Chemical Evolution of Groundwater in A Drainage Basin of Holocene Age, East-Central Alberta, Canada

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

.---701----1700-

HYDRAULIC HEAD

(METRES ABOVE SEA L E V E L )

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 from -4 cm/s in the most productive aquifer at the eroded driftbedrock contact zone, to -7 * cm/s in the sandy lenses of the upper cm/s in the deep sands of the Belly Belly River Formation, and t o -2 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

-

-

Y

4,

701

701

670

670

640

640

610

610

P 579

579

>

G

549

549

'

5181 488

1518 1.6

4.8

3.2

6.4

,

8.0

I

488

(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

LRR

I .'6

LEGEND

-

-600-

....

~

I 3.2

6.4

8.0

WATER TABLE

-

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

~

HYDROCHEHICAL ZONES

I

NaHCO ( G I )

I1

NaHC03S04

111

418 (km)

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. , Caago (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

Ho.

Lab

NO.

HC03

Sa rn p l ~n g

CI

SO4

Date

NO1

S102,

F1:Hd

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

1110

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

476 142 361

227 220 216

20

5 0

24 18

2 8 0.0 7.1

I 1 I 4 0 5

7 7 9.4 9 5

8 6 9 4 9.0

1245 1290 1250

8 0 2 8 9 5

107.9 107.9

7.9 4.8

2.6 1.0

12

6.8 4.1

8.0 8 6

-

8.1

I2

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

400

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

210 212

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

1.0

1.0

1.3

971i15 .. .

I

9.1 9 I

751061 760116 760256

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

751062 760139 760241

8/14/75 4130176 6 1 4/76

27.3 271 27.1

76.0 83.0 83.0

24.0

AH-5

29.0 29 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

l01.0 99.0

50 0

0.0

614 654

207 227

150

2.9 4.2

0.0

515 592

126

276 149 112

108

6.7 6 2 7.1

215

126f107

. 4.89.7

71

5.4

0.0

0.0

72.0

122.0

1270 55+7

--

0.0

451t161

-

8 1

1500

5.0

I75

10 28 26

1.4 I 1 0 6

0 8 1.6 I 0

9.5

10.1

9.6 9.8 9.7

I050 I015 830

6.0

I11

197.7?94 ~~

71.76

4

0.8

0 0 0.0 0.0

390

14.0

20 24 23.8

6 1 5.8 6.7

0.0 0.0 0.0

40.0 42.0 44.0

15

6.7 5 4 5 6

0 0

325

14.8

91.0

181

4

52

4 8

127 456 476

71 77 87

10

0.0

419 417 429

I10 112

0.0 0.0

159 600

78 66

0 0

2 01t2 4

71 82

6 7 5.4 7 9

6.7 14.6

8 0 5 0

0.5 1 0

25 20

95.0

5.8

8

41 . 3

12.5

4 1

6

6.5

183

155

4.2

0.0

1.8

1.8.4 10.9

8

X f I

760144 760240

4/10/76 6 1 4/76

26.8 26.8

15.0

21 0

68.1*27

11.9 7 7 23.I!9.5

291

143

73.8.78

6.7 5.4 5.7t2 6

0.0 0.0

0.0

447 669

__

400-91

365 256

-

8.0

5.0

..

9 I t . .~ W 1479f490 . 7 222 8.9r.6

10 2 11.0

8.0 7.5 7.8

7.4 7.8 7.5

600 700 750

8.0 16.0 8 0

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

180

9.0

-

0.0

14.0

9 5 14.8

7.4 7.6 7.2

7.8 7.4 7.2

575 780 850

10

6 8

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

16.8 181

7.5 7.7

675 725

11.0

750 1500

I5 0 9.0

2

4

8

4.0 0 0

1.6

28

24

7.4r6 2

0.0

0.8

1.6t1.7

7.4

9.8 12.8r4

855

7.5 7.7 ~

l07i71

1550

8.2

-

~~~~~~~~~~~~~~~

AH-IO(YI M-I0

-

8 I4

67 I50 96

98

80

I 9 1.1

118

137 176 400

0 0 0 0

17.5

188

2.5 1.3

5.8

31.0

195

155

180

9.8i19.6

AH-2IWl AH-? Iw-2

AH-5

i’c)

7 I 6 7 6 7

29.0 29.0 29.0

I

Temp

215 204 218

8/14/76 1/13/76 6/8/76

X

-11

19 I 14.6 13.7

751060 760019 760249

RAH-4lQl

IY mho Crn

I-IIPI 1-1 1-1

1-2 1-2

Cond

pH

~ a b

II/O/YR

7.7

8.1

7.5t.J

~

7.7 7.8 7.5’.2

1 8

12.0

~

740t190

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)

f

4 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 EQUILIBRIUM P

(Am)

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

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

60

('//oo, SMOW)

6IS0

(%,

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

X = -19.7

-140 -152 -132 -125 -133 -141 -138 -147 -161

+6

f4

+4

f3 f3

+3 f4 +3 f4

f4

X = -141

+ 11

f 0.1

f 0.2

(b)

f 0.2 f 0.1

f0.2 f 0.1 f 0.1 f 0.4 + 0.3 f 0.3

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

X = -20.5

3H (TU)

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 , and averaged 5yoo (H.R. Krouse, pers. Council ranged from -1 t o 13yoO 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 com-

positions 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

4.0

.DRIFT

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 10

GROUNMIATER FROM DRIFT AQUIFERS

x

8

-

0.18 f 0.12

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

ooo -1

-5

0

O

"

DL""

0 DRIF

O

-4

-0

0

-3 log

-2

I

-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

(

x 1 SMOW

- 549

I

549.

-16

518-

-518

(km) LEGEND AND RANGES OF 6 VALUES

I na mb,Ob JJa

Range 60-18 -18.5 to - 1 9 . 0

-WATER Range 6D

- 1 3 8 . 4 t o -147.1

-16.8 t o -22.9

-125

to -167.8

- 1 8 . 9 t o -20.3

-150.2

to

-19.9 t o -22.9

-141.4 t o -167.8

-153.9

_-

TABLE

-Boundaries

I

na

nb

ma

mb

Fig. 19. Local distribution of stable isotopes.

of dynamic zones

P o o r l y mixed, T = >3D,OOO YEP

Poorly mixed

Yell mixed P o o r l y mixed Yell Mixed

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

0 100 5 10 20 40 60 80

Total concentration in solution (meq./l) Ca

Mg

Na

K

C1

SO4

(HC03+C03)

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)

CaC03 precipitated mol)

Temper

Run A :

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 cm/s and 1.25 * cm (fine of p and dS0 are, respectively 2.77 * 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 mol/cm3) divided by the velocity, equal to 0.311 molcm-2 s-'. (2.08 * 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 to be 5.91 * molcm-2 s - l . This calculation, indicates that molecular 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).