteria for defining thermodynamic stability fields. Thermochemical measurements, which were utilized from time to time in these calculations, could of course form the entire basis of the argument if all the values needed were available (see, e.g., Bmton and Helgeson, 1980). The latter approach is particularly attractive for hydrates that occur as large, well-formed crystals, such as the zeolites.
ACKNOWLEDGMENTS
The opinions presented in this paper are an outgrowth of a continuing research program on transport processes in diagenetic environments funded by the U.S. Army Research Office and the National Science Foundation (Geochemistry Program). Comments by John Hower have been particularly helpful.
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313 Parham, W.E., 1969a. Formation of halloysite from feldspar: low temperature, artificial weathering versus natural weathering. Clays Clay Miner., 1 7 : 13-22. Parham, W.E., 196913. Halloysite-rich tropical weathering products of Hong Kong. In: L. Heller (Editor), Proccedings International Clay Conference, Tokyo. Israel Universities Press, Jerusalem, 1:403-416; 2: 104-105 (discussion). Perry, E.A. and Hower, J., 1970. Burial diagenesis in Gulf Coast pelitic sediments. Clays Clay Miner., 18: 165-178. Perry, E.A. and Hower, J., 1972. Late stage dehydration in deeply buried pelitic sediment. Am. Assoc. Pet. Geol. Bull., 56: 2013-2021. Plumley, W.J., 1980. Abnormally high fluid pressure: survey of some basic principles. Am. Assoc. Pet. Geol. Bull., 64: 414-430. Posner, A.M. and Quirk, J.P., 1964. Changes in basal spacing of montmorillonite in electrolyte solutions. J. Colloid Chem., 19: 798-812. Powers, M.C., 1959. Adjustment of clays t o chemical change and the concept of the equivalence level. In: A. Swineford (Editor), Proceedings Sixth National Conference Clays and Clay Minerals. Pergamon, New York, N.Y., pp. 309-326. Powers, M.C., 1967. Fluid release mechanism in compacting marine mudrocks and their importance in oil exploration. Am. Assoc. Pet. Geol. Bull., 51: 1240-1254. Quaide, W.L., 1956. Petrography and clay mineralogy of Pliocene sedimentary rocks from the Ventura Basin, California. Ph.D. Thesis, Department of Geology and Geophysics, University of California, Berkeley, Calif. Reynolds, R.C. and Hower, J., 1970. The nature of interlayering in mixed-layer illitemontmorillonite. Clays Clay Miner., 18: 25-36. Rowland, R.A., Weiss, E.J. and Bradley, W.F., 1956. Dehydration of monoionic montmorillonite. In: A. Swineford (Editor), Proceedings Fourth National Conference Clays and Clay Minerals. Natl. Acad. Sci.-Natl. Res. Counc., Publ. No. 456, pp. 85-95. Roy, R. and Osborn, E.F., 1954. The system A1203-Si02-HzO. Am. Mineral., 39: 853-885. Roy, R. and Romo, L.A., 1957. Weathering studies, 1. New data on vermiculite. J. Geol., 65: 603-610. Schoen, R. and White, D.E., 1965. Hydrothermal alteration of GS-3 and GS-4 drill holes, main terrace, Steamboat Springs, Nevada. Econ. Geol., 60: 1411-1421. Slabaugh, W.H., 1955. Heats of immersion of some clay systems in aqueous media. J. Phys. Chem., 59: 1022-1024. Steckler, M.S. and Watts, A.B., 1978. Subsidence of the Atlantic-type continental margin off New York. Earth Planet. Sci. Lett., 41: 1-13. Steiner, A., 1968. Clay minerals in hydrothermally altered rocks at Wairakei, New Zealand. Clays Clay Miner., 16: 193-213. Stewart, F. H., 1963. Marine evaporites. U.S. Geol. Surv., Prof. Pap. 440-Y, 52 pp. Stone, R.L. and Rowland, R.A., 1955. DTA of kaolinite and montmorillonite under water vapor pressures up t o six atmospheres. In: W.O. Milligan (Editor), Proceedings Third National Conference Clays and Clay Minerals. Natl. Acad. Sci.-Natl. Res. Counc., Publ. No. 395, pp. 103-116. Stul, M.S. and Mortier, W.J., 1974. The heterogeneity of the charge density in montmorillonites. Clays Clay Miner., 22: 391-396. Tardy, Y. and Garrels, R.M., 1974. A method of estimating the Gibbs energies of formation of layer silicates. Geochim. Cosmochim. Acta, 38: 1101-1116. van Moort, J.C., 1971. A comparative study of the diagenetic alteration of clay minerals in Mesozoic shales from Papua New Guinea, and in Tertiary shales from Louisiana, U.S.A. Clays Clay Miner., 19: 1-20. van Olphen, H., 1963. Compaction of clay sediments in the range of molecular particle distances. In: W.F. Bradley (Editor), Proceedings Eleventh National Conference Clays and Clay Minerals. Pergamon, New York, N.Y., pp. 178-187. Velde, B., 1969. The compositional join muscovite-pyrophyllite at moderate temperatures and pressures. Bull. SOC. Fr. Minbral. Ckistallogr., 92: 360-368.
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315
HfiTfiROG~NfiITfiCHIMIQUE ET HYDROLOGIQUE DES EAUX SOUTERRAINES D’UN KARST DU HAUT-JURA NEUCHATELOIS, SUISSE
YVES BOUYER et BERNARD KUBLER
Laborato ire d e Miniralogie, Pk trograph ie et Giochimie, Institut de Gkologie, Universiti d e NeuchGtel, CH-2000NeuchGtel 7 (Suisse) (Accept6 pour publication le 9 juin, 1981)
ABSTRACT Bouyer, Y. and Kubler, B., 1981. H&t6rog6n&itkchimique et hydrologique des eaux souterraines d’un karst du Haut-Jura neuchitelois, Suisse. (Chemical and hydrological heterogeneity of groundwaters in the Upper “Jura Neuchitelois”, Switzerland.) In: W. Back and R. Lbtolle (Guest-Editors), Symposium of Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54: 315-339. Chemical analyses were made of waters sampled in a valley of the folded Jura Mountains at three different points: (a) the surface, as running water; ( b ) four bore holes in the karstic zone, and ( c ) a karstic spring which collects (a) in (b). The chemical parameters have been treated by factor analysis according to Benzbcri, by comparing them with flow and precipitation. The results of the analysis, after occasional blind sampling and data treatment, show the individuality of the chemical composition of each bore hole and the relationship between the surface flow and karstic spring. The linear correlation coefficients have confirmed the difference between the bore holes and distinguished the relationship between chemical component, flow and precipitation. The success of these comparisons is due to the particular type of water of this specific syncline of the folded Jura where the highly-acid running water, drained from bogs, changes to hard water when it passes through the underlying limestone. The importance of the soil for the chemical composition of the groundwater is underlined.
RESUME Dans un synclinal du Haut-Jura (vall6e des Ponts), des analyses chimiques des eaux ont pu &re faites aux points suivants: (a) 6coulement superficiel; (b) quatre forages dans les calcaires; et (c) rbsurgence de ces eaux (source de la Noiraigue). Les param6tres chimiques ont &6 trait& par l’analyse factorielle des correspondances, selon Benzbcri, en les comparant avec les d6bits et les pr6cipitations. L’bchantillonnage aveugle dans le temps et dans l’espace et le traitement des donnies ont permis de mettre en 6vidence d’une part l’individualitb de la composition chimique de chaque forage et, d’autre part, la parent6 entre l’ckoulement superficiel et la rbsurgence. Les coefficients de corrglations lin6aires ont confirm6 les diff6rences entxe les eaux des divers forages et prbcisi les relations entre composition chimique, dbbits des ruisseaux et de la source, variation des niveaux pi6zom6triques et prkcipitations. Le succes de ces comparaisons est dii a la qr.alit6 des eaux particulihres de ce synclinal du Haut-Jura neuchiitelois qui passe des eaux de marais acides aux eaux dures percolant dans les calcaires. On met ainsi en bvidence le role d6terminant de la zone de marais sur la composition chimique de ces eaux souterraines.
316 INTRODUCTION
Les eaux souterraines du karst du HautJura neuchatelois ont ete etudiees au siecle passe dejB par Desor (1864), Jaccard (1883) et au debut de ce siecle par Schardt (1904) qui, par ses observations minutieuses, nombreuses et pertinentes, peut &re considere comme un des fondateurs de l’hydrogeologie. Depuis, Burger (1959), pour le bassin de l’Areuse, Tripet (1972), pour le bassin de la source de l’Areuse proprement dite et Morel (1976), pour le bassin de la Noiraigue qui fait l’objet de cet article, ont considerablement ameliore la connaissance hydrogeologique de ce karst. Mais c ’est surtout grace au memoire de Miserez (1973) que la gkochimie d e ces earn est mieux connue. Dans le cadre d’un projet d u Fonds National Suisse de la Recherche Scientifique du Laboratoire de Geochimie de I’Universite de Neuchatel (B. Kubler) et d’une operation concertee de la D.G.R.S.T. (M. Millot, Strasbourg) il a pu, avec la collaboration de Pochon (1978), confirmer l’importance et la nature de la couverture vegetale sur le chimisme. En utilisant les resultats de Persoz et Kubler (1973) il a montre I’influence des roches encaissantes sur l’augmentation des sulfates, des alcalins et de Sr. I1 a enfin presente quelques beaux exemples de relations directes ou inverses entre debit et ions dissous. Cependant de nombreuses questions restaient B resoudre, comme le transport par suspension, qui obligent immediatement A tenter de preciser l’heterogeneite du karst dans l’espace et le temps. Kiraly et Muller (1979) resument les grands traits de cette heterogeneite. Se referant A Schoeller, (1967), Drogue (1971), Tripet (1972), Mangin (1975) et Kiraly (1969,1973, 1975, 1978), ils distinguent un: “r&eau de drainage connexe tr&spermiable et des blocs capacitifs peu permiables, de volumes relativement importants”
ils admettent, avec Miserez ( 1973) une grande heterogeneite de l’alimentation et precisent que les maximums des crues ne peuvent s’expliquer que par une infiltration rapide, telles les pertes des eaux de surface dans les emposieux. En fait, le spectre de distribution des vitesses mesurees (au Centre d’Hydrogeologie de Neuchatel) par les traqages va de 800 m/hr. (cas exceptionnels) B 2 m/hr. (Kubler, 1972). Dans les sols, ces vitesses sont beaucoup plus petites: de 0,06 A 0,30 m/hr. (Boulaine, 1971). Dans les blocs capacitifs, des vitesses “raisonnables” de 0,2 B 2 m/hr. pourraient &re admises (Kiraly et Muller, 1979). Du reste ces vitesses dependent de l’amplitude des mises en charges, donc de l’importance des precipitations par rapport aux evenements antecedents. Ce phenomene, bien connu dans les sols sous le nom de ressuyage, s’exerce aussi dans le karst, et les suspensions que l’on peut observer au moment des crues en sont une preuve (Pochon et Simeoni, 1976). La notion de nappe en regime karstique est alors difficile B preciser et nous l’assimilerons A celle de blocs capacitifs et celles des reseaux de haute permeabilite B celles de l’exutoire en dehors des etiages. Comme les blocs
317
capacitifs occupent la majorite du volume du karst, il n’est des lors pas etonnant que les forages de reconnaissance d’eau ne tombent que tr&srarement dans le reseau de haute permeabilite. Or, comme chaque bassin karstique a son individualite, il est admissible que chaque bloc, meme avec des reactions hydrauliques peu differentes, posskde son propre chimisme. C’est ce que nous tenterons d’illustrer par les exemples du bassin de la No iraigue.
Le cadre hydrogeologique (Fig. 1) D’apres Morel (1976), la vallee des Ponts-de-Martel, qui forme le bassin alimentaire de la Noiraigue, a les caracteristiques qui sont donnees dans le Tableau I. Le synclinal peut etre divise en quatre ensembles calcaires
Fig. 1. Position des piezom&res: 1 4 a 17, des points d’eau de surface de 1 perte principale du Bied : 1 0 et de la resurgence de la Noiraigue: 13.
12, de la
Fig. 1 . Location of piezometers 14-1 7; sampling localities of surface waters 1-12, the main sink is at the Bied (10) and the karst spring at the Noiraigue ( 1 3 ) . TABLEAU I Caracteristiques hydrogeologiques de la vallee des Pontsde-Martel Superficie (km2 ) Altitude moyenne (m) Resurgence ( m ) Altitude maximale ( m ) For& (km’ ) Prairies et piturages (km’ ) Tourbibres ( km2 ) Volume des roches noyees (m3 ) Volume d’eau ( m 3 ) Temphature moyenne (OC) Prkcipitations moyennes annuelle (mm) Precipitations centenaire s k h e (mm) Prkcipitations centenaire humide (mm) Debit resurgence, moyen (m3/s) Debit pour des precipitations moyennes (mm)
69 1.065 740 1.337 19 36 14 9 109 46 * l o 6 5 94 468 870 2.191 de 1,18 2,88 de 523 a 1250
-
a
318
-
P e t i t Cachot innn
-
a
Martel Dernier .. . .
%Sourcede l a
,.’.
%/ Noiraigue
~
7 Areuse
ALL.GLACIAIRE8
PORTLRhOIEN
HAUTERIVIEU
KIMMERIDGIEN SEQUANIEN
VALANGINIEN
ARGOVIEN
PURBECKIEN
CPLLOVIEN
0 ~
\
1
-
-
2 Km
A C C I D E U T TECTONIQUE
Fig. 2. Profil gkologique simplifik du synclinal des Ponts-de-Martel passant par la rksurgence de la Noiraigue (d’apres Morel, 1976). Fig. 2. Schematic geological cross-section of the Pontsde-Martel geosyncline traversing the karst spring at the Noiraigue (after Morel, 1976).
separes chaque fois par des marnes: le Dogger, le Malm, 1’Hauterivien et le Valanginien: “Pierre jaune de Neuchatel”. Par endroits il subsiste des lambeaux de molasse, ailleurs depots quaternaires et tourbks occupent la surface (Figs. 2 et 3). Les prel&uements,analyses, parametres physiques et chimiques
Dix-sept points d’eau ont ete echantillonnes chaque mois, sur deux ans (1974--1975) recouvrant les quatre saisons: les douze premiers points, dans des eaux de surface, le treize dans la resurgence de la Noiraigue, les quatre derniers dans les piezom6tres donnes dans le Tableau 11. Debits de la resurgence et precipitations ont ete mesures par l’Office Federal de l’ficonomie Hydraulique et calcules par L. Kiraly du Centre d’Hydrogeologie de 1’Institut de Geologie de Neuchatel. Sur les eaux elles-memes, ont ete mesures: Ca2+, Mg2+,Fe2+--Fe3+,Sr2+,Na’, K+, SO:-, SiO, , NO,, C1-, O2 dissous, conductibilite, temperature, durete totale (DuTo), durete temporaire (DuTe) et mati6res organiques dissoutes (MO). Pour les essais de correlations entre parametres chimiques et parametres physiques, et pour se degager de tout B priori theorique, donc pour mieux TABLEAU I1 Les quatre pi&zom&respour comparaisons avec la rksurgence No.
Nom
Aquifere
Altitude pikzomitrique (m)
14 15
17
Petit-Martel
Dogger Malm Hauterivien i Valanginien i Malm Molasse Malm
1.140 840
16
Combe des Cugnets Martel-Dernier Brot-Dessus
741 7 49
319
LES CUGNETS
Q Bj -1100
MARDERN -1000 m. .
Q
Se
- 900 2.1.
Po - 800 Ki
- 700
- 600 FORMAT I ON>
-
Q
qudternalre
I,
vdlanginien
K1
klmeridgien
Mo
volilsse
PU
perbeckien
Se
sequanien
p
hdutErlvien
Po
portlandien
1.1
Zone lanilnee
Fig. 3. Profil lithologique et stratigraphique des pi&zom&res(in: Morel, 1976). Fig. 3. Lithologic and stratigraphic profile o f the piezometers (from Morel, 1976).
traiter les faits on a essaye de comparer precipitations ( P ) et debits ( D ) de la mani&redonnee dans le Tableau 111.
TRAITEMENT MATHEMATIQUE DES DONNEES
On a applique, au traitement de ces nombreuses donnees, l’analyse factorielle des correspondances (AFC) selon Benzecri (1970), Benzecri et al. (1973) que Bobke et al. (1977), et Lachance et al. (1979) ont utiliske avec succ&s aux probkmes de caracterisation des eaux. Le traitement a ete
320 TABLEAU I11 Les types du relev6 journalier des dkbits ( D ) et des prkcipitations (P) DI, D2, D3, D4, D5, D6, D7,
PI P2 P3 P4 P5 P6 P7
du lendemain de l’dchantillonnage du jour m6me du jour pr6c6dant le jour du prdhement moyenne des 26me et 36me jours prdc6dant le jour du pr616vement moyenne de la semaine pr6c6dant le jour du prdl6vement moyenne des 2&meet 36me semaines pr6ckdant le jour du pr6livement moyenne des 4kme, 5&meet 6kme semaines prdckdent le jour du prdl6vement
applique en plusieurs fois, permettant de distinguer des groupes et des sousgroupes. En complement et, pour chaque point d’eau, les coefficient de correlation lineaire ont 6t6 calcules entre les differents parametres.
Groupements d e toutes les eaux: surface, piezom6tres et exutoires par AFC Dans les eaux d e surface, le Bied, numero 10, echantillonne juste avant sa perte, a ete specialement identifie sur les figures; les autres eaux de surface viennent de marais draines et de sols sous pelouse. Dans ces premibres analyses des correspondances on a tenu compte de dix-sept points d’eau, seize variables (Ca2+, Mg2+, Fetot, Sr2+,Na’, K+, SO$-, SiO,, NO,, C1-, O,, H+, PO:-, DuTo, DuTe), et matieres organiques dissoutes (MO). Les plans des axes 1 et 2 (Fig. 4A), 3 et 4 (Fig. 4B) absorbent respectivement 66 et 19% de la variance totale. Dans le plan des axes 1 et 2, toutes les eaux sont regroupees en deux tendances qui se recoupent en un pole commun: PO:-, SO$- et MO. Les deux autres poles sont d’une part OH- et d’autre part H’, celles des piezometres la branche OH-. Les eaux de BrotDessus (No. 16) et de Petit-Martel(17) ont en commun surtout K+, Na’, C1-, Mg2+. Les eaux de la Noiraigue (13) et du Bied (10) ont en commun Ca2+, DuTo, DuTe et Fe2+, et sont bien groupees. Martel-Dernier (15)s’ecarte des autres groupes. I1 est en position intermediaire entre le groupe des eaux 10, 13 et 1 4 et l’intersection des tendances; il est proche des SiO, et NO;. Dans le plan des axes 3 et 4 (Fig. 4B) deux tendances apparaissent: (1)de Martel-Dernier (15) et NO;; (2) de Petit-Martel (17) et OH-, C1- et Fetot. A l’intersection de ces deux tendances se trouvent les eaux de surface, de la perte (emposieu), de la resurgence (exutoire) et des Cugnets (14). Les eaux de Brot-Dessus (16) sont intermediaires entre celles de la perte, de la resurgence et de Petit-Martel. Elles sont regroupees avec Mg2/ Na’, K+ et Sr 2+. De ce premier traitement, grace ii un echantillonnage aveugle sur plus de 24 mois, on doit constater: (1)Une parente etroite entre les eaux de la Noiraigue (resurgence No. 13) et du Bied (perte ponctuelle No. 10). (2) Une grande difference entre les eaux des piezometres et les eaux de surface aussi bien des marais acides que des marais elargis.
321 T
-1.06
..
_ _-3,07 __i Axe I
\e,
I
I
1'
\
Fig. 4.Plans des axes: (A) 1 et 2; et (B) 3 et 4. Dans le plan des axes factoriels, contributions des observations (petits points) et des variables (gros points); analyse des correspondances selon Benzkcri et al. (1973). Fig. 4.Correlation surfaces: (A) between axes 1 and 2; and (B) axes 3 and 4. Loadings of the objects (small dots) and of the variables (large dots) are shown in the surfaces through the factor axes, correlation analysis after Benzkcri et al. (1973).
322
( 3 ) Une meme tendance pour les eaux des piezometres de la resurgence et de la perte. (4) Une separation en deux groupes pour les piezometres, d’une part celui de Martel-Dernier ( 1 5 ) ,d’autre part celui de Petit-Martel (17). Les eaux de Brot-Dessus etant intermediaires entre 17 et le groupe perte-resurgence. Pour differencier les eaux souterraines, une seconde serie d’analyses des correspondances a ete faite en eliminant toutes les eaux de surface. Groupement par AFC des eaux souterraines: resurgence et piezometres On a traite, dans cette serie d’analyses des correspondances, les mgmes parametres physiques et chimiques mais sur les eaux des quatre piezometres et de la resurgence uniquement. Les axes 1, 2, 3 , 4 , 5 absorbent successivement 43, 23, 17, 7 et 5% soit le 95% de la variance totale. Dans le plan des axes 1 et 2 (Fig. 5) on distingue de nouveau deux tendances et leur intersection. La premiere tendance est celle des eaux de Martel-Dernier ( 1 5 ) (NO, et SO$-) et toutes les analyses d’eau de ce piezometre s’alignent selon cette tendance depuis le NO, (pole de la tendance) jusqu’au groupe central (intersection). Le groupe central, B l’intersection des deux tendances, regroupe la resurgence ( 1 3 ) ,DuTo, DuTe, Ca2+,SiO, , PO$-, H+ et 0, mais aussi les eaux du piezometre des Cugnets ( 1 4 ) dans l’aquifere du Dogger. La seconde tendance est jalonnee par toutes les analyses des eaux de BrotDessus (16) et celles de Petit-Martel ( 17) qui, avec Fetot, forment le pale de cette tendance. Donc grace B ce traitement par AFC des eaux souterraines, seules les tendances illustrees par les axes 3 et 4 (Fig. 4B), lors du traitement de toutes les eaux, sont confirmees et precisees, c’est-h-dire individualisation des eaux de Martel-Dernier ( 1 5 , tendance 1)’ de celles de Petit-Martel (17, tendance 2), convergence vers le groupe d’intersection represent6 par les eaux de la resurgence ( 1 3 ) , celles de la nappe du Dogger (14).Brot-Dessus ( 1 6 )occupe une position intermediaire entre Petit-Martel ( 17) et la resurgence (13 ) . Cependant le groupe, B l’intersection des deux tendances, n’est homogene que dans le plan des axes 1 et 2 (Fig. 5A). Dans celui des axes 1 et 3, ces deux eaux s’opposent surtout par Fetot. Les composantes (valeurs absolues) des Cugnets (14),sur l’axe 3, sont toujours plus grandes que celles de la resurgence ( 1 3 , voir Fig. 6). Cela reflete fidelement l’observation que les concentrations en Fetot sont toujours plus grandes dans les eaux du Dogger (Cugnets, 14) que dans celles de la resurgence ( 1 3 ) . Les petites differences entre les eaux de Petit-Martel (17) et Brot-Dessus ( 1 6 ) sont explicitees par l’axe 4 dans le plan des axes 1 et 4 (Fig. 5b). Les eaux de Petit-Martel(17) se distinguent par de faibles concentrations en Ca2+,DuTo, DuTe, SiOz et H+ et un enrichissement relatif en K+, Na’ et C1-, celles de Brot-Dessus par un enrichissement relatif en Mg2+, Sr2+, et SO$-. Toutes les donnees de la resurgence (13 ) placent ces eaux au milieu de ces oppositions.
323
1,147 I I I
Axe2
I I
I
Fe
-0'99
-0,46
Fig. 5. A. Plan des axes 1 et 2 en ne conservant que les eaux des piCzorn6tres et de la resurgence. B. Plan des axes 1 et 4 en ne conservant que les eaux des pikzom6tres et de la r6surgence. Fig. 5. A. Surface through axes 1 and 2 taking into account only soil water at the piezometers and karst water. B. Surface through axes 1 and 4 taking into account only soil water at the piezometers and karst water.
La distinction entre les eaux de la resurgence et des piezombtres etant operee, il s'agit des lors d'eliminer du traitement des donnees les eaux de la resurgence.
Essai de distinction, par AFC des regroupements duns les eaux des piezomktres uniq uement Dans les exemples ci-dessus l'analyse des correspondances a et6 appliquee sur les valeurs brutes. Pour mieux tenir compte de l'aspect quantitatif, la
324
I
I I I I I
I
I I
I I
I I
I I
I
Fig. 6. Plan des axes 1 et 3 en ne conservant que les eaux des piizom6tres et de la rksurgence. Fig. 6. Surface through axes 1 and 3 taking into account only soil water at the piezometers and karst water.
variation de chaque parametre a ete divisee en trois classes. Les limites des classes ont ete etablies & partir des histogrammes, selon un codage logique, sous forme disjonctive complete (Benzecri et al., 1973; Bob6e et al., 1977). Pour chaque parametre la classe faible est affich6e: 1,moyenne: 2, et forte: 3 . Cependant pour mieux 6tablir la distinction entre les eaux on a tenu compte, en plus des parametres physico-chimiques habituels, des parametres complementaires comme le fer filtre", le rapport du fer filtre" au fer total (FeFI/FetOt), des precipitations PI A P7, de la profondeur des prelevements et de la temperature. Le nombre total de variables atteint ainsi 27 et, par le codage logique, & chaque pikzometre correspond un tableau de donnees de 3 x 27. Apres les premiers traitements exploratoires il est apparu qu'il fallait eliminer SO$- (3), Si02 ( 3 ) , conductibilite ( 2 ) , pH (I), PO:- ( 3 ) ,DuTe (1) et MO (1, 2 et 3 ) , soit parce qu'ils n'apparaissaient jamais dans les eaux des piezometres, soit parce qu'ils etaient toujours en faible quantite.
* On entend sous le terme de fer filtr6 (FeFI), toutes les esp&es de fer passant un filtre membrane de 0,45pm.
325
Axe 1
c>
-1J1
Fig. 7. Plan des axes 1 et 2 des observations des eaux des pi&zom&res uniquement: chaque pi&zom&re, d'apr& sa composition chimique, occupe une aire distincte et non s&cant e (observations u niq uement ). Fig. 7. Surface through axes 1 and 2 only loaded with the objects for soil water at each of the piezometers; the data of each piezometer, according to the en lieu chemical composition, occupy a distinct area which does not intersect other areas consisting of objects only.
Les eaux des quatre piezometres se distribuent en quatre groupes bien separ6s (cf. Fig. 7) (observations) et Fig. 8 (variables). (1)Les eaux des Cugnets (14), aquifere du Dogger par: Sr2+moyen, profondeur faible & moyenne, Mg2+moyen, T ("C) moyen et NO, moyen. (2) Les eaux de Martel-Dernier (15) par: le rapport Fe,,/Fe,,, et fort, DuTe moyenne, SO',- moyen, Mg2+ faible mais surtout NO, et O 2 fort. (3) Les eaux de Petit-Martel ( 1 7 ) : par une DuTo et 0 2 ,Ca2+, conductibilite et Si02 faibles, par C1- et K+ moyens et surtout Na' et Mg2+eleves. (4)Les eaux de Brot-Dessus (16)forment en quelque sorte le centre de gravite des trois autres groupes.
Les resultats d u traitement par I'analyse des wrrespondances Sur la base d'une echantillonnage aveugle (une prise d'eau par mois sur deux ans & 27 mois), de 1 7 & 27 parametres physico-chimiques, donc quelle que soit la saison et dans celle-ci les episodes meteorologiques et hydrodynamiques, le traitement par l'analyse des correspondances selon Benzecri nous permet les conclusions provisoires suivantes:
326
I
Fig. 8. Plan des axes 1 et 2 des variables dans les eaux des piizom6tres uniquement. Chaque variable (par exemple, Febt, DuTo, etc.) a 6t6 divis6e en trois classes 6 partir des histogrammes des valeurs, selon un codage logique sous forme disjonctive compl6te. Fig. 8. Surface through axes 1 and 2 only loaded with the variables of the soil waters at each of the piezometers. Each variable (for example, Febt, total hardness (DuTo), etc.) is classified into three groups according to data frequency analysis with logical and total subdividing into three groups.
Le chimisme global et moyen (sur deux ans) des eaux de la resurgence de la Noiraigue est tres proche de celui du Bied confluent de tous les bieds de surface: groupe quaternaire de surface. Le chimisme des eaux de surface accuse de grandes differences (pH, durete, etc.). Des marais acides aux marais elargis et aux anciens marais drafnes et cultives, ce chimisme converge vers celui du Bied et vers celui de la resurgence. Les eaux des piezometres s'ecartent notablement de celles de surface et de celles de la resurgence. Cependant, fait plus important, chaque piezometre se distingue par le chimisme different de ses eaux. Cela confirme l'heterogeneit6 d u chimisme
327
des blocs capacitifs A faible permeabilite. Dans le cas du Bassin de la Noiraigue et avec les quatre piezometres seulement on devrait distinguer quatre blocs capacitifs.
COMPARAISON SUCCINTE ET SEMI-QUANTITATIVE DES PRECIPITATIONS, NIVEAUX PIEZOMETRIQUES E T DEBITS DE L’EXUTOIRE
Les hydrogrammes des precipitations, des debits B la resurgence et des quatre niveaux piezometriques sont reproduits B la Fig. 9. Les limnigrammes des piezometres, veritables empreintes digitales de la reaction des blocs capacitifs 8 la mise en charge et B la vidange, seront decomposes plus tard (Bouyer, en prep.) en plusieurs exponentielles selon les conseils de Kiraly et Muller (1979). D’un point de vue plus terre B terre, le limnigramme du piezom6tre des Cugnets (14) accuse un battement du niveau de plus de 48 m, une remontee tres rapide en un ou deux jours, une lente redescente en 28 B 32 jours et une stagnation du niveau le plus bas en periode d’etiage. En dehors des periodes de fonte des neiges, la reaction est du jour meme ou du lendemain. Le niveau piezometrique de Martel-Dernier (15 ) presente un battement de 1 9 B 20 m, des decroissances exponentielles rapides d’une dizaine de jours. Cependant toutes les “pointes de crue” sont ecretees B la cote constante de -172 m, c’est-&-direB l’altitude de 8 5 1 m. Toute eau en exces, par rapport B cette cote, est sans doute evacuee par des systemes de trop-plein. Dans ces circonstances il est difficile d’apprecier le decalage entre les precipitations et les altitudes piezometriques, il semble toutefois qu’un jour les &pare. Petit-Martel et Brot-Dessus (respect. 17 et 16), contrairement aux autres forages, ne presentent de paliers ni dans les altitudes elevees, ni dans les basses. Le battement est plus eleve: 45 m et la decroissance exponentielle plus lente B Petit-Martel (17 ) qu’8 BrobDessus oh le battement n’est que de 24m. Dans les deux cas l’augmentation des niveaux jusqu’B leur maximum est du lendemain, du surlendemain ou de trois jours apres les precipitations. On note, par contre, une concordance chronologique du jour meme entre les hydrogrammes des deux piezometres. La determination des permeabilites, par essais d’injection, n’a ete pratiquee que sur Martel-Dernier (15) (Simeoni, 1971). Dans la zone desaturee les valeurs sont souvent plus grandes que 6 lo-’ m/s, alors que dans la zone saturee elles restent souvent inferieures B cette limite. Les permeabilites sont donc tres faibles. Leur ordre de grandeur explique l’allure des variations des niveaux piezometriques mais confirment la notion de blocs capacitifs et l’originalite chimique de chacune de ces nappes. La montee de l’exutoire ( 13) se correle avec les precipitations de la veille, voire du jour meme. La decroissance des debits est rapide. Ces faits sont le reflet d’une participation importante de l’alimentation ponctuelle (Kiraly et Morel, 1976; Kiraly et’Muller, 1979). D’aprbs les premiers calculs de bilan,
-
328
400 I 300 I’ETIT hl,\RTCL 200
,
100
NOIR41CIJE
80
o e b j t jrwrnalier
60
(m3ir.)
40
20
- 0
-220
I
I
Fig. 9. Tableau synoptique: des prbcipitations journaliGres, des d6bits journaliers de la Noiraigue et des fluctuations des niveaux piCzom6triques. Fig. 9. Synoptic table: daily rainfall and daily runoff at the Noiraigue and the variation o f the piezometers heads.
Morel (1976) estime que la perte du Bied (10) contribue pour 25% au volume annuel de la resurgence. Ces proprietes hydrauliques particulieres, m$me au stade semi-quantitatif de l’interpretation, montrent une relation directe et rapide entre les eaux de surface et l’exutoire, plus lente entre ces eaux et les niveaux pic5zomc5trique.s et une individualite de reactions des piezometres. I1 n’est des lors pas
329
etonnant que par analyse des correspondances, les eaux de chaque piezometre forment un groupe physico-chimique independant et que les eaux de surface, particulierement celles de la perte et celles de la resurgence soient etroitement apparentees. On peut illustrer ces relations plus quantitativement en passant par la matrice des coefficients de correlation.
INTERPRETATIONDES CORRELATIONS, PRECIPITATIONS, DEBITS ET CHIMISME
Lorsque la concentration d’une substance chimique augmente avec le debit, on y voit un “ e f f e t piston”, si au contraire cette concentration diminue, c’est l’effet de dilution. L’AFC permet dejja de separer les groupes d’eau selon ces deux lois de concentration ou de diminution avec le debit, cependant pour avoir une approche plus quantitative, les coefficients de correlation lineaires sur la somme des donnees des dix-sept points d’eau (15.600 valeurs) ont ete calcules. Si Miserez (1973) recommandait de mesurer, si possible en continu, le maximum de parametres sur plusieurs crues, il n’est pas toujours possible d’equiper dix-sept points d’eau de chaines de mesure en continu. Les resultats et interpretations presentes ici, s’appuient sur un enregistrement journalier des precipitations et debits des points d’eau les plus importants, sur l’analyse en continu de quelques crues de la resurgence principale (Bouyer, en prkp.), sur le calcul de correlations canoniques et sur le dessin de graphes X-Y pour tester, surveiller et controler la validite des correlations. Correlations dans les eaux d e surface Aux precipitations et au debit du ruisseau principal: le Bied, ont ete correles tous les parametres physico-chimiques des six petits ruisseaux affluents du Bied (points d’eau 1 ja 6, Tableau IV). Les meilleures correlations sont entre le K+ et les precipitations du jour meme (entre 0,83 et + 0,93, Tableau IV) de meme qu’entre la MO et les m6mes precipitations. Pour les debits, par contre, les meilleures correlations sont entre le K + et la MO et les debits du lendemain. Or c’est aussi le debit du lendemain qui est le mieux correle aux precipitations ( r = 0,88). On explique aisement ces faits. En surface, dans le marais acide et son systeme de drainage, les precipitations lavent le K+ et la MO soluble des sols et des tourbes; ces eaux, enrichies en ces substances, arrivent B la perte principale ou est installe le limnigraphe avec un decalage d’un jour. L’augmentation du K + et de la MO avec les precipitations, est decrit par les pedologues par le phenomene de “ressuyage” qui n’est autre que l’effet piston au niveau des sols et des tourbes. Des que l’on quitte le marais acide, les correlations deviennent mauvaises et les coefficients changent meme de signe, signifiant que K+ et MO ont
+
+
330 TABLEAU IV Les coefficients de corrdlation*’ des param6tres physico-chimiques aux prckipitations (P)*’ et dbbits (D)*’ des six petits ruisseaux affluents du Bied
r ( K+ - P2)
r(K+ - D l )
r ( M 0 - P2)
r(M0 - D l )
0,85 0,91 0,83 0,83 0,82 0,93
0,77 0,82 0,77 0,78 0,74 0,84
(0,341 0,61 0,31 0,66 0,61 0,49
(0,271 0,65 0,52 0,58 0,52 0,49
Pointe d’eau
*l *2
Le niveau de confiance de ces diffirents coefficients de corrdation = 20 = 99,99%. Signification de P et D cf. Tableau 111.
tendance i~ diminuer avec les precipitations et les debits, marquant ainsi des signes de l’effet de dilution.
Correlations eaux d e surface--resurgence A la rksurgence, un examen des corrklations du Tableau V montre que le K + et la MO sont le mieux correles avec la moyenne des precipitations de la 0,63 et 0,56). Pour les relations semaine precedente (respect., r = precipitations-debits de la resurgence, les coefficients de correlation entre precipitations d u jour mkme, du jour precdent ou des 2-3 jours precedents, sont tous mauvais avec n’importe quel debit. 11s deviennent par contre
+
+
TABLEAU V Covariances et corrilations entre param6tres chimiquespricipitations (P)et dibits ( D ) a la source de la Noiraigue
Mg P1 P2 P3 P4 P5 P6 P7
-0,14 0,28 0,24 0,23 -0,16 -0,40 -0,28
D1 D2 D3 P4 ,435 D6 D7
-0,53 -0,55 -0,47 -0,74 -0,78 -0,67 -0,44
Ca
’+
0,27 0,17 0,08
0,07 0,09 0,43 0,48 0,03 -0,09 -0,03 -0,04 -0,05
0,17 0,05
K+
MO
P5
0,43 0,12 0,36 0,52 0,63 -0,lB 0,29
0,13 -0,03 0,05 0,48 0,56 0,07 0,52
1,oo -0,Ol. 0,38
0,40 0,50 0,48 0,39 0,12 -0,lB -0,20
0,36 0,40 0,62 0,46 0,27 -0,Ol -0,07
0,68 0,71 0,69 0,62 0,43 0,05
0,20
331
excellents entre precipitations de la semaine precedente et debits du jour du prelevement. Par ailleurs, les correlations entre K+, MO et O2 d’une part et debits d’autre part sont les meilleures pour K + et debits du jour meme ( r = 0,50) 0’62) et O2 et debits du jour MO et debits du jour precaent ( r = precedent ( r = 0’72). Cela signifie que la concentration de trois substances chimiques augmente statistiquement avec le debit. En m6me temps, si debit K+, MO sont le mieux correles avec la moyenne des precipitations de la semaine precedente, l’interpretation est alors Claire: les eaux du marais acide s’etant chargees en K’, MO et O2 se retrouvent A la resurgence principale selon un retard moyen d’une semaine, du moins pendant les 24 A 27 mois qu’a dure la campagne de prelevement mensuel. Plus exactement la “vague” d’arrivee de ces eaux se correle le mieux avec la moyenne des precipitations des sept jours p r e c a a n t le prelevement. Ces faits sont contraires aux interpretations que nous donnions (Kubler et al., 1978). L’ “effet piston” se produit tout-A-fait en surface dans les tourbes et sols organiques. I1 est impossible A distinguer dans les zones non saturees et saturees. 11 pourrait exister pour les cations mineraux comme le Ca2+, Mg2+, Sr2+, par exemple. Aucune de nos correlations ne permet de le demontrer. L’effet de dilution est reconnu quand la concentration d’une substance diminue avec l’accroissement d u debit, donc quand les coefficients de correlation sont negatifs. Dans le karst jurassien Miserez (1973) avait trouve une loi de dilution entre le logarithme du debit instantane et la concentration, le cation le plus significatif &ant le Mg2+,A la Noiraigue. D’apres nos resultats cette loi s’explique car le coefficient de correlation entre le Mg2+et le debit du jour meme est dejh de 0,55 (Tableau V). Cependant le coefficient atteint son maximum -0,78 avec le debit moyen de la semaine precedant le prelevement. L’effet de dilution est confirme par la conductibilite dont le coefficient de correlation est aussi le plus fort -0’66 avec le debit moyen de la semaine precedente. La composition chimique des eaux de la resurgence est donc largement influencee par les evenements hydrauliques qui precedent le prelevement.
+
+
+
Le cas d u calcium Pour le Ca2+,les coefficients de correlation sont toujours non significatifs tant avec les precipitations, les debits qu’avec la conductibilite. I1 apparait comme un cation non significatif de l’ecoulement des eaux ou le comportement hydrodynamique de la resurgence. CelA s’explique par le fait qu’une partie du Ca2+ n’est pas transportee en solution waie, mais sous forme solide particulaire. En effet, lors de l’examen detaille de crues (Bouyer, en prep.) le Ca2+particulaire se correle m i e n avec le debit que le Ca2+waiment libre. Par ailleurs, aux pH mesures dans le karst, et d’apres les produits de solubilite des sels de Mg2+et des sels de Ca2+, on est beaucoup plus eloigne
332
par sous-saturation de l’equilibre des mineraux magnesiens que des mineraux calciques. Du reste, les mauvaises correlations entre la conductibilite et le Ca2+ et au contraire les bonnes correlations avec le Mg2+( + 0,52) confirment ce fait. Pour retrouver donc les lois qui regissent les teneurs en Ca2+dans les eaux karstiques, un dosage chimique n’est pas suffisant, il faut separer la phase solide particulaire de la solution. La phase solide, dans certaines crues, peut augmenter avec le debit en raison de l’effet-piston alors que dans le meme temps le Ca2+en solution subit probablement l’effet de dissolution comme le Mg2+. I1 faut trouver dans la combinaison de ces deux lois contraires le manque de correlation d u Ca2+avec les autres parametres.
Les correlations dans les eaux des piezomktres: des blocs capacitifs Nous avons vu plus haut que K+, MO et O2 etaient de bons marqueurs de 1’“effet-piston” dans les sols, que le Mg2+ etait un bon marqueur de l’effet de dilution et que d’apres 1’AFC les eaux des blocs capacitifs traverses par les piezometres formaient des groupes separes. Les coefficients de correlation n’ont trouve une signification que pour chaque piezometre. Pour tous les piezometres, la MO hydrosoluble a disparu. Ce fait illustre dejA que d’une part les ecoulements surface-resurgence sont beaucoup plus rapides que les ecoulements dans les blocs capacitifs et que d’autre part la nature des sols o u d e l’absence de sols est determinante pour la composition chimique des eaux karstiques. Cependant, si la MO a disparu dans les piezometres le K + subsiste; on l’utilisera pour detecter 1’“effet-piston” et le Mg2+ pour la dilution. Aux Cugnets (No. 1 4 ) et 51. Brot-Dessus (16) pourtant deux piezometres A regime hydraulique tres different, le K+ subit un effet de dilution. Son coefficient de correlation avec les precipitations moyennes des 2eme et 3eme semaines precedant le jour de prelevement (P6) est negatif,
0,54
No. 14:
(Kf-P6)
=
-
NO. 1 6 :
(K+-P6)
1
-0,69
et
Cet effet de dilution est confirme par les correlations Mg 2+, conductibilite precipitations,
NO. 14:
(Mg2+- P6) = - 0’47
NO. 1 6 :
(Mg2+- P6) =
-
et
0’69
[(conductibilite) - P6] = - 0’70 L’homogeneite de ces correlations met bien en evidence le temps beaucoup plus lent de l’infiltration dans les blocs capacitifs que celui de la surface A la resurgence. Si le K + obeit ici A une loi de dilution cel51. vient d e la nature des sols. Aux Cugnets (No. 14) les eaux de ruissellement s’enfoncent rapidement dans des calcaires denoyes, depourvus de bonne couverture de sols. A Brof-
333
Dessus (No. 16) la couverture des flancs anticlinaux est surtout formee de sols bruns depourvus de MO soluble; les eaux sont en contact avec des calcites dolomitiques, la concentration en Mg2+ est plus forte que dans les autres piezometres. Les eaux d’infiltration, plus pauwes en Mg2+,diluent les eaux du piezometre. C’est ce qui explique le bon coefficient de correlation de signe negatif. A Martel-Dernier (No. 15), K+, Mg2+ et conductibilite sont correles positivement avec les precipitations. Pour le Mg2+ la correlation est la meilleure avec les precipitations du jour precedant (P3): r = 0’72, pour le K+, avec la moyenne des precipitations des 2eme et 3eme jours precedents (P4) r = t 0’68 et pour la conductibilite avec la moyenne des precipitations de la semaine precedente. L’alimentation de ce bloc capacitif est mixte; statistiquement on enregistre d’abord l’arrivee d’eau plus riches en Mg2+ (effet piston de surface) puis celle d u marais acide. Petit Martel (No. 17) est un cas B part. Son pH est eleve (pH 9); aucune correlation n’apparait entre K+, Mg2+ et precipitations. Cependant c’est le pH qui devient un marqueur de l’effet de dilution: r (pH-P2) = -0,75, il est en effet correle negativement avec les precipitations du jour meme; la conductibilite se correle le mieux avec les precipitations du jour precedant r = -0,66. L’infiltration des eaux de precipitation est assez rapide dans ce piezometre. Si le Mg2+ ne presente aucune correlation, bien que, comme B BrotDessus, les eaux soient en contact avec des calcaires dolomitiques, c’est qu’8 pH 9 le produit de solubilitk est toujours atteint malgre l’apport des eaux de surface et ainsi les concentrations en Mg2+restent toujours faibles.
+
COMPARAISON D U CHIMISME MOYEN DES PRINCE’AUX POINTS D’EAU
Au Tableau VI figurent les teneurs moyennes et leur estimee de l’ecarttype des eaux des quatre piezometres de la resurgence et de la perte principale. Les eaux de la perte, de la resurgence et des blocs capacitifs des Cugnets (No. 1 4 ) et de Martel-Dernier (15) sont du type carbonate-calcique; celles de Brot-Dessus (16) sont carbonatees magnesiennes, tandis que celles de Petit-Martel, avec un pH moyen de 9,2, ont une charge dissoute faible. D’apres le chimisme moyen, les eaux de la perte (No. 1 0 ) sont les plus chargees; l’exutoire No. 13 a une composition situee entre celle de la perte et celle des blocs capacitifs. On aurait tendance 8 de conclure que les eaux de la resurgence sont un melange des eaux de la perte et des blocs capacitifs. Cependant par les coefficients de correlation, avec les donnees temporelles (precipitations, debits du jour meme, du jour precedent ou du lendemain, etc.) par les comparaisons debits-precipitations-chimisme, nous avons vu que les variations de l’exutoire n’avaient aucune correlation avec celles enregistrees dans les blocs capacitifs. Nous avons essaye pourtant de decaler dans le temps chimisme des blocs capacitifs et chimisme de l’exutoire, sans succes. Au contraire, le chimisme de la perte et de la resurgence varient
w w Ip
TABLEAU VI Moyenne des valeurs chimiques et physiques (S, estimee d e I'ecart-type, N nombre de rnesures retenues; la periode d'analyse: 1974-1975) Variables
1 3 Noiraigue
S, 5 6 7 8 9 10
11 12 13 14 15 16 17 18
19 20 21 22 23 24 39
Profondeur (m) Ca2+(ppm) Mg2+(ppm) Fer total, Fe,,(ppm) Fer filtre, FeFI (ppm) Sr2+(ppm) Na+(ppm) K+(ppm) so$-(ppm) SiO, (ppm) NO5 (ppm) Cl-(ppm) Oxyghne dissous (ppm) Conductibilite ( p S ) pH Temperature ("C) Po: (ppm) Dureti totale, DuTo (ppm) Dureti temporaire, DuTe (ppm) Matieres organiques dissoutes, MO (ppm) FeFI/Fet,,t (%)
300 81,5 3,9 0,18 0,06 0,23 1,23 0,95 4,56 3,06 5,05 3,05 9,20 379,8 7,32 7,71 0,33 225,87 215,64 6,22 36,9
0 6,7 1,0 0,17 0,06 0.12 0,23 0,26 3,25 0,62 2,90 1,17 1,50 16,89 0,21 0,60 0,35 13,98 12,69 0,14 28,2
15 Martel Dernier
1 4 Cugnets
N 27 27 27 27 25 27 27 27 27 27 27 27 27 27 27 27 27 27 27 27 25
POI 34 76,7
1,s 3,32 0,09 0,18 0,58 0,58 5,52 2,79 2,35 1,05 10,25 344,5 7,31 5,65 0,30 221,40 202,8 5,2 17,4
So
N
17
14 14 14 14 12 14 14 14 14 14 14 14 14 14 14 14 14 14 14 14 12
1,6
0,4 637 0,13 0,04 0,15 0,31 1,96 0,84 1,41 0,65 1,21 21,5 0,15 0,53 0,23 19,5 7,4 0,22 16,6
16 Brot-Dessus
so
184 80,6 0,89 0,655 0,14 0,08 0,81 0,58 14,50 1,74 26,87 2,89 11,68 373,8 7,45 7,64 0,29 217,46 167,03 5,6 61,2
8,2 1,9 0,15 0,95 0,11 0,05 0,14 1,08 6,81 0,59 15J5
1,48 0,98 113,7 0,32 0,27 0,27 26,4 17,4 3,8 87,5
14 14 14 14 12 14 14 14 14 14 14 14 14 14 14 14 14 14 14 14 12
279 71,2 18,54 3,16 0,25 0,73 2,35 1,27 11,12 3,31 1,47 8,31 10,53 411,62 7,54 7,17 0,33 260,9 230,4 5,O 27,7
.
10 Emposieu
17 Petit-Martel
So
N
5,14 6,8 2,68 3,37 0,54 0,53 3,46 0,132 8,09 0,99 1,OO 10,7 1,3 22,l 0,11 0,22 0,29 17 22,8 0 51,2
12 12 12 12 12 12 12 12 12 12 12 12 12 12 12 12 12 12 12 12 12
263 16,7 5,71 2,94 0,lO 0,lO 2,14 2,12 1,28 0,60 0,29 5,53 3,655
111,6 9,2 8,2 0,20 64,32 50,45 5,07
so
N
so
N
3,75 10,17 0,638 3,53 0,12 0,05 0,79 0,44 1,lO 0,47 0,23 1,58 0,54 12,2 5,5 0,25 0,24 17,4 5,O 2,16
14 0 1 4 90,04 14 4,43 14 0,83 12 0,40 14 0,43 14 3,16 14 1,61 14 7,53 14 3,71 14 3,64 14 6,62 14 8,65 1 4 432,23 14 7,51 14 8,88 14 0,93 1 4 256,67 1 4 243,92 1 4 10,95 12 56,7
0 11,46 2,06 0,37 0,13 0,24
26 26 26 26 23 26 26 26 26 26 26 26 26 26 26
1,81 0,48 1,90 1.40 1,92 8,72 2,38 5,83 0,33 4,02 0,652 3,40 3,60 2,81
26 26 26 26 26
335
remarquablement en phase, comme du reste leurs debits [voir les Conclusions, l’objet (Z)] .
Interprktation d e la saturation e n Pcoz Pour les equilibres: C02
+ H 2 0 =+H’ + HC05
et
C02
+ H 2 0 + CO-;” + 2H’
B partir de I’energie standard d e reaction de Gibbs et des comparaisons dans les eaux, on calcule:
Q =
[H’]
[CO-;”]
1
[H,Ol ou Q est le coefficient reactionnel. Si log K - log Q est inferieur A 0, les eaux sont sous-saturees en COz, si cette difference est positive, les eaux sont sursaturees par rapport A Pco2 . La Pcoz occupe une position cardinale dans les equilibres des eaux karstiques. C’est une grandeur de premiere importance pour la comparaison entre les diverses eaux. Nous n’avons trouve aucune correlation entre Pco, des piezomgtres et de la resurgence. La correlation est par contre excellente entre la variation de Pco2 de la perte et celle de l’exutoire (Fig. 10, r = 0’8). C’est une preuve supplementaire de 1’6troite dependance d u chimisme des eaux de la resurgence et de celles de la perte. Cette preuve est encore soulignee par l’excellente correlation entre Mg2+ de la perte et de la resurgence ( r = 0’7, voir Fig. 11). [PCO,
CONCLUSIONS
Dans un bassin karstique jurassien (altitude 700 A 1100 m) avec une partie du ruissellement concentree sur une perte totale, equipe de quatre piezombtres et deux limnigraphes ( A la resurgence et A la perte), sur la base d’un releve journalier des debits, precipitations et niveaux dans les piezometres et par un echantillonnage mensuel des eaux sur 24 A 27 mois, il est possible, par un traitement correct des donnees, d’apporter les precisions suivantes. Dans le bassin 6tudie et dont l’originalite consiste en la presence d’un marais tourbeux important: (1) L’effet de dilution par l’augmentation du debit est le mieux illustre par le Mg”. (2) L’effet-piston est localise aux eaux de surface, au niveau des sols organiques et des tourbes. (3) L’heterogeneite des eaux de surface, leur groupement par rapport aux eaux de la perte, de la resurgence et des piezom&tres,la parente des eaux de la perte et de la resurgence est bien mise en evidence par l’analyse factorielle des correspondances selon Benzecri et al. (1973).
336
r = 0,8 I
I
I A
I B
I C
- log f
I D
I
l
l
1
1
E
F
G
H
I
J
- log
Q
[Con1 :tog K
1
1
K
1
L
1
1
M
N
1
I
I
i
I
I
I
I
I
I
I
I
O
P
Q
R
1
2
3
4
5
6
7
8
R
1
2
3
4
5
6
7
8
Point 13
A
B
C
D
E
F
G
H
I
J
K
L
M
N
O
P
Q
Fig. 10. Sous-saturation en pression partielle de C 0 2 , comparaison entre la perte (10) et la r6surgence (13) mois par mois (A a R ) et semaine par semaine (1 8). Les variations de sous-saturations sont parfaitement identiques entre les points (pour le calcul de la soussaturation, explication dans le texte). Fig. 10. Undersaturation of the partial C 0 2 pressure, comparison between sink (10)and karst (13) water - sampled month to month ( A - R ) and week t o week (1-8). The variation of the carbon dioxide undersaturation is entirely identical between the different data points (for calculation of the carbon dioxide undersaturation, see explanation in the text).
(4)L’individualite des eaux de chaque piezomhtre, dejja sensible par la cornparaison du chimisme moyen, est clairement dernontrbe par 1’AFC. (5) Par le calcul des coefficients de correlation sur chaque groupe d’eau,
337
en m g l l 8- [Mg2+]
1-
6-
I P Point 10
5-
4-
3-
2-
r = 0,7
1-
1
1
1
1
1
1
1
1
1
1
1
1
1
1
1
1
1
A
E
C
D
E
F
G
H
I
J
K
L
M
N
O
P
1 Q
F
G
H
I
J
K
L
M
N
O
P
Q
1 R
I
I
I
I
I
I
I
I
1
2
3
4
5
6
7
8
2
3
4
5
6
7
8
[Mg2+] en mgl I
A
B
C
D
E
1
Fig. 11. Comparaison des teneurs en Mg2+ de la rksurgence (1 3 ) et de la perte (1 0 ) mois par mois et semaine par semaine. Fig. 11. Comparison of the Mg2+ concentration of karst (13) and sink (10) water - sampled month to month and week to week.
l’etroite relation eaux de la perte eaux de la resurgence est quantifiee, l’independance des eaux des piezometres, donc des blocs capacitifs, est confirmee les relations precipitations, debits, chimisme sont etablies precisant le dephasage de la “vague” hydraulique et des divers dephasages de la vague chimique. En termes globaux on doit conclure que l’heterogeneite de la nappe se situe ti trois niveaux: heterogeneite de l’alimentation : sols calcaires, lapiez, sols bmns calciques, sols organiques, tourbes, marais, marais cultive; hetero-
338
geneite des permeabilites de la zone denoyee et heterogeneite de la nappe perenne: blocs capacitifs. Dans le bassin de la Noiraigue, les excellentes correlations perteresurgence, les mauvaises correlations blocs capacitifs entre eux-memes et la resurgence indiquent que les forages d’exploration n’ont pas rencontre le reseau de grands chenaux et que par comparaison avec d’autres karsts plus meridionaux, le karst etudie doit etre en partie colmate.
REMERCIEMENTS
Les remerciements des auteurs vont 8 : Moret, statisticienne-conseil, Centre de Calcul de 1’Universite de Neuchatel, qui s’est chargee de 1’AFC. -C. Wittwer, Institut de Geologie de 1’Universite de Neuchatel, grace 8 laquelle les niveaux piezometriques ont ete dessines minutieusement. - A. Burger, Professeur, Directeur du Centre d’Hydrogeologie de 1’Universite de Neuchatel, qui a bien voulu relire ce manuscrit et nous aider de ses conseils. - B. Fritz, C.N.R.S. Strasbourg, qui, 8 l’aide des methodes d’Hegelson, nous a programme les indices de saturations. - L. Kiraly, Directeur de recherche au Centre d’Hydrologie de 1’Universite de Neuchatel. - Groupe de Recherche en Methodes Quantitatives de 1’Universite de Neuchatel (Suisse): Prof. Strohmeyer. - Fonds National Suisse de la Recherche Scientifique, requetes No. 2.542.76, 2.768.77 et 2.045.78. - J.
BIBLIOGRAPHIE Benzecri, J.P., 1970. Distance distributionelle et mkthode du chideux en analyse factorielle des correspondances. Lab. Stat. Math., Univ. Paris VI, Paris, 3e kd. Benzkcri, J.P. et al., 1973. L’analyse des donnkes, Tome 2. L’analyse des correspondances. Dunod, Paris, 619 pp. Bobee, B., Cluis, D., Goulet, M., Lachance, M., Potvin, L. et Tessier, A., 1977. Evaluation du rkseau de qualit6 des eaux. Analyses et interpr6tation des donnkes de la pkriode 1967-75.I.N.R.S.-Eau, Rapp. Min. Rich. Nat., Qukbec, Que., Rapp. Sci., No. 78. Boulaine, J., 1971. L’agrologie. Coll. “Que sais-je?” Paris, No. 1412, 1 2 3 pp. Bouyer, J., en prhp. Diplacement du fer dans un karst en partie couvert de tourbes. Universite de Neuchitel, Neuchitel. Burger, A., 1959. Hydrogeologie du bassin d e 1’Areuse. Th&se, Universitk de Neuchitel, Neuchitel, 304 pp. DCsor, E., 1864. Expbriences sur la dur6e du parcours souterrain des eaux d e la Noiraigue. Bull. S O C . Neuchitel. Sci. Nat., 7: 37-59. Drogue, C., 1971. De l’eau dans le calcaire. Sci., Prog. Dkcouvertes, No. 3433, pp. 39-46. Jaccard, A,, 1883. Note sur le changement du r6gime des sources dans le Jura neuchatelois. Bull Soc. Neuchitel., No. 13.
339 Kiraly, L., 1969. Anisotropie et hktkrogknkitk de la permkabilitk dans les calcaires fissurks. Eclogae Geol. Helv., 62(2): 613-619. Kiraly, L., 1973. Notice explicative de la carte hydrogkologique du canton de Neuchitel. Suppl. Bull. SOC.Neuchitel. Sci. Nat., 96, 1 5 pp.(une carte). Kiraly, L., 1975. Rapport sur l’ktat actuel des connaissances dans le domaine des caract&es physiques des roches karstiques. Dans: A. Burger et L. Dubertret (Editors). A.I.H., Paris, pp. 53-67. Kiraly, L., 1978. La notion d’unitk hydrogkologique. Bull. Cent. Hydrogkol. Neuchitel, NO. 2, pp. 83-216. Kiraly, L. et Morel, G., 1976. Remarques sur l’hydrogramme des sources karstiques, simulk par modhles mathkmatiques. Univ. Neuchitel, Inst. Gkol., Bull. Cent. Hydrogeol., No. 1, pp. 37-60. Kiraly, L. et Muller, I., 1979. Hdtkrogknkitk de la permkabilitk et de I’alimentation dans le karst: effet sur la variation du chimisme des sources karstiques. Univ. Neuchitel, Inst. Gkol., Bull. Cent. Hydrogkol., No. 3, pp. 237-285. Kubler, B., 1972. Le sel, agresseur mkconnu de notre environnement. Bull. SOC.Neuchitel. Sci. Nat., 95: 133-163. Kubler, B., Pochon, M. et Simeoni, J.-P., 1978. Les troubles des eaux karstiques: un exempfe d’implication de l’hydrog6ologie et de la minkralogie, pkdologie, skdimentologie et gkochimie. Symp. Int. sur Implications de I’Hydrogkologie dans les autres Sciences de la Terre, Montpellier, 11-16 Sept. 1978. Lachance, M., Bobee, B. et Gouin, D., 1979. Characterization of the water quality in the St. Lawrence River: determination of homogeneous zones by correspondence analysis. Water Resour. Res., 15(6): 1451-1462. Mangin, A., 1975. Contribution i l’ktude hydrodynamique des aquif6res karstiques. ThGse, Lab. Souterr., Cent. Natl. Rech. Sci. (C.N.R.S.), Moulis, 124 pp. Miserez, J.J., 1973. Geochimie des eaux du karst jurassien. (Contribution physicochimique il’ktude des alterations). Thhse, Universitk de Neuchitel, Neuchitel, 313 pp. Morel, G., 1976. Etude hydrogkologique du bassin de la source de la Noiraigue (Jura). Univ. Neuchitel, Cent. Hydrogkol. Persoz, F. et Kubler, B., 197 3. Compositions minkralogiques et chimiques des roches. Bull. SOC.Neuchitel. Sci. Nat., Suppl., 96: 11, pl. 1. Pochon, M., 1978. Origine et evolution des sols du Haut-Jura suisse. Thhse, Universitk de Neuchitel, Neuchitel, 190 pp. (aussi Mem. SOC. Helv. Sci. Nat., 90, 190 pp.). Pochon, M. et Simeoni, J.P., 1976. Comportement hydrodynamique, nature et r6le traceur des troubles argileux dans deux sources karstiques: (Jura tabulaire suisse). Coll. sur I’Hydrologie en Pays Calcaire. Ann. Sci. Univ. Besanpn. Geol. Fasc. 25, 3hme skrie, pp. 321-339. Schardt, H., 1904. Origine de la source de 1’Areuse (La Doux); Bull. SOC.Neuchitel. Sci. Nat. Mkm. Gkol., 5e Fasc. 32: 118-139. Schoeller, H., 1967. Hydrodynamique dans le karst. Chron. Hydrogkol., 10: 7-21. Tripet, J.-P., 1972. Etude hydrogkologique du bassin de 1’Areuse (Jura neuchitelois). Thhse, Universitk de Neuchitel, Neuchitel, 1 8 3 pp.
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341
EFFECT OF LEACHATE SOLUTIONS FROM FLY AND BOTTOM ASH ON GROUNDWATER QUALITY DEBORAH A. KOPSICK and ERNEST E. ANGINO*
Department of Geology, University o f Kansas, Lawrence, KS 66044 (U.S.A.) (Accepted for publication February 27, 1981)
ABSTRACT Kopsick, D.A. and Angino, E.E., 1981. Effect of leachate solutions from fly and bottom ash on groundwater quality. In: W. Back and R. Lhtolle (Guest-Editors), Symposium on Geochemistry of Groundwater - 26th International Geological Congress. J. Hydrol., 54: 341-356. Leaching experiments on fly and bottom ash for Ca, Mg, Na, K, Fe, Mn, Zn, Cu and Pb indicated a potential for contamination of ground- and surface-water supplies. Due to the variability in chemical composition of coals, it is difficult to make generalizations concerning the chemistry of leachate solutions from the ashes of the coals. A decrease in concentration with time of leaching was observed for all elements, except for Ca which was released at a constant rate. Fly ash from a Missouri coal generated a leachate enriched in Pb, Zn, Cu, Fe, Mn and Cd, reflective of the high Pb-Zn mineralization present in the surrounding area, With a pH of 3.0 this ash has the greatest potential for groundwater contamination. Conversely, leachates from Wyoming fly and bottom ashes exhibited low trace-metal concentrations. These same solutions were high in K, Na, Ca and Mg, and also showed strong pozzolanic behvior, which will reduce the leachability of these ashes. In most instances, fly and bottom ash from Kentucky and Illinois coals yielded leachates intermediate in elemental composition to leachates of Missouri and Wyoming coal ashes. Leaching experiments indicate that it is not valid to predict the chemistry of leachates from fly and bottom ash based solely on the chemical composition of the ash. From the limited number of parameters and sites examined in this study, it is clear that many of the problems related to leachates from fly and bottom ash and gob piles are site specific as well as specific to the source of coal burned. These results are, nevertheless, indicative of problems likely to be encountered in working with these materials.
INTRODUCTION
Estimates for usage of coal in the U.S.A. for the year 1980 range up to 8 * 10' metric tons (t) (Surprenant et al., 1975). Waste products from coal combustion include fly ash and bottom ash. Projected production of these ashes in 1980 is estimated at 75*106t (Faber et al., 1976). Disposal of the ash material is one of the major environmental problems associated with the use of coal-fired power plants.
* Person to be contacted for reprints.
342
Based on current trends in energy production, it is apparent that there will be a sharp increase in the amount of coal ash produced. Many power plants in the U S A . that are able to convert from oil and gas t o coal are being ordered to do so by the US. Department of Energy. Emission standards have become more stringent for sulfur dioxide (SO,) emission and many power plants have switched from high-sulfur eastern U.S.A. coals to lowsulfur western U.S.A. coals. Western U.S.A. coals generally produce more ash than eastern U.S.A. coals. Due to more stringent particulate emission control regulations, more efficient ash collection devices have been developed and utilized; consequently, many new coal-fired plants now have the capacity of removing up to 99% of the fly ash produced. Therefore, even if the amount of coal burned remained constant, the amount of ash produced and collected will increase. By controlling the two major air-pollution problems, production of SO, and air-borne particulates, the amount of solid waste produced from coal-fired power plants has increased. Many power plants are located adjacent t o rivers, which provide a source of cooling water. Floodplain soils and sediments are permeable and may be subject to flooding and seepage from surface impoundments. Therefore, it is important that ash pond disposal sites constructed in floodplain sediments be designed so that the waste is isolated from the environment. The obvious way to limit the quantity of leachate produced is to minimize the contact of the ash with water. However, the most common method of transferring coal ash from the collection areas t o the final disposal site is by sluicing with water. Dry disposal has the advantage of limiting the amount of water that comes in contact with the ash; however, there are problems with wind erosion and stabilization of the ash pile. Fly ash and bottom ash has been declared a “special waste” by the U S . Environmental Protection Agency, in concordance with the Federal Resource Conservation and Recovery Act (R.C.R.A.). This designation requires that coal ash disposal ponds be sealed t o prevent downward percolation of leachate and that groundwater monitoring systems be installed at the site. Many disposal ponds constructed prior to 1976 are unlined and may allow the infiltration of leachate into the groundwater aquifer. While newer ponds may be lined, the life of the lining before it is breached is unknown. Therefore, information is needed on the chemical composition of ash leachates t o evaluate the effect of this leachate on groundwater quality. The physical and chemical characteristics of these ashes, combined with the operational parameters of the power plant and the disposal environments in which the ashes are placed, control the leaching susceptibility of these wastes and determine the potential for contamination t o groundwater aquifers. Little information is known on the effect of coal ash leachates on groundwater quality (Theis, 1975; Theis et al., 1978, 1979). This paper discusses the concentrations of selected elements that can be expected t o be released under ash disposal conditions, as simulated by laboratory experiments.
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Description of ash In order t o understand the leaching behaviour of fly ash and bottom ash, we need to understand the mechanics of formation and the morphology of the ash. Bottom ash is the coarse-grained material (greater than 100pm in diameter) that is collected at the bottom of the boiler and is then wet-sluiced to ponds. Fly ash consists of finer sized particles, ranging from 0.5 to 200 pm in diameter, but is generally in the 0.5-100-pm range. The fly ash is entrained in the gas stream and carried up the stack following combustion. The majority of this ash is recovered by collection devices. Coal ash is an aluminosilicate glass, consisting of the oxides of Si, Al, Fe and Ca, with minor amounts of Mg, Ti, Na, K and S, and various trace elements. The trace-element concentrations associated with the ash may be either adsorbed onto the surface of the particle or incorporated into the matrix (Kaakinen et al., 1975; Cambell et al., 1978). Collin (1974) has proposed a model whereby acidic, neutral and basic layers are formed on the surface of the fly ash particle as it is caqied through the gas stream. Ash particles are formed through the thermal decomposition or dehydration of inorganic minerals associated with the coal. Calcium carbonate and clay are the most abundant mineral impurities, with lesser amounts of sulfides, chlorides and oxides also present. The shape of the ash particle is dependent on many factors, two of which are the amount of time and temperature t o which the coal is exposed in the combustion chamber (Fisher et al., 1978). The spherical shape, most commonly associated with fly ash particles, shows that complete melting of silicates occurs at high temperature. These spheres may be solid, hollow (cenospheres) or encapsulating spheres (plerospheres). SCOPE OF STUDY
An evaluation was made of: (1)the relation of the whole ash composition (unleached sample) to composition of leachate generated from the same ash for Ca, Mg, Na, K, Fe, Mn, Pb, Cu and Zn; (2) the chemical composition of the ash leachate on a temporal basis and the effect of repeated leachings; (3) the physical changes that occur within the sample during leaching; and (4)the contamination potential of the ash leachates in relation t o groundwater quality. Material sump led Nine different ashes (six fly ashes and three bottom ashes) were evaluated. These samples represent the coal ash from four U.S.A. coal-producing provinces and are a cross-section of the more widely utilized coals in the U S A . The location and origin of the coals from which these ashes were derived are shown in Fig. 1. The method of combustion of the power plant from which
344
345
TABLE I Identification of ash samples Sample identification
Origin of coal
Type of combustion
Method of collection
1 WYO-FA*'
Powder River Basin, Wyoming
pulverized (unwashed)
dry (ESP)*3
2 CCWY-FA
Carbon County, Wyoming (Hanna Field)
pulverized
dry (ESP)
ccWY - B A * ~ 3 OK-FA
wet (pond) northeast Oklahoma
pulverized (washed)
OK-BA
dry (ESP) wet (pond)
4 NEMO-FA
northeast Missouri ( Wheeler-Beevier bed)
cyclone washed)
dry (ESP)
5 SI-FA
blend of southern Illinois coals
cyclone (unwashed)
dry (ESP)
(4
SI-BA 6 EKY-FA
*1
wet (pond) eastern Kentucky (Kanawha River Valley)
pulverized (unwashed)
dry (ESP)
FA = fly ash; *' BA = bottom ash; * 3 ESP = electrostatic precipitator.
the ashes were collected is shown in Table I. The fly ash samples were collected dry from the electrostatic precipitator hoppers, while the bottom ash samples were collected wet from the disposal pond (Table I). Since bottom ash is the material that drops to the bottom of the boiler, it is very difficult, if not impossible at most plants, to obtain samples of bottom ash that have not been in contact with sluicing water.
METHODS OF ANALYSIS
Chemical methods Whole ash analyses were performed using flameless atomic absorption spectrophotometric techniques. The lithium metaborate-nitric acid fusion techniques was used in the preparation of the sample for analysis of Ca, Mg, Na and K. Hydrofluoric-nitric acid dissolution was employed in the preparation of the samples for measurement of Fe, Mn, Cu, Pb, Zn and Cd. Leachates generated from these ashes were then analyzed for Ca, Mg, Fe, Mn, Cu, Pb and Zn, using standard flame atomic absorption methods. The concentrations of Na and K in the leachates were analyzed by flame photometric methods.
346
Leaching methods Leachates were derived from the coal ash samples, using a column apparatus. Glass columns, 1.0 m in length and 5 cm internal diameter, were packed to a height of 46 cm with ash. Distilled, deionized water was passed through the column from the bottom. This direction of flow was chosen to prevent channelization and to insure complete saturation. Ponding of the leaching solution in the tube was prevented through the maintenance of a hydrostatic head above the entry point of the tube. Leachate was collected as it emerged from the top of the column, with seven samples collected within a 2-hr. period. This sequence was repeated one week and two weeks after the original samples were taken. Leachings for each ash sample were done in triplicate giving a total of nine data sets per sample. This leaching procedure is described in more detail in Kopsick (1980). Between leachings, the sample in the column was allowed to remain saturated. This method of leaching was chosen in order to simulate natural leaching conditions as would occur at an ash disposal site. Data from the three leachings, taken at one-week intervals, helped in the evaluation of any flushing that would occur at the onset of rainfall and also t o determine the effect of prolonged contact of the rain water with the ash between rainfall events. Column tests also allow the observation of changes that occur in the physical characteristics of the wastes during leaching.
RESULTS
Observed leaching trends Three distinct leaching patterns were observed for the samples examined. The most prevalent trend was characterized by a large initial release of elements in the sample, and a leveling off of concentrations at later time periods. Subsequent leachings also exhibited this trend; however, the initial release was lower for each successive run. The majority of trace metals exhibited this response to leaching. This pattern represents the immediate leaching of soluble constituents on the surface of the ash particles, as well as extraneous mineral matter in the ash. Theis et al. (1978) noted a substantial increase in the concentration of Fe, Ni, Pb, Cu, Zn and As in local groundwater when fly ash was placed in a disposal pond at a power plant in Indiana following a shutdown of the plant. This field observation is analagous t o the initial release curve we observed in the laboratory. A constant release pattern was observed for Ca in the Oklahoma bottom ash and fly ash (OK-BA, OK-FA) and Wyoming bottom ash and fly ash (CCWY-BA, CCWY-FA). Subsequent leaching of these samples did not significantly lower the concentration of the Ca in the leachate. Soluble Ca-sulfates and -oxides present in the ash represent a constant source of Ca for release.
347
1501
A
I N I T I A L RELEASE CCWY-BA
\,
25
_----. _
.
@ CONSTANT
I soot.,
I
,
,
,
.
)
,
25-
(mgil) Mg 20: 15
@DELAYED
/PI--
RELEASE NEMO - FA
RELEASE EKY-FA
-.\
10
I 10
30
50
70
90
110
T I M E IN MINUTES
Fig. 2. Leaching trends observed. The 1 , 2, 3 designation refers to consecutive leaching tests that were run on the same sample at one-week intervals.
The third leaching trend observed involved a delayed release curve in which a short (1--5-min.) period of time elapsed before the maximum concentration in the leachate was observed. The eastern Kentucky fly ash (EKY-FA) sample exhibited this trend. This leaching behavior appeared to us to be related t o the morphology of this ash. Large-diameter hollow cenopheres dominate this sample. These particles have a smaller total surface area available for reaction with the leaching solution. Reaction times are therefore delayed or slowed down. Fig. 2 shows these three leaching trends. Leachate composition us. whole ash composition Leaching experiments show that the elemental composition of the leachate may not proportionally reflect the elemental composition of the whole ash sample (Table 11). The statistical means of the maximum concentrations for elements in the leachate d o not mirror the progression of concentrations for the same elements in the whole ash sample, as shown in Fig. 3. This
TABLE I1 Whole ash analyses (ppm) of fly ash and bottom ash for selected major and minor elements Sample
Ca
Mg
Na
K
Fe
cu
Mn
Pb
Zn
Cd
CCWY-FA CCWY-BA SI-FA SI-BA OK-FA OK-BA WYO-FA EKY-FA NEMO-FA
87,400 105,900 14,000 40,000 56,000 68,000 200,200 5,000 14,800
17,400 12,000 2,700 48,000 5,500 5,400 25,500 23,500 4,500
2,900 1,600 1,300 2,000 3,700 1,900 12,900 2,000 3,500
15,600 13,000 10,200 15,800 19,000 10,900 3,000 19,400 19,200
41,900 38,200 78,000 206,500 187,600 186,500 37,900 26,000 225,800
139.2 55.3 50.9 34.3 220.4 133.3 243.5 187.7 244.7
382.4 404.8 88.9 654.1 988.0 1,320 355.5 24.5 171.3
104.0 42.1 446.7 43.2 757.9 242.9 53.8 54.0 720.5
309.7 67.2 1,701 209.9 5,104 1.796 262.1 116.4 2,226
4.22 1.60 8.14 3.23 26.7 5.29 6.34 8.10 15.3
349
I WHOLE
ASH
M A X LEACWATE
mean range
A
mean
&j
range
I '
105
'
10'
103
(vim)
10'
10'
I Na
Mg
K
10"
Zn
Pb
Fig. 3. Range and mean of selected elements in whole ash samples and mean maximum leachate values generated from these ashes. Cross-hatching represents an area where the range for the leachate and whole ash composition overlaps. Selected elements are arranged in order of increasing atomic number.
suggests, that for some elements, a correlation of leachate quality t o whole ash composition cannot be made. The rate at which these elements will leach from the ash sample is dependent on the form in which the element is present and the location of the element within the ash matrix or adsorbed onto the ash particle surface. The ash spheres are chemically stable in the environment and are resistant t o weathering due to the aluminosilicate matrix. Any element present in this matrix will not be readily avaiable for leaching. Elements adsorbed onto the surface of the ash spheres will be more readily leached. Uncombusted mineral matter may account for presence of high concentrations of certain elements in the whole ash analyses. However, leachates generated from these ashes may not reflect the high concentrations because the extraneous material associated with the ash is not in a form that is susceptible t o leaching within the time span of these experiments.
350
t
_1
\
-,om \i
.
\
Fig. 4. Whole ash composition (solid line) vs. mean maximum leachate composition (dashed line) for selected elements in four coal ashes.
Fig. 4 shows the elemental composition of four coal ashes, along with the mean maximum concentration leached from selected ashes. For the four samples, Ca, Mg, Na and K were the elements most readily released by leaching with water. These concentrations will rapidly decrease as the time since generation of the first leachate increases. Owing t o the volatility of Mn, Pb, Zn and Cu, it is assumed that these elements are released in the gaseous state during combustion and condense onto the fly ash spheres as cooling occurs (Davison e t al., 1974; Kaakinen e t al., 1975; Linton et al., 1976; Smith et al., 1979). Therefore, these elements make initial contact with the leaching solution. The chemical complexing of these elements and the method of adsorption and attenuation are important factors to consider when estimating leachate composition from ash composition. From the data, it is concluded that Fe is located primarily in the ash
351
matrix. Fig. 4 shows that although high concentrations of Fe are recorded for the whole ash sample, only minor amounts of this Fe are found in the leachate. Uncombusted pyritic material is another source of Fe in the whole ash sample, particularly bottom ashes which contain the dense material that drops to the bottom of the boiler during combustion.
Northern Great Plains Coal Province One sample of Wyoming fly ash (WYO-FA) from the Northern Great Plains Coal Province was subjected t o leaching. Mn, Pb, Zn and Cu were detected in very minor amounts (G 0.15 mg/l). The initial leachate solution had an Fe concentration of 6.2mg/l, with no Fe being detected in later samples. Of the elements analyzed, the major constituents in the leachate were Ca, Na and K.
Rocky Mountain Coal Province Leachate samples from Hanna Basin coal ash (CCWY-BA, CCWY-FA) Carbon County, Wyoming were analyzed. Fe, Mn, Pb, Zn and Cu were not detected in the bottom ash leachate. Of the major elements, K had the highest concentrations (11-367 mg/l) followed by Ca (36-142 mg/l), Na (1.1-172 mg/l) and minor amounts of Mg (0.07-1.64mg/l). No Pb was detected in the fly ash leachate. Minor amounts of Fe (G2.9 mg/l), Mn (GO.l5mg/l), Zn (G0.34mg/l) and Cu (<0.65mg/l) were detected. Na concentrations were enriched seven-fold in the fly ash relative t o the bottom ash. Ca was released consistently in the 455-535-mg/l range.
Interior Coal Province Leachates from fly and bottom ash samples from northeast Missouri (NEMO-FA), Oklahoma (OK-FA, OK-BA) and Illinois (SI-FA, SI-BA) were found to vary considerably in chemical composition. The NEMO-FA produced leachates with pH values of 2.8-3.7. Fe, Mn, Pb, Zn and Cu were found in elevated concentrations with mean maximum values of 2381, 39.7, 5.4, 504 and 444 mg/l, respectively. These values were the highest recorded for these elements for all the ash samples analyzed. The northeast Oklahoma fly ash (OK-FA) sample resembled the NEMOFA sample in Mg, Na, K, Fe, Cu and Pb concentrations and had significantly higher concentrations of Ca, Mn and Zn. The leachates from OK-FA, however, exhibited only traces of Mn (G0.48mg/l), Pb (<0.67mg/l), Zn (G2.09 mg/l) and Cu (G0.17mg/l), and a maximum of 18.1mg/l Fe. Ca concentrations of the NEMO-FA and OK-FA were similar, but the OK-FA sample was deficient in Mg, Na and K relative to the NEMO-FA. The pH of the OK-FA leachates were 10.5-11.8. If a prediction concerning leachate quality based on whole ash analysis were to be made, these two samples
352
should have similar leachate compositions. The low pH of the NEMO-FA samples are most likely due to fine-grain disseminated pyrite associated with the ash. The oxidation of this pyrite upon expossure t o oxygen and water will produce sulfuric acid, and this acidic soluion will mobilize the Mn, Pb, Zn and Cu, allowing these elements t o remain in solution. The OK-BA sample yielded a leachate with a pH of 8.8-9.6, with traces of Fe (<1.93mg/l), Mn (<0.034mg/l), Pb (<0.17mg/l) and Zn (GO.11 mg/l), and not detectable Cu. Ca concentrations were comparable with those of OK-FA, but Mg concentrations were lower. Na and K were slightly more concentrated in the bottom ash than the fly ash. The SI-BA sample produced an acidic leachate, with a pH of 3.7-4.2. A mean maximum value of 157mg/l Fe was measured. The concentrations of Mn, Pb, Zn and Cu, as well as Ca, Mg, Na and K were low. The fly ash from this coal (SI-FA) showed lower Fe concentrations (<2.4mg/l), Mn (<2.87 mg/l) and Cu (<0.10 mg/l) than the bottom as leachate. The concentration of Zn was elevated in the fly ash, with a maximum concentration of 21.4 mg/l being recorded. The ash samples from Oklahoma, Missouri and Illinois contain the highest concentration of Pb, Zn and Fe of the nine samples analyzed. This likely represents secondary enrichment of the coals as a result of a regional Pb-Zn mineralization that occurs in this province. Eastern Coal Province One coal ash from the Kanawha River Valley of eastern Kentucky (EKYFA) was evaluated. Trace-metal concentrations were low in this leachate and also low in the whole ash analysis. Ca, Mg, Na and K concentrations were deficient in the leachate in comparison to the other ash leachates.
p H of ash leachates The pH of the ash leachate is one parameter that controls the migration potential of the elements leached from the ash. Fig. 5 shows the variation in pH that exists between the various leachates collected. Due to the high concentrations of Ca and Mg in most ashes, the majority of leachates are alkaline. Notable exceptions are the slightly acidic leachate of the EKY-FA sample and the acidic leachate of the NEMO-FA and SI-BA samples. These acidic leachates will allow metals leached from the ash t o remain in solution, increasing the mobility of these elements. This mobility contributes greatly to the potential for contamination of groundwater aquifers by such solutions. The SI-FA leachate had an initial pH of 5.5, with a rapid increase t o a pH of 11. The coal from which this ash was formed had not been washed prior t o burning. The initial acidic pH is due to the rapid oxidation of sulfides (e.g., pyrite) present in the ash as a result of incomplete combustion of the
353
10
20
30 40 50 60
7 0 80
90 100 110 120
Time, in minutes,since generation of first leachate
Fig. 5. pH of ash leachates during the initial 2-hr. leaching period. Values represent means of three replications.
coal impurities. The bottom ash from this southern Illinois coal also exhibits an acidic leachate; however, the pH remains acidic throughout the leaching period. This can be related t o the concentration of the mineral matter associated with the coarse-grained bottom ash and slag. Physical alteration of ash during leaching The physical structure of some of the ash samples was noted to change during the column leaching studies. The WYO-FA sample exhibited a strong pozzolanic reaction. A pozzolan is aluminosilicate material that in itself is not a cement, but when put in the presence with moisture will react with calcium hydroxide to form a cement-like material. Expansion of the ash material was also noted and enough pressure was developed t o crack the glass leaching tubes. This pozzolanic reaction produced an impermeable material in which leaching was possible only along channels and expansion features. This greatly reduced the amount of surface area that is available for leaching and prevented the complete saturation of the sample. The EKY-FA sample exhibited a density separation during leaching, with
354
the low-density cenospheres collecting on the top. This behavior has also been noted at disposal sites (Pedlow, 1974) where these “floaters” do not settle in the disposal pond. This can lead t o contamination of surface water during runoff following heavy rainfall. Channelization developed in tests of the EKY-FA, OK-FA and NEMO-FA samples. This behavior during leaching will reduce the amount of ash particles actively leached, although the ash may remain fully saturated. Gas bubbles of unknown origin were expelled along channels in the EKY-FA sample. The physical reaction of the ash to aqueous leaching is an important factor when considering the potential contamination of groundwater by these leachates. Concentrations of elements in leachates will be determined by the amount of ash exposed t o leaching solutions. Observations of this type are useful in determining the mode of disposal most suited for the ash being produced.For example, the WYO-FA sample would be better suited t o a dry disposal system, with an occasional wetting, rather than the pond system which is most frequently used.
CONCLUSIONS
Before discussing the conclusions of our study, the following comments need consideration. We studied only a few selected samples from widely scattered localities. The samples examined cannot, in any way, be considered representative of either a region or a single mining district. Considerable variation undoubtedly exists from mine to mine and region to region. The results, however, are interesting and instructive in what they show of the leaching phenomenon on fly and bottom ash piles. Of the nine coal ash samples subjected to leaching, one bottom ash (SI-BA) and one fly ash (NEMO-FA) were found t o produce a leachate that would be capable of degrading groundwater quality. Both these ashes produced acidic leachates, which would enable metals t o remain in solution and migrate from the disposal site. Three prominent leaching trends were observed for the samples, these being an initial flush curve, a constant release curve and a delayed flush curve. The most commonly noted curve for metal release was the initial flush curve, in which the highest concentrations released are in the initial leachage samples, with a rapidly decreasing concentration with increasing time. Therefore, the input of large amounts of fresh ash will produce a leachate with elevated concentrations of metals, which may overpower the exchange sites of the soils and sediments, and allow contamination to occur. Inactive ponds, providing reducing conditions do not exist, should produce a leachate with low concentrations of metals, which would pose little threat t o groundwater quality. The elemental composition of the leachate will not proportionally reflect
355
the elemental composition of the whole ash sample. Ca, Mg, Na and K are released readily, while Fe, Mn, Pb, Zn and Cu are detected in proportionally lower concentrations, with the notable exceptions of SI-BA and NEMO-FA. From the leaching data, it is concluded that Fe occurs primarily in the matrix or as extraneous sulfide material that is not susceptible t o leaching. The high concentrations of Fe in the whole ash samples are not reflected in the chemistry of the leachate. Therefore, whole ash analysis does not provide sufficient information for prediction of the chemical quality of the leachate. A number of alterations in morphology were noted during the leaching process. Pozzolanization, channelization and incomplete saturation will reduce the amount of ash that is exposed t o leaching solutions. For pozzolanic ashes, such as the WYO-FA sample, dry disposal with an occasional wetting would be the preferred method of disposal, as it would reduce the quantity of leachate produced. Except for the NEMO-FA and EKY-FA samples, the fly ash samples generated leachates with a pH range of 9.1-11.8. Bottom ash leachates were more acidic than their fly ash counterparts, due to the concentration of sulfide impurities in the bottom ash. The ash samples from Wyoming coal (WYO-FA, CCWY-FA and CCWY-BA) produce a leachate that is least likely t o degrade groundwater supplies. Samples from the Interior Coal Province (SI-BA, SI-FA, NEMO-FA, OK-BA and OK-FA) present the greatest potential for contamination to groundwater, particularly NEMO-FA. The EKY-FA sample from the Eastern Coal Province is intermediate between these two groups.
ACKNOWLEDGMENTS
This research was supported by a grant through the Kansas Water Resources Research Institute and the Office of Water Research and Technology, U S . Department of the Interior, Washington, D.C. The authors wish t o thank the Geochemistry Section of the Kansas Geological Survey, Lawrence, Kansas for the analyses of the whole ash. The assistance of A. Paul Kinkella in preparation of the graphics is greatly appreciated.
REFERENCES Cambell, J., Laul, J., Neilson, K. and Smith, R., 1978. Separation and chemical characterization of finely-sized fly ash particles. Anal. Chem., 50(8): 1032-1040. Collin, P.J., 1974. Some aspects of the chemistry of fly-ash surfaces. Symp. on Changing Technology of Electrostatic Precipitation. Instit. Fuel, Adelaide, S.A., 16 pp. Davison, R., Natusch, D., Wallace, J. and Evans, C., 1974. Trace elements in fly ash: dependence of concentration on particle size. Environ. Sci. Technol. 8(13): 1107-1113. Dickey, H.P., Zimmerman, J.L., Plinsky, R.O. and Davis, R.D., 1977. Soil survey of Douglas County, Kansas. U.S. Dep. Agric., Soil Conserv. Serv. in cooperation with Kans. Agric. Exp. Stn., Jul. 1977, 74 pp.
356 Faber, J.H., Babcock, A.W. and Spencer, J.D. (Editors), 1976. Proceedings of the Fourth International Ash Utilization Symposium, St. Louis, Miss. Morgantown Energy Res. Cent. E.R.D.A., Morgantown, W.Va. Fisher, G., Prentice, B., Silberman, D., Ondov, J., Biermann, A., Ragaini, R. and McFarland, A., 1978. Physical and morphological studies of size-classified coal fly ash. Environ. Sci. Technol., 12(4): 447-451. Kaakinen, J., Jorden, R.M., Lawasani, M. and West, R., 1975. Trace element behavior in a coal-fired power plant. Prepr. Pap., Natl. Meet., Div. Environ., Chem., Am. Chem. Soc., 15(1): 136-138. Kopsick, D.A., 1980. Geochemistry of leachates from selected coal mining and combustion wastes. M. Thesis, University of Kansas, Lawrence, Kans. (unpublished). Kopsick, D.A. and Angino, E.E., 1980. Solubility of selected major and minor elements from coal gob and fly ash assimulations. Univ. Kans., Lawrence, Kans., Kans. Water Resour. Res. Inst., Contrib. No. 211, 2 1 pp. Linton, R., Loh, A., Natusch, D., Evans, C.A. and Williams, P., 1976. Surface predominance of trace elements in airborne particles. Science, 191: 852-854. Pedlow, T.W., 1974. Cenospheres. U.S. Bur. Mines, Info. Circ. No. 9640, pp. 33-35. Smith, R.D., Campbell, J.A. and Nielson, K.K., 1979. Concentration dependence upon particle size of volatilized elements in fly ash. Environ. Technol., 13(5): 553-558. Surprenant, N., Hall, R., Slater, S., Suza, T., Sussman, M. and Young, C., 1975. Preliminary environmental assessment of conventional combustion systems, Vol. 1. U.S. Environ. Prot. Agency, G.C.A. Corp., Bedford, Mass., Publ. No. GCA-TR-75-26-G, Aug. 1975, 265 pp. Theis, T.L., 1975. Contamination of groundwater by heavy metals from land disposal of fly ash. Natl. Tech. Info. Serv., Progr. Rep. Jun. 1, 1975-Sep. 30, 1975, No. 9416, 1 2 PP. Theis, T.L. and Richter, R.O., 1979. Chemical speciation of heavy metals in power plant ash pond leachate. Environ. Sci. Technol. 14(2): 219-224. Theis, T.L., Westrick, J.D., Hsu, C.L. and Marley, J.J., 1978. Field investigation of trace metals in groundwater from fly ash disposal. J. Water Pollut. Control Fed., 50(11): 24 57-24 6 9.
351
GEOLOGICAL CONSIDERATIONS IN HAZARDOUSWASTE DISPOSAL K. CARTWRIGHT, R.H. GILKESON and T.M. JOHNSON Illinois State Geological Survey, Champaign, IL 61 820 (U.S.A.) (Accepted for publication February 27, 1981)
ABSTRACT Cartwright, K., Gilkeson, R.H. and Johnson, T.M., 1981. Geological considerations in hazardous-waste disposal. In: W. Back and R. Letolle (Guest-Editors), Symposium on Geochemistry of Groundwater - 26th International Geological Congress. J. Hydro]., 54: 357-369. Present regulations assume that long-term isolation of hazardous wastes - including toxic chemical, biological, radioactive, flammable and explosive wastes - may be effected by disposal in landfills that have liners of very low hydraulic conductivity. In reality, total isolation of wastes in humid areas is not possible; some migration of leachate from wastes buried in the gound will always occur. Regulations should provide performance standards applicable on a site-by-site basis rather than rigid criteria for site selection and design. The performance standards should take into account several factors: (1)the categories, segregation, degradation and toxicity of the wastes; ( 2 ) the site hydrogeology, which governs the direction and rate of contaminant transport; ( 3 ) the attenuation of contaminants by geochemical interactions with geologic materials; and (4) the release rate of unattenuated pollutants to surface or groundwater. An adequate monitoring system is essential. The system should both test the extent to which the operation of the site meets performance standards and provide sufficient warning of pollution problems to allow implementation of remedial measures. In recent years there has been a trend away from numerous, small disposal sites toward fewer and larger sites. The size of a disposal site should be based on the attenuation capacity of the geologic material, which has a finite, though generally not well-defined, limit. For slowly degradable wastes, engineered sites with leachate-collection systems appear to be only a temporary solution since the leachate collected will also require final disposal.
INTRODUCTION
The geology and hydrogeology of a site must be carefully considered in planning the disposal of wastes and in assessing the potential problems that may result from some current regulatory criteria. An evaluation process is necessary that considers both the specific character of the wastes for disposal and the specific geologic conditions at the proposed disposal site. Wastes are frequently categorized as nonhazardous and hazardous. Nonhazardous wastes vary from construction debris to municipal waste and sludges, and each affords a very different hazard to the environment. Hazardous waste has not
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yet been satisfactorily defined in total by the U S . Environmental Protection Agency (U.S.E.P.A.), and none is offered here - only suggestions on an approach. This paper deals with wastes that are liquid or will produce aleachate that must be controlled to prevent “environmental degradation”. The term “hazardous wastes” includes the general categories of toxic chemical, biological, radioactive, flammable and explosive wastes. Such wastes, if improperly managed, obviously threaten public health and welfare. There are many sources of hazardous wastes. Approximately 10%of all nonradioactive wastes generated by industry are considered hazardous wastes. About 90% of the hazardous wastes are in the liquid form, of which about 40% are inorganic and 60% are organic (Garland and Mosher, 1975). The quantity of hazardous wastes has been growing at the rate of 5--10% annually; the volume of solid wastes, sludges and liquid concentrations of pollutants from industry that will be disposed of in landfills or lagoons is expected to double during the next ten years. The problems of disposal of wastes into geologic materials are a relatively recent subject of research. Research has, until recently, concentrated on the disposal of municipal refuse (e.g., Andersen and Dornbush, 1967; Apgar and Langmuir, 1971; Hughes et al., 1971; Palmquist and Sendlein, P975;Gilkeson et al., 1978; Baedecker and Back, 1979). Although criteria for disposal of hazardous wastes may differ in some detail from those for disposal of general refuse, the principles developed from this research are applicable to both. Current practices of landfill disposal of wastes, particularly in humid climates, usually generate a noxious liquid leachate. An extensive bibliography on the subject can be found in Cartwright et al. (1981). Studies of the hydrologic systems in fine-grained geologic materials, into which current practices direct most wastes, were almost nonexistent twenty years ago; however, during the past two decades procedures for study of fine-grained materials have been developed to provide the data required to study waste disposal sites (e.g., Farquhar and Rovers, 1975; Griffin et al., 1976, 1977; Cartwright et al., 1981). Case histories also have been developed to help predict chemical and physical changes that may occur as a result of the burial and surface disposal of waste. Investigation of the waste attenuation characteristics of geologic materials is now one of the major areas of research. Attenuation capacity is the material’s ability to remove contaminants from percolating fluids. Approximations of the attenuation capacities of some geologic materials are known for some contaminants in leachates. These approximations provide the general relationships and principles on which judgments can be made; however, it probably will be a number of years before the mechanisms of attenuation are fully understood and the attenuation characteristics of most geologic materials for various contaminants and combination of contaminants are known.
3 59 PRESENT CRITERIA USED BY REGULATORY AGENCIES
Each U.S.A. State has different rules, regulations and guidelines concerning the siting of waste lagoons and landfills. Many categorize sites according to the type of wastes to be received. As an example, landfill sites in Illinois (I.P.C.B., 1973) are divided into five classes on the basis of geologic and groundwater conditions. Hazardous wastes may only be accepted at sites that have the strictest requirements. Illinois regulations require disposal of hazardous wastes in class-I and possibly class-I1 sites, the distinction being made on the basis of the hydraulic conductivity at the bottom and sides. Class-I sites require 3 m of material with a hydraulic conductivity of cm/s or less, and class-I1 sites require the same thickness but a hydraulic conductivity of 5 * cm/s or less. There are no standard specifications for making this measurement, although A.S.T.M. (1970) does have a suggested laboratory method. Laboratory measurement of such low values is difficult; the error in measurement may be quite large, possibly greater than the difference that distinguishes the sites. Field measurement is also difficult, timeconsuming, costly and may be no more accurate. The current guidelines require, in addition to a permeability barrier, a minimum distance of commonly 1 5 0 m from the nearest water well. To protect surface water, siting on a floodplain is prohibited, surface runoff must be controlled, and the site must be at least 1 5 0 m from any body of surface water. If a site does not meet these criteria, engineered modifications in accordance with certain guidelines may be implemented to enhance site conditions. Such modifications frequently involve a liner and a leachatecollection system. U.S.E.P.A. (1980) recently published criteria for disposal of hazardous wastes in Secure Chemical Waste Landfills. The rules provide two alternatives for landfill design: natural geologic containment, or artificial containment using engineered features. Conditions in at least one-half to two-thirds of the U.S.A. will not allow natural geologic containment of hazardous wastes in accordance with U.S.E.P.A. rules for two reasons: (1) the mean annual precipitation is too great; and (2) the water table is too high, especially in geologic materials with low hydraulic conductivity. Therefore, the Federal guidelines require an artificial liner and a leachate-collection system for such a facility. However, on the basis of our understanding of the behaviour of leachate movement in geologic materials, natural containment, with attenuation of leachates by earth materials, is preferable to artificial containment; the reasons for this statement are explained later in this paper (pp. 364365). Additional U.S.E.P.A. requirements for all disposal sites include low hydraulic conductivities cm/s), a minimum of 1.5m between the base of the artificial liner and the water table, and a 1 5 0 m separation from any functioning public or private water supply. The rules also prohibit direct contact between the wastes and surface water, location on a wetland, on a floodplain, in a fault zone, or in the recharge zone of a single-source aquifer.
360 PERFORMANCE STANDARDS FOR DISPOSAL
The objective of regulations governing hazardous waste disposal is protection of surface-water and groundwater resources. Regulations with rigid specifications of geological and hydrological criteria for sites cannot be applied over the entire U.S., or even a region the size of most states. Strict application of some criteria, such as depth t o water table, can actually lead to selection of less suitable sites. Rather, regulations should provide performance standards that the site design must meet to be acceptable and should be applied on a site-by-site basis. In the evaluation of a site, it is the possible effect upon the environment which must be considered. The specific character of the wastes, the geologic materials at the proposed site, and the interaction between the two must be carefully examined. A performance standard should stipulate the effect that a disposal site can have on the surrounding land. The standard must be written t o limit the amount (volume and concentration) of contaminant allowed to be discharged by specifying: (1)the degree of water quality that is required, such as that it must be fit for human consumption or that it must meet specific ion concentrations; (2) the degree t o which the quality of ambient water can be altered; and (3) the exact area that must meet these requirements, such as within certain property lines and including the nearest aquifer or surface water body. If a mixing zone (as for point-source discharge to surface water) is acceptable, the size of that zone should be specified. These specifications must be realistic; that is, a specification of “zero discharge measured a t the waste boundary” cannot be accomplished. These performance standards are an alternative approach t o the use of liners and “total containment”; they permit the selection of disposal sites that will provide for an acceptable amount of attenuation of the toxic constituents from the leachate by interaction with geologic materials at the site. The sites must be carefully designed and engineered t o minimize differential compaction that may occur, and trench covers must be constructed t o control infiltration so that it is equal to or less than the possible migration rate from the site. An understanding of both unsaturated and saturated groundwater flow is required by those people who both design and review sites. To use this approach will minimize the “bathtub” effect (see section on “Site hydrogeology”) and may allow the refuse in the landfill t o leach and compact sooner and shorten the required monitoring time. Lagoons and landfills designed to meet performance standards should take into account five factors: (1) the type of waste to be disposed; (2) the site hydrogeology that governs the direction and rate of contaminant travel; (3) the attenuation of contaminants by geochemical interactions with the geologic materials; (4) the release rate of unattenuated pollutants t o surface or groundwater; and (5) the character of the receiving waters.
361 CONSIDERATION OF THE WASTES
Introduction Regulatory agencies commonly classify wastes as hazardous and nonhazardous, a categorization that is not easily made and often must be arbitrary. In addition, mixed wastes, such as building debris and general municipal refuse, may contain some hazardous materials. The hazardous wastes in building debris may present a problem because current regulations regarding the disposal of building debris are lenient. The hazardous materials that are frequently present in general municipal refuse also may pose problems; general municipal refuse is a mixture of different types of wastes that may promote reactions that enhance the mobility of certain toxic constituents. Presently, the disposal of industrial wastes and sewage sludges into sanitary landfills constructed to receive general municipal refuse is common.
Segregation of wastes The present authors believe that hazardous wastes should be segregated by type where possible; this may sometimes be accomplished by designating sites for particular types of waste disposal. Segregation of wastes allows for better prediction of attenuation characteristics of the geologic material for geochemically similar wastes. This procedure is less complicated than for mixed wastes, and it prevents interaction among incompatible wastes. Chemical reactions between some mixed wastes may increase the mobility of certain toxic constituents. For example, the mobility of most heavy metals is directly related to the pH of the solution; and polychlorinated biphenyls (PCB’s), which are nearly immobile in aqueous solution, become highly mobile in organic solvents (such as carbon tetrachloride). In some instances, an immobile ion may complex with a more mobile ion and migrate with it. Presently, the kinds of complex species that can form and their types of mobility are not very well understood for many hazardous wastes.
Degradation of wastes The nature of degradation of wastes must be considered, i.e. whether by some natural process the waste may change from its present form to some less complex chemical compound and, hopefully, a less noxious form. Categorization into degradable and nondegradable wastes is desirable for all types of waste; it allows the addition of a time factor t o geological and geochemical considerations. The decay/decomposition rate governs the duration of time that is required for isolating the hazardous waste from the environment; the time can range from a few days t o thousands of years. The decay/decomposition rate governs the duration of time that is required for isolating the hazardous waste from the environment; the time can range from a few days
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to thousands of years. The decay/decomposition process may be the result of radioactive decay, organic decomposition or some other process. Wastes that require a long decay/decomposition period (thousands of years) generally should, from a practical hydrogeological point of view, be considered nondegradable. Obviously, this distinction is arbitrary and perhaps should differ from site to site, depending upon the site characteristics. Some geologic materials provide both containment time and attenuation of the wastes; however, wastes buried in landfills ultimately will return byproducts to the environment in some form and concentration. Toxicity of wastes
Because of the ultimate return of waste by-products from disposal sites to the environment, consideration of the toxicity of the wastes is essential. The toxicity of many wastes is not well known and the assigned values are often arbitrary. One approach to this problem of waste classification is to consider the level at which toxicity occurs, i.e. as parts per thousand, million, billion, etc. In evaluating waste for disposal by landfill, the toxicity of the waste should also be related to its decomposition/decay rate. Geologic conditions in many areas may be unsuitable for landfill disposal of some wastes that have slow decomposition/decay rates and contain constituents that are toxic in low concentrations. These wastes may require destructive treatment, deep-well disposal, or shipment t o a site with unique geologic conditions that may make it suitable for landfill disposal of the wastes.
CONSIDERATION O F SITE
Site hydrogeology
The objective of existing regulations that require disposal of hazardous wastes in trenches or lagoons in natural clay materials or with artificial clay liners of very low hydraulic conductivity is to contain the wastes and thereby protect groundwater resources. This approach is valid; however, it can create problems in humid climates where natural precipitation and infiltration of water from the surface exceeds the hydraulic conductivity of the surrounding natural material or liner. When this excess infiltration occurs the disposal site fills with leachate and overflows, a phenomenon called the “bathtub” effect (Hughes et al., 1976; Cartwright et al., 1977). The water table rises into the disposal excavations, even if the original water table was located well below the bottom, eventually filling the trenches or lagoons and sometimes “spilling” out the sides as springs. Thus a site that, on the basis of standards designed for groundwater protection, was suitable for disposal of hazardous wastes may become a hazard. The leachate will then have to be collected, treated, and redisposed of.
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The “bathtub” effect occurs, in part, because most wastes have much higher hydraulic conductivities than the natural material into which they are placed ; they may also have very different unsaturated soil-moisture characteristics. The hydraulic conductivity of some wastes can be reduced by compaction. The “bathtub” effect also occurs because more infiltration enters the disposal excavation than would under normal undisturbed conditions. Trench covers may be constructed to achieve the desired hydraulic conductivity and to limit infiltration for the required period of containment or until compaction of the wastes occurs; however, it is difficult to maintain the trench covers. The covers must withstand attack by plants, weather (freeze-thaw, wet-dry), erosion, and strains caused by consolidation within the trench. Most trench covers are not capable of meeting these demanding requirements without costly long-term maintenance programs. The cover should be designed t o allow for expected consolidation and t o utilize hydrogeological concepts of saturated and unsaturated flow systems present at the site. The present authors believe that the importance of the water table is exaggerated in most regulations (Hughes and Cartwright, 1972). With the goal of protecting groundwater, regulations commonly require that the base of the disposal trenches and lagoons must be situated a specified distance above the water table; therefore, a relatively deep water table is required. However, as pointed out earlier, the water table may be altered by the disposal operations. In much of the humid areas of the U.S.A., deep water tables usually occur in the coarse-grained deposits with relatively high hydraulic conductivities; this is especially true in areas of low topographic relief. These materials may be a potential groundwater resource, rather than a suitable medium for burial of wastes. In these materials, the location of the wastes above the water table does not ensure protection for groundwater from leachate contamination. Research has shown that infiltration through refuse buried in these materials rapidly moves contaminants down to the water table. Shallow water tables generally occur in fine-grained geologic materials having low hydraulic conductivity ; however, the fine-grained material is not a groundwater resource. In the proper hydrogeologic setting, fine-grained materials are well suited for the disposal of hazardous wastes because they are more effective than coarse-grained materials in containing and attenuating wastes and isolating them from aquifers. Site geology
The geologic setting at the disposal site determines whether the leachate will discharge near the trench or flow for great distances through natural geologic materials (Bergstrom, 1968). The pathway to a point of concern and the materials through which the contaminants must pass are critical. Fine-grained matei-ials of low hydraulic conductivity have been found
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to be the most suitable medium for burial and attenuation of wastes. These materials will be more effective in slowing the movement of the leachate and removing contaminants than coarse-grained or fractured materials. Water wells in such materials are usually unsuccessful because the materials, though saturated, do not transmit water fast enough to supply the pump. Where substantial deposits of these materials are present and water-yielding deposits (aquifers) are absent or isolated, conditions are most suitable for attenuation by physicochemical processes for disposal of wastes. Fine-grained geologic materials are frequently found in glacial drift, alluvium, colluvium and other unconsolidated deposits; also shale, claystones, mudstones and other materials of the bedrock are characterized by their fine grain. In addition, many crystalline rocks, evaporites, and such have extremely low hydraulic conductivities. These materials may protect aquifers, a primary objective when establishing criteria for disposal facilities. However, coarse-textured materials have been shown to be acceptable under certain circumstances. It may be possible to dispose of wastes with very low solubility in sand or other coarsetextured materials. Disposal may also be acceptable in areas where migration of contaminants is over moderate distances and attenuation by dilution and limited cation exchange can occur prior to public contact with the contaminants via groundwater development. The geology of the site should be studied in sufficient detail to provide the information required for the site design and to predict the fate of the waste by-products. For some sites, areal geological mapping with limited or no exploratory borings may be sufficient, whereas other sites may require considerable drilling and laboratory testing. Exploration methods making use of piezometers, lysimeters, tensiometers and drilling cores have been developed to provide the required data. Although some drilling will generally be necessary at most hazardous-waste disposal sites, we believe it should be held to a reasonable minimum because each boring represents a possible manmade conduit for the waste by-products to follow. (Strict plugging specifications should be required.) As was mentioned earlier, the hydrology of the site may be altered by disposal trenches. Such consequences should be considered in the proposed site design and operation plan.
Waste attenuation capacity of site Individual constituents in waste leachate may have markedly different mobilities in different geologic materials (Griffin et al., 1976, 1977; Fuller, 1978). Geochemical mechanisms that strongly inhibit the migration of one constituent may have little or no effect on other constituents. The chemical composition of waste leachates vary widely (U.S.E.P.A., 1974) and the interactions with geological materials are complicated; however, the leachates can be considered in terms of their basic constituents. This permits the evaluation of factors affecting attenuation and their disposal in landfills with regard to their impact on the environment and the public health. Although
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only generalizations can be made, it should be recognized that the particular combination of leachate and site will be unique. The attenuating characteristics of fine-grained geologic materials are considered favorable for waste disposal. The properties of geologic materials that are considered most important for attenuation are texture (grain size and structure), pore-size distribution, clay composition and chemical composition. The high percentage of clay content and the distribution of small pores in the fine-grained materials provided a low flux of solution and gases, and long contact times and extensive contact areas between the earth materials and the contaminants. The chemical composition of the geologic material includes its solublespecies-adsorption capacity and matrix composition. The surficial materials are generally low in soluble species that could contribute to the pollutant load. This condition differs from that of arid regions where, in some instances, the natural salt content of surficial materials is sufficient to degrade the groundwater below waste disposal sites. Surficial geologic materials in the eastern half of the U.S.A. frequently contain moderate to high amounts of hydrous oxides and carbonate minerals which, along with clay minerals, provide good adsorption properties for a wide range of chemical species. Some important geochemical mechanisms that attenuate waste leachates are exchange processes in which contaminants are selectively removed from the leachate and replaced by nontoxic constituents from the enclosing geologic materials. Since this is an exchange process, the total concentration of dissolved solids in the leachate does not change greatly. This is a reversible process; adsorbed ions may later be released. Thus, this process may be considered as dilution. The capacity of the geochemical mechanisms in the geologic materials to renovate contaminants from leachates is finite and, if exceeded, will allow the leachate to pass with little change. Therefore, the attenuation capacity of the site’s geologic materials must be the limiting factor for volume of wastes for disposal.
Release rate o f unattenuated contaminants Determining the release rate of unattenuated or poorly attenuated contaminants from the disposal site to surface water or groundwater (aquifer) is a necessary step in evaluating a waste disposal site. A decision must be made as to which ions must be attenuated “totally” and which could be released to the environment in the primary movement of contaminants (Cartwright et al., 1977). A properly designed and operated site promotes the dilution of contaminants by restricting the rate of their release into the environment to some acceptable level at which the concentrations in the receiving natural waters will remain below an acceptable maximum. The calculation of the release rate of leachate from the bottom and sides of a landfill or lagoon and its flow path presents a complex problem. The use
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of a high-speed digital computer to predict the rate and path of fluids may be required; however, preliminary estimates of leakage can be made by use of the Darcy equation. If we consider the accuracy of most of the input data, this may be almost as reliable as more complex models. Alterations in the hydrogeologic system caused by the presence of the landfill must be considered, and the leakage calculated for the landfill must then be compared in volume to the receiving waters. The time required for contaminants leached from the wastes to arrive at some point away from the source may be important under certain circumstances. The rate of water movement provides an estimate of the rate of travel of nonattenuated contaminants. The rate calculation must take into account the pore volume and structure of the geologic materials in which flow occurs. The attenuation characteristics of the geologic materials will retard this rate for a specific contaminant by an amount related to its attenuation factor. This factor is of great importance for wastes that undergo decay or decomposition during flow through a porous medium. Geochemical mechanisms that attenuate waste by-products as they migrate through fine-grained geologic materials have been discussed briefly. If the composition of the leachate is known, an approximation of the attenuation factor can be made in the laboratory. The distribution coefficient, or the less complicated retardation factor, can be measured for the samples of geologic materials from the disposal site and for the leachate from the waste. The values measured in the laboratory, in practice, present some difficulty, as they are not constant but vary as the concentration of waste byproducts changes in the leachate; nevertheless, these factors do provide data upon which a judgment can be made. However, the composition of the leachate is rarely known prior to disposal. An approximate attenuation factor for the by-products of some wastes may also be calculated from known general relationships and principles. More research will be necessary before the mechanisms of attenuation can be understood and the attenuation characteristics for various contaminants by most geologic materials are known.
DISCUSSION
For both hazardous and municipal wastes, there has been a trend in recent years from numerous, widely dispersed, small disposal sites to few and larger sites. This strategy should be used with extreme caution, especially if both large and small sites are judged by the same design standards. The authors believe that the use of performance standards rather than design standards are essential under these circumstances. The attenuation capacity of any geologic materials has a limit which, if exceeded by the volume of leachate that enters the material, will allow contaminants to pass almost unretarded. Unfortunately, there are insufficient data on the attenuation capacities of
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geologic materials for most leachate constituents to clearly define this limit. Larger land disposal sites are more likely than the smaller sites to exceed this limit. For slowly degradable or nondegradable wastes, we view the trend to engineered sites with leachate collection systems as an interim disposal technique. Eventually the leachate collected will have t o be redisposed of at a final disposal site and perhaps at great expense. Such engineered sites may be suitable for the disposal of degradable wastes where isolation of wastes from the environment is not necessary for long periods of time. Sites may be engineered so as to reduce the volume of wastes that need to be transferred for final disposal; such a site may be appropriate in densely populated regions. The leachate collected from these sites eventually presents a disposal problem. Destructive treatment may be difficult and processing the leachate in a standard municipal waste treatment plant may only dilute hazardous substances, possibly causing the sludge from the treatment plant to become hazardous. This discussion has considered waste in general. Radioactive materials, mentioned several times in the paper, represent a special type of hazardous waste that is often given special consideration. In the authors’ opinion, such special consideration is not always necessary ; the principles discussed in this paper apply equally to low-level radioactive-waste materials and some shortlived high-level wastes. An adequate monitoring system is essential to the operation of hazardouswaste disposal sites. The monitoring system should test the extent to which the operation of the site meets performance standards of the site design. Also, the monitoring system should provide sufficient warning of potential pollution problems so that remedial measures, called for in a contingency plan, can be instituted. A contingency plan should be part of all disposal site designs. These measures should be specified in the site design to ensure against environmental degradation in the event that operation of the disposal site fails to meet performance standards. Most regulations prohibit burying wastes in floodplains of rivers; however, these regulations do not recognize that often the floodplain does not occupy the entire river valley. The term floodplain and its application should be clearly defined. Where river valleys are underlain by shallow high-capacity aquifers, geologic conditions are generally not suited for the disposal of hazardous wastes. Sites that are located in the valley, well out of the reach of erosion by flood waters and are not underlain by coarse sands and gravel with high conductivities would, under some circumstances, be suited to the disposal of certain types of hazardous wastes. These sites should be underlain by fine-grained geologic materials that have adequate attenuation capacity and should permit only the slow release of contaminants to the environment at an acceptable rate. The migration of contaminants from disposal sites in this hydrogeologic setting will follow along short well-defined flow paths, and consequently there will be a limited area of contamination that
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may be relatively easy to monitor. At these sites the slow release of poorly attenuated or unattenuated contaminants t o the environment where there is a comparatively high volume of receiving water will provide very high dilution rates. The proper operation of a waste disposal site in the floodplain setting is very similar in concept to the current operation of waste treatment plants that discharge directly to surface water. Deep bedrock disposal wells were mentioned briefly in this paper as a possible method for disposal of some wastes. These disposal wells are drilled into deep bedrock formations containing saline groundwater (greater than lo4 ppm of total dissolved solids), and the wastes for disposal are pumped into these formations. Engineering aspects of constructing deep disposal wells are well understood ; however, determining the feasibility of deep disposal wells requires careful consideration of the nature and volume of wastes for disposal and the geologic and hydrologic conditions of the disposal zone. With thorough site investigation, careful operation and adequate monitoring, deep-well disposal has the potential to provide excellent isolation from the environment for limited quantities of highly toxic wastes.
REFERENCES Andersen, J.R. and Dornbush, J.N., 1967. Influence of sanitary landfill on groundwater quality. J.Am. Water Works Assoc., 59: 457-470. Apgar, M.A. and Langmuir, D., 1971. Groundwater pollution potential of a landfill above the water table. Ground Water, 9: 6. A.S.T.M. (American Society for Testing Materials), 1970. Special procedures for testing soil and rock for engineering purposes. Am. SOC.Test. Mater., Spec. Tech. Publ., 479: 1 41-1 4 5. Baedecker, M.J. and Back, W., 1979. Hydrogeological processes and chemical reactions at a landfill. Ground Water, 17(5): 429-437. Bergstrom, R.E., 1968. Disposal of wastes: scientific and administrative considerations. Ill. State Geol. Surv., Environ. Geol. Note 20, 1 2 p. Cartwright, K., Griffin, R.A. and Gilkeson, R.H., 1977. Migration of landfill leachate through glacial tills. Ground Water, 15(4): 294-305. Cartwright, K., Gilkeson, R.H., Griffin, R.A., Johnson, T.M., Lindorff, D.E. and DuMontelle, P.B., 1981. Hydrogeologic considerations in hazardous-wastes disposal in Illinois. Ill. State Geol. Surv., Environ. Note 94, 20 pp. Farquhar, G.J. and Rovers, F.A., 1975. Leachate attenuation in undisturbed and remoulded soils. Proc. Res. Symp. Gas and Leachate from Landfills: Formation, Collection, and Treatment, New Brunswick, N.J., March 15-16, 1975, U.S. Environ. Prot. Agency, Natl. Environ. Res. Cent., Cincinnati, Ohio. Fuller, W.H., 1978. Investigation of landfill leachate pollutant attenuation by soils. U.S. Environ. Prot. Agency, Cincinnati, Ohio, EPA-60012-78-158. Garland, G.A. and Mosher, D.C., 1975. Leachate effects from improper land disposal. Waste Age, 6 : 42-48. Gilkeson, R.H., Cartwright, K., Follmer, L.R. and Johnson, T.M., 1978. Hydrogeologic investigation of groundwater contamination from land disposal of toxic wastes in Ogle County, Illinois. Ill. State Geol. Surv., Repr. 1978-D (reprint from Proc. 15th Annu. Eng. Geol. Soils Eng., Symp., Pocatello, Idaho, 1977, p. 17-28). Griffin, R.A., Cartwright, K., Shimp, N.F., Steele, J.D., Ruch, R.R., White, W.A.,
369 Hughes, G.M. and Gilkeson, R.H., 1976. Attenuation of pollutants in municipal landfill leachate by clay minerals, Part 1. Column leaching and field verification. Ill. State Geol. Surv., Environ. Geol. Note 78, 34 pp. Griffin, R.A., Frost, R.R., Au, A.K., Robinson, G. and Shimp, N.F., 1977. Attenuation of pollutants in municipal landfill leachate by clay minerals, Part 2. Heavy metal adsorption. Ill. State Geol. Surv., Environ. Geol. Note 79. 47 pp. Hughes, G.M. and Cartwright, K., 1972. Scientific and administrative criteria for shallow waste disposal. Ciu. Eng., Am. SOC.Civ. Eng., 42(3): 70-73. Hughes, G.M., Landon, R.A. and Farvolden, R.N., 1971. Hydrogeology of solid waste disposal sites in northeastern Illinois. U.S. Environ. Prot. Agency, Solid Waste Manage. Ser., Rep. SW-l2d, 1 5 4 pp. Hughes, G.M., Schleicher, J.A. and Cartwright, K., 1976. Supplement to the final report on the hydrogeology of solid waste disposal sites in northeastern Illinoise. Ill. State Geol. Surv., Environ. Geol. Note 80. 24 pp. I.P.C.B. (Illinois Pollution Control Board), 1973. Rules and regulations. In: Solid Waste, Ch. 7. Ill. Pollut. Control Board, Springfield, 111. Palmquist, R. and Sendlein, L.V.A., 1975. The configuration of contamination enclaves from refuse disposal sites on floodplains. Ground Water, 13(2): 167-181. U.S.E.P.A. (U.S. Environmental Protection Agency), 1974. Summary report: Gas and leachate from land disposal of municipal solid waste. Solid Hazard. Waste Res. Div., Munic. Environ. Res. Lab., Cincinnati, Ohio. U.S.E.P.A. (U.S. Environmental Protection Agency), 1977. Waste Disposal practices and their effects on ground water. U.S. Environ. Prot. Agency, Off. Water Supply-off. Solid Waste Manage. Progr., Rep. to U.S. Congress, edited by D.W. Miller (Rep. also published in book form as Waste Disposal Effects on Ground Water, 1980, Premier Press, Berkeley, Calif., 512 pp.)
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