Andre Et Al 2005

Journal of Hydrology 305 (2005) 4062

www.elsevier.com/locate/jhydrol

Using geochemical data and modelling to enhance the understanding of groundwater flow in a regional deep aquifer, Aquitaine Basin, south-west of France

L. Andre*, M. Franceschi, P. Pouchan, O. AtteiaInstitut EGID Bordeaux 3, 1 Allee Daguin, 33607 Pessac, France

Received 15 July 2003; revised 9 August 2004; accepted 19 August 2004

Abstract

In deep aquifers the complex flow pattern originating from the geological structure often leads to difficult predictions of water origin, determination of the main flow paths, potential mixing of waters. All these uncertainties prevent an efficient management of the resource. In the context of the Aquitaine basin an original modelling approach suggests that geochemical data can be used to identify flow directions where geological and hydrogeological data are too scarce to provide sufficient information.

In the Eocene sands aquifer, the major patterns of groundwater geochemistry suggest the presence of two distinct areas within the aquifer. In the north and the east, waters exhibit sodium bicarbonate or sodium sulphate facies, and moderate total dissolved solids related to high sulphate concentrations. In the south, waters are characterised by calcium bicarbonate facies and low total dissolved solids.

Sulphur isotopic ratios provided key information on the origin of sulphur in solution (meteoric, gypsum dissolution, pyrite oxidation) and also on the intensity of the geochemical processes involved in the dissolution of minerals and the concentration evolution.

A geochemical model was developed to analyse the processes generating the chemical composition of each sampled water. At the aquifer scale, four main geochemical processesdissolution, redox, acidbase reaction, exchangeof varying intensity could explain most of the observed spatial variability in groundwater composition.

Among several potential reaction schemes at each point, only one allowed to reproduce the independent variables (pH and 13C). The developed model was used to select the most probable water pathways at the aquifer scale. In this context, geochemistry clearly demonstrates the role played by subsurface structures on water flow velocities and residence time in their vicinity. In addition, the concentrations of several ions could only be justified by the aquitardaquifer interactions.

q 2004 Elsevier B.V. All rights reserved.

Keywords: Water geochemistry; Groundwater hydrogeology; Isotopes; Modelling; Ion exchange; Leakage; Aquitaine

* Corresponding author. Address: Institut EGID Bordeaux 3, 1 Allee Duguin, Pessac 33607, France. Tel.:C33 5 57 12 10 17; fax: C33 5 57 12 10 01.

E-mail address: [email protected] (L. Andre).

0022-1694/$ - see front matter q 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.jhydrol.2004.08.027

L. Andre et al. / Journal of Hydrology 305 (2005) 406241

1. Introduction

The chemical composition of groundwater is controlled by many factors that include composition of precipitation, geological structure and mineralogy of the watersheds and aquifers, and geochemical processes within the aquifer. The interaction of all factors leads to various water facies. Usually, major ions studies are used to define hydrochemical facies of waters and the spatial variability can provide insight into aquifer heterogeneity and connectivity (Murray, 1996; Rosen and Jones, 1998). With the development of geochemical modelling, trace, major and isotopic elements are used to infer the physical and chemical processes controlling the water chemistry and to delineate flow paths in aquifer (Eberts and George, 2000; Plummer and Sprinckle, 2001; Guler and Thyne, 2004; Andre, 2002).

The Eocene sands aquifer, located in the south Aquitaine Basin, constitutes an important water resource used for various purposes (thermal water, potable water.). This aquifer has been mostly studied since the end of the 50s along with the development of oil exploration in this part of France. The production wells and exploratory borings were used to draw structural maps and to define the hydrodynamic properties of the aquifer. More recently, the conflicts for water use have led to detailed studies of the hydrogeology of the aquifer (Labat, 1998). The hydrodynamic characterisation allowed identification of some main flow lines and the approximate behaviour of the aquifer near geologic structures.

The water chemistry has been first approached by a description of the spatial variation of some ions and by the calculation of the residence time (Blavoux et al., 1993). The approach presented here uses geochemical indicators and modelling, both linked to the structural or deep sedimentologic features in order to locate potential mass transfers. After a detailed analysis of the regional geology and the sampling locations used for observations, the physico-chemical, chemical and isotopic data are presented. Specific attention will be paid to isotopic data, particularly 34S.

Geochemical modelling is needed to deal with the interactions of several elements and dissolution processes. A pre-requisite step is devoted to a detailed inspection of the reactions that can occur, or not, in

the studied aquifer. The modelling results are then validated using independent variables. Particular attention is paid to the points having an original behaviour or water composition that could not be predicted using the first group of assumptions.

The results of the geochemical modelling step are then interpreted in terms of potential flow at the aquifer scale. Similar regional approaches were used by several authors to interpret the geochemical evolution in regional aquifers (e.g. Hendry and Schwartz, 1990; Weaver and Bahr, 1991; Gerla, 1992; Sracek and Hirata, 2002). The quantitative approach of the reaction linked with the information on the aquifer solid and isotope data allowed to enhance the confidence in geochemical models. The modelling results were able to locate flow barriers at the regional scale and to elucidate the circulations in complex geological structure. The developed approach, involving the analysis of geochemical reactions, validation variables and aquitardaquifer interactions, seems to be applicable to other basins.

2. Hydrogeologic and geologic settings

The Aquitaine Basin is limited in the east by the foothills of Montagne Noire, in the south by the North Pyrenean Piedmont, in the west by the Atlantic Ocean and in the north by the Poitou Plateau. We will be interested here in the southern sector of the basin located to the south of the Garonne river.

The Eocene sands aquifer, located in this part of the Aquitaine Basin, constitutes a major aquifer used for drinking water, agriculture, gas storage and as a thermal resource. This aquifer extends over 150 km from east to west and 200 km from south to north and constitutes a part of a multi-layer system. To manage this resource, the interactions with the under- and over-lying aquifers are fairly important because the Eocene sands aquifer has very restricted outcrops. The average thickness of the quartz sand deposit is estimated as 50 m, with a high porosity of 2035%, and average permeability, estimated from aquifer testing and modelling results to approximately 3!10K5 m/s (Labat, 1998). The average interstitial velocity, using a gradient of 0.001 and an effective porosity of 20%, is close to 5 m/y. Groundwater flow is mainly oriented from SE towards NW but outflow

42L. Andre et al. / Journal of Hydrology 305 (2005) 4062

Fig. 1. The study region and the piezometric map of the Eocene sands aquifer. (Institut EGID, 1999)

from the aquifer is not completely identified (Fig. 1). In the centre of the basin, the estimated age of the groundwater (using 14C data) is close to 2025 ky. This is consistent with the effective advection calculated above.

Sands deposits mineralogy is very poor, containing mainly quartz, augmented with calcite, and, at places, dolomite and K-feldspars (Andre, 2002). Detrital sediments were eroded from the Massif Central, the Montagne Noire, and the emerging Pyrenees moun-tains. These eroded sediments were deposited in vast marshy plains. Geological structures, identified by the exploration, have affected this surface. They include local structures (e.g. the domes of Garlin or Saint-Medard) as well as larger structures (e.g. the Audignon anticline, located at the western border of

the aquifer, and the CeltAquitaine flexure which seems to divide the aquifer in two quite distinct zones) (Fig. 2). It seems that most of the structures are related to the deformation of deep Triassic sands (Rey, 1995; Serrano, 2001).

The DanoPaleocene calcareous and dolomitic aquifer, lying beneath the Eocene sands, is separated from them by clay deposits of variable thickness, from 10 to 100 m. The Eocene sands aquifer is covered by several hundreds of meters of Tertiary molasses. Aquitard mineralogy, although less well investigated, is richer, with quartz, feldspars, mica and several clay types (detrital limestone levels with sandy-clay deposits). At the bottom of the molasse deposits, crystallised gypsum has been observed in cuttings at many places.


Fig. 2. Location of the sampling points and major faults. The numbers with symbols correspond to sampling points having used to the geochemical modelling.

3. Water chemistry

3.1. Sampling and analyses

Waters were sampled during various field cam-paigns distributed over two years. This is due to the irregular use of some wells (irrigation, geothermal power heat). Owing to the residence time of the water in the aquifer, the time lags may not play a dominant role on the chemical composition of the waters and consecutive samplings at some wells showed a constant chemical composition.

Several parameters were measured in the field: temperature, pH, alkalinity and redox potential measurements were made, shielded from the air, in a flow through cell during sampling.

Electrodes were calibrated in the laboratory and in the field. For titration of the total alkalinity

and the total sulphides, the temperature of the measuring cell was hold at the sample temperature. Alkalinity, determined by using a fixed endpoint methodology (pH 4.3), was titrated by 0.1 M hydro-chloric acid with a precision of 2%. For titration of the species of sulphur, the method developed by Boulegue and Popoff (1979) was used (Ohayon-Courtes, 1992). The sulphur species were measured in strongly basic environment by a solution of HgCl2. The potentio-metric follow-up, between a specific electrode Ag/Ag2S and a double junction reference electrode (AgCl) allowed the determination of an inflexion point corresponding to the sum H2SCHSK. This titration allows the determination of concentrations with a precision of 56%.

Analyses of sulphur isotopes within reduced and oxidised forms of sulphur in solution was undertaken. The reduced forms were precipitated in the form of


zinc sulphide by addition of zinc acetate. This procedure precludes any later oxidation of sulphides to sulphates and allows a single sample to be used for 34S analyses of both sulphides and sulphates.

Anions (ClK, SO24K, F, BrK) and cations (Ca2C, Mg2C, NaC, KC) concentrations in filtered samples (acidified for cations) were measured by ionic chromatography on a Dionex DX-120. Trace elements (Al, Mn, Fe, Ba, Sr, Rb,) were measured in filtered and acidified samples by graphite-furnace atomic absorption (Perkin Elmer SIMAA 6000).

Isotope analyses were conducted during two sampling campaigns in 20002001 as part of this study. Analyses were performed by the Environment Isotopes laboratory of Waterloo, in Canada. Isotopes analyses from another study (Blavoux et al., 1993) were also used for the interpretation.

3.2. Results

The analytical results are given in Tables 1 and 2. Waters from the Eocene sands aquifer mostly show a calcium bicarbonate facies, a common facies for waters stemming from deep aquifers (Marini et al., 2000). Sodium bicarbonate waters remain excep-tional and sodium sulphate waters occur at a few wells. Most of the waters have a total dissolved solid concentration lower than 1000 mg/L. This is uncom-mon considering the depth of the aquifer and the residence times. The chemical composition interpre-tation clearly evidences three main geochemical facies (Fig. 3), a calcium bicarbonate facies, a sodium bicarbonate facies with low total dissolved solids (e.g. sample from well 13) or average mineralised (e.g. sample from well 17) and a sodium sulphate facies (e.g. sample from well 4).

The chloride concentrations vary over the studied area from 0.20 to more than 14 meq/L. Two zones can be identified in this aquifer. One zone is located in the south of the basin, with chloride concentrations on the order of 0.150.30 meq/L with a few samples reaching concentrations of about 1.50 meq/L (Wells 19, 21, 22). Concentrations show no clear trend, with values close to the concentrations observed for waters from the recharge area. Schoeller (1956) quotes, as an average chloride content of rain water, a value close to 0.08 meq/L. With a factor of concentration close to 2.5, corresponding to local evaporation rates,

we would obtain an average concentration of 0.20 meq/L. The other zone includes all the northern part of the aquifer. The chloride concentrations range from 1.4 to 16 meq/L, which seems to indicate another processes occurring in solution.

Sulphur is present under various forms, the most common being dissolved sulphate. Sulphide can also

be present but the concentrationsof this species

are always relatively low (lowerthan 1 mg/L),

and represent at most only 10% of the quantity of total sulphur. Sulphide arises, mostly, from local phenomena of bioreduction.

The distribution of the sulphate concentrations shows the same regional zonation as the one observed for chloride. In the south, weakly mineralised waters have concentrations lower than 0.50 meq/L, with, contrasting with chloride, an increase from south to north, the concentration going from 0.20 meq/L near the outcrops to 1.50 meq/L at well 11. In the north, the concentrations lie between 2 and 23 meq/L.

Sulphur isotope analyses (Andre et al., 2002) were made on dissolved sulphate. Gypsum samples, collected from the top of the aquifer, were also analysed. The mean values of d34S and d18O measured on this gypsum are respectively 12.72G1.20 CDT and 14.9 SMOW. These values agree with those quoted in the literature for sulphate evaporates (Claypool et al., 1980).

The spatial distribution of the d34S values observed in sulphates from waters indicates four distinct zones (Fig. 4):

Zone A, along the edge of the Pyrenees, in the zone of outcrop of the aquifer, where the d34S values are

close to values found for sulphates from precipi-tation (C3.2!d34S!C8.2 CDT) (Pearson and Rightmire, 1980). The value of d18O(SO4) is

slightly lower than the expected precipitation values (C9!d18O(SO4)!C13.34 with respect to SMOW) (Trembaczowski and Halas, 1984) and results probably from mixing of various contributions in this zone of outcrops (Pelissier-Hermitte et al., 2000).

Zone B, the largest zone, which extends from the Montagne Noire, in the East, towards the centre of the basin. Values of d34S and 18O(SO4) are close to those measured in the gypsum crystals sampled in the molasse cover, at the top of the aquifer.

Table 1

Chemical composition of waters from the Eocene sands aquiferNo.Wells namesT (8C)Cond.pHEhCa2CMg2CNaCKCHCO3KClKNO3KSO42K[H2S]

(mS/cm)corrected(mg/L)(mg/L)(mg/L)(mg/L)(mg/L(mg/L)(mg/L(mg/L(!10K6

(mV)HCO3)NO3)SO4)mol/L)

1Barbotan 10237.23907.0766.950.606.8812.165.83170.309.48040.01NM

2Barbotan 10336.93727.2573.148.296.7011.745.99176.008.71031.62NM

3Barbotan Lotus 2314107.106.8154.488.1011.375.80183.708.38051.39NM

4Lectoure42.530107.2072.644.6413.49548.2020.30367.70321.200651.800.00

5Lussagnet 5745.43467.29K1243.717.129.375.52178.546.59020.322.23

6Izaute 535.53197.35K7.841.855.239.334.87163.697.95015.2215.57

7Nogaro 251.32957.35K5938.464.9414.305.02155.068.16028.1912.25

8Gondrin42.43497.4535.731.067.4423.195.76163.007.10027.670.00

9Demu52.32967.32K13032.874.6417.116.24159.707.4107.939.54

10Beaucaire34.322807.449.18134.350.631618205.4449.93010691.20

11Castera-Verduzan30.44607.4876.344.3214.5521.386.85178.607.09079.200.76

12Plehaut27.93917.4530.342.4810.6218.588.98167.116.98049.110.00

13Eugenie les Bains 219.03577.94K4722.786.2642.887.16201.527.44010.9826.30

(ELB 2)

14Geaune30.93367.43K8.153.085.099.204.32181.635.83013.530.00

15Geaune Pecorade 10132.43307.4138.749.844.136.083.52185.305.65012.770.00

16Geaune Bats32.53347.317.6150.604.275.273.69184.355.27011.320.83

17Blagnac (piscine)49.613777.20K2412.223.23284.6010.28425.46126.300131.003.06

18Lalbarede25.78577.77K1812.226.82165.205.40281.4080.15058.380.00

19Garlin27.23517.5028.834.398.5825.503.63188.1615.0304.631.87

20Lespielle 126.52617.39NM42.195.109.932.82173.557.46010.8814.82

21Lamaze`re50.14947.3134132.189.9757.1510.33292.387.26025.900.00

22Saint Medard20.63058.42K4522.3011.9115.605.26143.0611.71019.161.04

23Bordes 313.44847.4636287.164.085.021.40223.7411.9241.6914.430.00

24Grignols21.416597.08174123.945.6141.9014.7202.52167.400423.50.00


45


Table 2

Isotopic composition of waters from the Eocene sands aquifer

Wellsd18O(H2O)2H3H (U.T.)d13CA14Cd18O(SO4)d34S(SO4)d34S(H2S)

number( SMOW)( SMOW)( PDB)(p.c.m.)( SMOW)( CDT)( CDT)

1K8.90K55.70!0.8K12.05.5G0.34.41K20.124.28

2K8.90K57.400.2G0.3K12.05.1G0.55.33K16.18NM

3K8.60K57.001.3K11.924.2G0.43.95K15.80K32.25

4K7.20K45.000G0.4K9.982.5G0.216.1312.23/

5K9.40K61.50!0.8K11.22.3G0.39.19K3.21K15.48

6K8.78K57.42!0.8G0.4K10.968.93G0.119.11K1.63K24.05

7K8.70K49.30!0.8K13.53.3G0.212.618.76K20.17

8K8.00K50.400G0.3K12.92.1G0.315.411.56/

9K7.69K51.30!0.7K14.44.5G0.412.8418.93K16.33

10K7.22K49.44!0.8G0.5K6.814.94G0.0813.6615.05NM

11K8.18K53.37!0.8G0.4K10.054.75G0.1315.1510.32NM

12K8.20K50.20!0.8K13.84.5G0.515.289.76/

13K9.56K63.67!0.8G0.4K12.585.24G0.1214.1433.18K10.55

14K8.60K54.200.5G0.3K11.46.9G0.58.30K3.20/

15K7.29K53.03! 0.8G0.5K11.299.72G0.115.98K4.01NM

16K7.56K52.54!0.8G0.4K11.512.81G0.136.63K2.98NM

17K7.35K50.37!0.8G0.4K8.411.44G0.0715.2313.237.86

18K7.32K48.86!0.8G0.4K11.982.89G0.0617.2313.0610.78

19K8.10K53.900.80G0.3K14.5!0.712.8513.01K28.05

20K8.00K50.801.5G0.5K14.33.0G0.811.965.79NM

21K8.50K52.40!0.8K6.03.6G0.214.9317.50/

22K7.31K51.720.9G0.5K8.4813.09G0.214.2724.37NM

23K7.30K46.1016.0G0.8K11.977.3G1.05.134.87/

24K7.53K51.10!0.8G0.5K4.150.96G0.0712.8315.70/

NM, not measured.

Zone C, a stretched shape oriented SWNE, shows negative to very negative values (from K0.4 to

K20.1 CDT), with a decrease from the south northward. The values of d18O(SO4) follow the same trend with a decrease along the same axis. These values seem to reflect a different origin of dissolved sulphates than in zone B, a plausible explanation being the oxidation of sulphide minerals (Dazy et al., 1980).

One particular point in the west (e.g. Well 13Zone D), close to the aquifer boundary, shows significant enrichments in sulphur 34 (d34SO33 CDT). A proposed explanation for this is the action of bioreduction processes.

4. Modelling strategy

A geochemical model is needed to quantify the relative importance of each chemical reaction at

the sampled wells. During this first geochemical modelling step, only the system Na, Ca, Mg, Cl, S, CO3, H is considered. Si and K concentrations are not included due to the lack of information on the composition of the clay fraction of the aquifer, not accessible from cuttings samples. Moreover, the silicate equilibrium does not significantly modify the acidbase and redox equilibrium, and can thus be neglected. Trace elements such as Sr, Br, F and Ba might be integrated in future works after modelling the behaviour of the major species in solution.

4.1. Choice of mineral phases

The Eocene sands appeared, in some drillings, clean, with translucent grains, pyrite and lignite traces. At places the sandy series include argillaceous or argillaceous-sandstones interbeds. The few min-erals identified are quartz, calcite, feldspars, kaolinite, mixture of iron oxides/hydroxides, pyrite in some


Fig. 3. Trilinear Piper diagram of water chemistry of the Eocene sands aquifer.

samples while gypsum and anhydrite are present at only one place in the extreme east of the basin. The presence of few minerals agrees with the fairly dilute water samples. Although mineralogical data are not sufficient for such a large area, they are unable to explain the widespread existence of sulphate in groundwaters.

In the water samples, the saturation indices have been calculated with the geochemical code PHREEQC (Parkhurst and Appelo, 1999). Taking into account the accuracy of the field measurements of pH and alkalinity and the uncertainties of other analytical data, Nordstrom and Ball (1989) or Busby et al. (1991) estimate that waters are in equilibrium with respect to calcite if the saturation index ranges from K0.1 to C0.1. For dolomite a saturation index ranging from K0.2 to C0.2 may indicate equilibrium. According to this, our results show that the waters are generally at equilibrium with respect to calcium carbonate mineral and slightly undersaturated com-pared to dolomite (Table 3). All the waters are

undersaturated with respect to sulphur bearing minerals (gypsum, anhydrite and pyrite) and near saturation with respect to siderite. All others minerals phases which could control NaCl in a sedimentary system are largely undersaturated.

According to these results, calcite, dolomite, gypsum, halite, siderite and pyrite are the mineral phases taken into account to represent our system. Water interactions with these phases, including redox reactions, can provide a basis for the reactions that may occur in the aquifer.

4.2. Infiltration waters

It is difficult to identify a composition that might be typical of the average infiltrating groundwater during the past 25,000 years. In the study area, the only one well (Fig. 2, Well 23) situated near the outcrops is influenced by anthropic activities (NOK3 close to 50 mg/L) and the water does not represent the original water. The simplest assumption to obtain


Fig. 4. Map of sulphur-34 values in dissolved sulphates (Andre et al., 2002).

the original infiltrating water composition is gener-ally to take the lowest observed concentration of each ion within all groundwater samples. This relies on the fact that most of the water-rock interactions processes tend to increase ion concentrations. So, the average concentrations for the infiltration water, is deduced from both these lowest observed concentrations and elemental concentrations in precipitation through historical times (Schoeller, 1956): water infiltration temperature of 10 8C, [NaC]Z[ClK]Z0.2 meq/L, [SO24K]Z0.04 meq/L, [Mg2C]Z0.4 meq/L.

The initial equilibrium pCO2 must be addressed with more details. In soils, two extreme conditions can prevail: calcite is either in contact with a fixed CO2 partial pressure, the open system, or it is dissolved by water containing a given amount of dissolved CO2, the closed system. Two arguments suggest that water infiltrating Eocene sands equilibrated with

a fixed CO2 pressure in an open system: (i) water infiltrates mostly in alluvial plains and mountain piedmont with colluvial soils often saturated with calcite and (ii) equilibrium with calcite at depth, i.e. with a finite source of CO2, would have led to alkalinity 510 times lower than the one observed in groundwater samples. Considering an open system, soil pCO2 largely varies from 10K1 to 10K3 atm in active to cold soils (Clark and Fritz, 1997); it is thus difficult to estimate an average equilibrium pCO2 for the infiltration water over geologic time.

Fig. 5 shows the theoretical evolution of pH versus temperature for an initial infiltration water having a temperature of 10 8C. First, this water reaches equilibrium with CaCO3, at 10 8C in an open system (for different initial pCO2), and then, the temperature increases from 10 to 60 8C and the water is maintained in equilibrium with CaCO3, in a closed system.


Table 3

Computed saturation indices of waters (PHREEQC)

WellsAragonite Calcite Dolomite Anhydrite Gypsum FluoriteHalite QuartzAmorphousSideritePCO2

numbersilica

1K0.41K0.27K0.96K2.23K2.07

2K0.43K0.30K1.00K2.34K2.18

3K0.38K0.24K0.90K2.14K1.94

4K0.24K0.10K0.29K1.44K1.31

5K0.040.09K0.08K2.48K2.38

6K0.15K0.01K0.50K2.72K2.53

7K0.050.08K0.23K2.79K2.72

8K0.18K0.04K0.22K2.52K2.39

9K0.060.06K0.21K2.94K2.89

100.080.210.40K0.88K0.70

11K0.16K0.02K0.13K2.07K1.86

12K0.27K0.12K0.47K2.26K2.05

13K0.20K0.05K0.37K3.18K2.94

14K0.16K0.02K0.66K2.62K2.42

15K0.090.05K0.56K2.73K2.54

16K0.18K0.04K0.73K2.77K2.59

17K0.27K0.15K0.34K2.18K2.17

18K0.33K0.18K0.27K2.79K2.56

19K0.22K0.08K0.39K3.35K3.13

20K0.28K0.14K0.84K2.88K2.66

210.020.140.29K2.53K2.46

220.270.420.86K2.92K2.68

23K0.010.14K0.84K2.53K2.28

24K0.36K0.22K0.58K1.20K0.96

K3.08K8.520.26K0.920.14K1.76

K3.02K8.570.27K0.920.25K1.73

K3.00K8.600.33K0.890.07K1.83

K0.69K5.44K0.09K1.26K0.03K1.56

K3.12K8.810.18K0.950.52K1.96

K3.00K8.750.37K0.840.58K2.18

K3.04K8.600.16K0.950.56K2.04

K2.72K8.370.19K0.960.29K2.12

K2.84K8.430.15K0.940.18K2.00

K0.64K6.470.29K0.920.32K2.12

K1.46K8.400.28K0.940.01K2.20

K1.97K8.460.32K0.930.25K2.22

K2.65K7.970.40K0.90K0.84K2.60

K3.23K8.950.27K0.95K0.65K2.06

K3.08K9.030.32K0.890.58K2.10

K3.25K9.130.31K0.900.52K2.00

K0.74K5.820.08K0.990.96K1.32

K0.69K6.460.12K1.140.25K2.33

K2.41K7.970.26K0.980.26K2.21

K2.52K8.680.31K0.94K4.88K2.14

K1.18K8.000.15K0.96K4.89K1.68

K1.51K8.280.15K1.141.07K3.31

K3.31K8.650.46K0.88K3.54K2.18

0.01K6.230.34K0.950.20K1.83

Plotting pH versus temperature for all the sampling points (Fig. 5) reveals that most of the points lies on the same theoretical line; they correspond to the same model and thus to the same original pCO2 of approximately 10K2.5 atm. This unique value of pCO2 will be considered as the one of the infiltrating water for the whole basin.

4.3. Elementary processes controlling water composition

4.3.1. CaHCO3 system

As described in part 3.2, CaHCO3 water is the predominant type in the Eocene sands aquifer, produced by dissolution of carbonate minerals such as calcite. The CO2 produced by the oxidation of organic matter and root respiration in the unsaturated zone and dissolved by the recharge water is at the origin of such dissolution. On Fig. 6, plotting pH measured in the field, versus pCO2, we can note a good correlation with the calcite saturation line. So, equilibrium with calcite and potential dissolution

of dolomite will be used in the geochemical modelling to set the CaMgCO3 system.

4.3.2. Sulphur case

The evolution of water composition towards CaSO4 type can be explained by using isotope

Fig. 5. Variation of pH versus temperature of emergence waters. Lines represent pH computed from evolution of infiltration water at 10 8C in equilibrium with calcite and with different initial pCO2 versus temperature and points represent waters of the Eocene sands aquifer.


Fig. 6. Variation of pH versus pCO2. Full line, pH evolution of a solution at equilibrium with calcite. Dashed line, pH evolution of a solution at equilibrium with dolomite.

values of sulphur. Data allowed us to distinguish two reactions:

Gypsum or anhydrite dissolution. Where it occurs, this process has been applied in a closed system consisting of a confined aquifer with calcite remaining at saturation. During the gypsum (or anhydrite) dissolution, delivering sulphate and

calcium to the water, calcite becomes over-saturated and, as it precipitates, the HCOK3 concentration decreases and the pH increases.

Pyrite oxidation. This process must be detailed in order to model it. Waters from the Eocene sands aquifer revealed small quantities of oxygen (00.15 mg/L) and redox potentials ranging between K50 mV and 0 mV, signifying slightly reduced media, and suggesting that pyrite oxi-dation is not associated with oxygen reduction. Among several other possibilities, the following reactions involving iron oxides and pyrite (Holmes

and Crundwell, 2000) were the only combination able to reproduce the concentrations of Fe, H2S, SO4 in solution

FeOH3 C3HC4Fe3C C3H2O

FeS2 C14Fe3C C8H2O415Fe2C C2SO24K C16HC

Fe2C CHCOK34HC CFeCO3

The overall reaction is written as

FeS2 C14FeOH3 C11HC C15HCOK342SO24K

C34H2O C15FeCO3

This reaction shows that carbonate equilibrium is modified by this process. Two arguments are in agreement with this process: in zone C, the d18O(SO4) values express an origin of dissolved sulphates from an oxidation of sulphurized compounds and a Eh-pH diagram of the sampled waters (Fig. 7) clearly indicates that this reaction should be the prominent one for these samples. In our approach, pyrite and siderite being at equilibrium, the reaction process is limited by the availability or dissolution rate of iron hydroxides. The dissolution of these solids is quite slow and they can thus remain in the solid phase while slowly dissolving. One may argue that the presence of pyrite and iron hydroxides in the same sediment can be surprising as one species occurs in reduced media while the second one is present in oxidised media. However, in sediments, reduced and oxidised periods can occur (Dubreuilh, 1987; Rey, 1995) and cuttings from wells drilled in zone C (Cazal et al., 1967; Rechiniac, 1962) report important quantities of


Fig. 7. Stability fieldsof oxides, carbonates and iron sulphur, at 25 8C, [H2S]Z[HSK]Z10K6 mol/L, [Fe3C]Z[Fe2C]Z10K5 mol/L, [Ctotal]Z3.10K3 mol/L, [SO24K]Z2!10K4 mol/L Diagram obtained with equilibrium constants from Garrels and Christ (1965). Points represent waters in which sulphates come from pyrite oxidation and gypsum dissolution and crosses symbolise waters in which sulphates come from gypsum dissolution exclusively.

pyrite in the aquifer as well as at the top and bottom of the unit. These investigators reported also the presence of mixtures of iron hydroxides. The choice of iron hydroxides instead of others iron oxides will be discuss below.

4.3.3. Cation exchange

This process explains the observed increase in NaC concentration without an associated increase in ClK concentration. Many studies have shown this kind of process (Back, 1966; Freeze and Cherry, 1979; Thorstenton et al., 1979; Chapelle and Knobel, 1983; Appelo and Postma, 1999). In an aquifer where carbonates minerals are present, cation exchange may be accompanied by calcite dissolution (and dolomite), as Ca2C (Mg2C) is removed from solution and replaced by NaC. The carbonate mineral

dissolves and provides more Ca2C (and Mg2C) in solution to exchange with NaC and causing HCOK3 concentration to increase. In the Eocene sands, the series sometimes include argillaceous or argillaceous-sandstones interbeds, able to play a role as ion exchangers. Depending on location, these layers show variable thickness. Under these conditions, ion exchangers will be considered in the modelling approach and will be used to explain Na-HCO3 type of some waters.

Under the chemical conditions considered here, excluding redox reactions in a first approach, seven conditions are imposed: equilibrium with calcite, equilibrium or addition of dolomite, addition of gypsum and halite, ion exchange, starting pCO2 of infiltration water and electro neutrality. Seven ion concentrations are considered, i.e. Ca, Mg, Na, CO3, SO4, Cl and H. We are therefore in a fully determined system which should have a unique solution, provided the data are of good quality. Although the modelling system by itself might be consistent, the assumptions might be wrong; for instance we might be able to solve a system by adding a small amount of dolomite, while Mg may come from another source. To overcome this difficulty it is important to have validation variables that may confirm or reject our hypotheses.

4.4. Model variables validation

The pCO2 is different from the other variables as an initial value is specified over the whole area. In fact all other processes or added minerals are adjusted at each point to reproduce the local concentration of the ions of interest, as it is done in any modelling study (Plummer and Sprinckle, 2001; Guler and Thyne, 2004). As all the processes may change carbonates equilibrium, the value of pCO2 may vary along the water pathway and can be used to validate the differences between the points. But this value is difficult to measure. So, in this condition, we decide to use an other variable which depends of all processes: the pH. Moreover, the attention brought on the measure of pH in the field allows us to consider this variable as a good validation variable.

d13C can also be used as a validation variable because its relative concentration in groundwater will depend only on the geochemical processes cited above (Clark and Fritz, 1997). Thus if the calculated amounts


of calcite and dolomite dissolved are the right ones, the 13C values should match. The underlying assump-tion is that 13C values come from the two end members: calcite and soil water. Assuming here the classical literature values, i.e. 0 for 13CO3 in calcite and dolomite and K23 for soil 13CO2 gas, we consider that all the waters infiltrate under the same conditions (soil pCO2Z10K2.5 atm) and that calcite dissolution at equilibrium with CO2 of the soil gives a value of d13C of the dissolved CO2 equal or near to K12 (Clark and Fritz, 1997). The further evolution of this value depends only on the different processes (dissolution or precipitation) affecting carbonate minerals during the groundwater flow in a closed system. The precipitating calcite has a d13C value starting from the isotopic equilibrium with the groundwater and slightly enriched by C2 (Clark and Fritz, 1997); the same hypothesis will be used for siderite.

4.5. Comparison of modelling approaches

All the chemical processes detailed above are well known and are used in studies in literature (Murray, 1996; Rosen and Jones, 1998; Edmunds et al., 2000; Marini et al., 2000). As the chemical composition of groundwater in such systems mainly results from salt dissolution and carbonate equilibrium, it is possible to fit most of the groundwater composition by adjusting the dissolved amount of the various minerals. This approach generally gives results in good agreement with observed data for major chemical parameters. Modelling results are thus used to suggest hypotheses on the water pathway within the aquifer. Although other geochemical processes occurring within the aquifer (exchanges with aquitards or mixing with others aquifers) are assumed to explain anomalous chemical composition, these hypothesis are often not confirmed by additional data. A more detailed approach, incoming recently, implies the use of inverse models including mixing, dissolution/precipi-tation salts, cation exchange that may explain changes of groundwater chemistry between points lying on the same flow line. These models often concerned major ions present in solution (Gosselin et al., 2001; Sracek and Hirata, 2002; Martinez and Bocanegra, 2002). Inverse modelling, although including more variables, can still consist to fit the variable amounts of minerals added to the solution, if validation variables are not

used. A typical adjustment can consist in using a variable equilibrium pCO2 for each modelled point. Other authors validate their models using additional data for model calibration. They use isotope data, like 13C or 34S, or elements in solution like total dissolved carbon (Walvoord et al., 1999; Plummer and Sprinckle, 2001; Guler and Thyne, 2004). Our model was inspired from this approach but, in our case, the system is much more constrained. The origin of sulphate by the pyrite dissolution may be used to illustrate this purpose: the reactions must satisfy conditions on the concentrations of sulphur and iron, the redox potential, the availability of O2, sulphur isotope values and finally pH. All these parameters could not be matched by simple dissolution of pyrite; the complex reaction scheme, involving two dis-solving and one precipitating solid, seems to be the only one to fulfil all conditions. The few available informations on the sediment solid phase tend to confirm our hypotheses. We also constrain the carbonates equilibrium by setting a unique pCO2 for all infiltrating waters. An inverse modelling of this value would have led to an easy fit of pH. But, in turn, this might have hidden the effects of redox reactions on pH or other origins of CO2. In addition, being a good tool for validation, d13C also allowed to identify potential sources of additional CO2. In that case again, several hypothesis were tested and, for the points where anomalous CO2 occur, only a direct addition of CO2 satisfies carbonate equilibrium, pH and d13C.

During water-rock interactions numerous equili-bria are involved but they can be ordered according to the dependency among the different reactions, in the considered case it is: redox reaction, gypsum dissolution, carbonate equilibrium, halite dissolution. At each step, a complete validation of all parameters influenced by the geochemical processes is done with a justification by the composition of the solid phase. The major objective is to have as less as possible variable phase amounts.

5. Modelling results

All groundwater samples were modelled using the approach described above (Fig. 8). It was possible to reproduce most of the water compositions (Tables 46).


Fig. 8. Geochemical processes governing the chemical composition of waters from the Eocene sands aquifer.

Table 4

Comparison between measured and modelled chemical composition in south-western zone

GeauneLussagnetWell 13Barbotan

Initial PCO2 (atm)10K2.510K2.510K2.510K2.5

NaX5!10K25!10K21.74!10K1

Halite////

Dolomite////

GypsumK6!10K2K1!10K1K6!10K2/

PyriteK2.86!10K2K4.29!10K2K2.86!10K2K1.43!10K1

Fe(OH)3K4!10K1K6!10K1K4!10K1K1

CalciteC2.73!10K1C3.92!10K1C1.77!10K1C4.36!10K1

SideriteC4.23!10K1C6.38!10K1C4.25!10K1C1.14

O2(g)///2.5!10K1

MeasuredComputedMeasuredComputedMeasuredComputedMeasuredComputed

valuesvaluesvaluesvaluesvaluesvaluesvaluesvalues

PH7.357.547.377.637.868.127.107.85

HCO3K2.992.252.841.963.212.432.771.47

Ca2C1.241.011.090.930.520.380.870.63

ClK0.160.200.200.200.240.200.250.20

KC0.100.000.140.000.170.140.160.04

Mg2C0.180.190.260.190.260.090.280.16

NaC0.240.250.260.251.851.830.570.60

SO42K0.140.120.190.180.120.130.310.27

Altotal! LD8.3!10K52.8!10K41!10K4! LD7.1!10K5! LD2.7!10K4

Sitotal0.250.130.300.130.220.080.290.15

Log(PCO2)K2.03K2.35K2.08K2.50K2.60K2.99K1.80K2.81

13C ( PDB)K11.4K12.5K11.2K13.0K10.05K12.5K12.0K14.0

LD, limit of detection (C); mineral precipitation from solution; (K), mineral dissolution from solid to solution. All data for minerals or exchangers quoted in the upper part of the table indicate the amounts of these minerals dissolved/precipitated in solution to reach the measured concentration. They are expressed as mmol/L of solution. In the low part of table, pH are expressed in standard units whereas measured and computed values are expressed in mmol/L. Some wells are very near and the chemical compositions of waters are similar. In these conditions, Geaune represents waters from wells 1416, Lussagnet represents waters from wells 5 and 6 and Barbotan represent waters from wells 13.

Table 5

Comparison between measured and modelled chemical composition in south-eastern zone

Well 9Well 11Well 7Well 12Well 8Well 20Well 19

Initial P(atm)10K2.510K2.510K2.510K2.510K2.510K2.510K2.5

CO2

NaX5.5!10K17.5!10K14.50!10K15.50!10K18!10K12.5!10K16!10K1

Halite//////K2.25!10K1

Dolomite/K6!10K1/K3!10K1K2!10K1/K2!10K1

GypsumK8!10K2K1K9!10K2K5.5!10K1K3!10K1K1!10K1K3!10K2

Pyrite/K2.8!10K2K7!10K3K2.14!10K2/K1!10K2/

Fe(OH)3/K4!10K1K1!10K1K3!10K1/K1.4!10K1/

CalciteC1.10!10K1C1.44C1.55!10K1C8.14!10K1C4.12!10K1C1.12!10K1C3.82!10K1

Siderite/C4.23!10K1C1.01!10K1C3.15!10K1/C1.44!10K1/

MeasuredCompu-MeasuredCompu-MeasuredCompu-MeasuredCompu-MeasuredCompu-MeasuredCompu-MeasuredCompu-

valuested valuesvaluested valuesvaluested valuesvaluested valuesvaluested valuesvaluested valuesvaluested values

Ph7.337.237.487.617.407.287.457.597.507.487.397.517.507.58

HCOK2.622.672.932.282.562.562.782.432.672.862.802.642.992.92

3

Ca2C0.800.951.110.940.880.961.001.030.780.961.051.110.750.85

ClK0.220.200.200.200.130.200.210.200.200.200.210.200.430.43

KC0.150.050.180.070.140.040.210.050.150.080.070.020.090.06

Mg2C0.190.160.600.600.180.170.430.400.310.300.210.180.320.32

NaC0.740.740.920.940.580.640.750.741.000.990.430.451.021.01

SO2K0.080.080.830.820.100.110.540.520.290.280.110.120.040.04

4

Altotal9.6281.19!LD2.41.300.577.47.31.24.77.45.5

!10K5!10K5!10K3!10K5!10K4!10K4!10K4!10K5!10K5!10K5!10K5!10K5!10K5

Sitotal0.360.250.250.130.340.230.240.110.290.130.230.110.210.11

Log(PC-K2.00K1.81K2.21K2.43K2.05K1.90K2.22K2.41K2.13K2.19K2.15K2.28K2.22K2.31

O2)

13C (K14.4K12.1K12.6K10.2K13.5K12.2K13.8K11.0K12.9K10.9K14.3K12.2K14.5K11.1

PDB)

(C): mineral precipitation from solution; (K): mineral dissolution from solid to solution. pH are expressed in standard units whereas amounts of minerals (upper part of table) and measured and computed values (low part of table) are expressed in mmol/L.54



Table 6

Comparison between measured and modelled chemical composition in northern zone

Well 17Well 4Well 24Well 18Well 10

Initial PCO2 (atm)10K2.510K2.510K2.510K2.510K2.5

NaX112015514

HaliteK3.4K8.9K4.4K2K1.2

DolomiteK9!10K1K2.62K2.5K5.6!10K1K5.7

GypsumK1.5K7K6K6.5!10K1K15

CalciteK2.11!10K1C3.6C4.77C9.05!10K1C11.27

CO2(g)2.225!10K1/5!10K1

MeasuredComputedMeasuredComputedMeasuredComputedMeasuredComputedMeasuredComputed

valuesvaluesvaluesvaluesvaluesvaluesvaluesvaluesvaluesvalues

pH7.207.477.207.257.087.177.778.507.447.06

HCO3K6.986.836.036.033.323.264.613.223.373.08

Ca2C0.310.301.120.943.13.450.310.113.363.72

ClK3.563.609.059.104.724.602.262.201.401.40

KC0.260.530.520.960.380.140.140.280.460.96

Mg2C0.130.120.560.511.891.900.280.082.102.19

NaC12.3211.9423.7323.676.145.967.156.2113.6813.86

SO42K1.371.396.796.004.414.430.610.6411.1310.94

Altotal2.2!10K44!10K45.6!10K41.0!10K42.6!10K4 1.3!10K54.30!10K4 4!10K4! LD3.8!10K5

Sitotal0.30.230.140.170.200.090.140.110.290.14

Log(PC-K1.32K1.7K1.56K1.62K1.84K1.93K2.33K3.20K2.12K1.7

O2)

13C (K8.41K6.6K9.98K5.7K4.15K7.1K11.8K9.4K6.81K8.69

PDB)

(C): mineral precipitation from solution; (K): mineral dissolution from solid to solution. pH are expressed in standard units whereas amounts of minerals (upper part of table) and measured and computed values (low part of table) are expressed in mmol/L.

Among the processes, calcite equilibrium is the major factor influencing pH evolution, the pH value being also influenced by pyrite oxidation where it occurs. A good agreement is found between calcu-lated and measured in situ pH, the maximum variations reaching only 0.3 pH unit except for two points having variation near 0.7 pH unit (Fig. 9). This result is very convincing because the implied processes are numerous and the pH is a quite sensitive variable. Moreover, the range of variation (between 7 and 8.3) is quite wide. Contrary to other regions, the western zone is the only one where the modelled pHs are approximately 0.2 pH units higher than the analytical data. All proportions of sulphate originating from pyrite dissolution were calculated using the only available value, d34SpyriteZK20, obtained from literature (pyrite samples were too small to measure 34S). A change in this isotopic composition of pyrite modifies significantly the amount of dissolved pyrite

Fig. 9. Comparison between pH measured and computed. Error bars on measured pH correspond to measurements errors (5%). Computed pH have been obtained from three different data bases (WATEQ, MINTEQ and PHREEQC).


which may, in turn, modify the pH value by 0.2 to 0.3 units. This may be the source of the observed small discrepancy.

Measured and computed d13C values are similar (Tables 46). The variations do not exceed 4. Owing to the uncertainty in values presented in literature (calcite and dolomite: 0G2, CO2 gas in soil: C23 G2) (Clark and Fritz, 1997), the results can be considered to agree well. More precision could be obtained in future studies if d13C of the calcite from the aquifer would be measured.

The overall dissolved-solids concentrations can be obtained mainly from gypsum dissolution and slight halite (NaCl) dissolution at some locations. Although this process is in agreement with the water compo-sition, waters are largely undersaturated with respect to both minerals. This means that gypsum and halite minerals are not present inside the active part of aquifer, a conclusion consistent with the deposit age of the aquifer and its fairly fast renewal. Then another process may deliver to the aquifer water that has previously dissolved gypsum (gypsum dissolution is needed to reach the d34S values of sulphate existing in groundwater). The source of gypsum must then be search in the aquitard. Throughout the aquifer small amounts of dissolved dolomite were needed to model the measured concentrations of Mg. Again, most of the waters being undersaturated with respect to dolomite, the Mg rich waters may come from vertical leakage and not from the local solid phase. However, for Mg content, calcite can also play an important role as non-negligible amounts of Mg are often present in calcite. With the existing data, we cannot answer this question.

As expected, the variations in Na concentrations are influenced by the ion exchange process. Despite our poor knowledge of the argillaceous phase, the typical ion exchange constant used seems to give results in agreement with measured concen-trations (Tables 46). Ion exchange occurs at only few places, but significantly modifies the groundwater composition at such places.

As said above, the measured alkalinity could not be modelled at a few wells (wells 4, 10, 17). Among other possibilities the only process that has been able to reproduce the measured data is the addition of CO2. For these wells d13C was not used as a validation

variable but as an understanding tool. In fact the d13C

values of these waters were only compatible with a deep origin of CO2 with a mean d13C of K5

(Clark and Fritz, 1997), which could result from metamorphic gas production. This process has been observed elsewhere (Chiodini et al., 1999) and may be explained by the presence of important structures facilitating upward gas flow. It is interesting to notice that at two sites (wells 4 and 10) anomalous radon content has also been observed (Franceschi, 2005). As radon is known to come from deep granitic bedrock, a deep origin of CO2 seems to be plausible.

6. Geochemical arguments for flow delineation

The developed model can also help to understand the connections between different sampled points. In fact a similar composition of two waters is not a proof of hydrodynamic connection. However, if an hydraulic connection is assumed, one must be able to explain the chemical processes that lead from one point to the other. This is the basis of how we used the geochemical model to delineate groundwater flow paths: some water paths are not possible and some preferential flow directions can be drawn if one compound is increasing along the flow line. A flow barrier constraint is present at several places; it is not possible to go from Demu (well 9) to Nogaro 2 (well 7), nor from Beaucaire (well 10) to Gondrin (well 8). Furthermore, the major concentrations trends are from east to west in the northern area and from south to north in the southern area. Fig. 10 depicts the major flow directions derived from the modelling results; the four major flow lines with the associated chemical composition evolution and models are given in Tables 46.

6.1. Major flow directions

One main conclusion from the reconstruction of the groundwater composition by modelling is that only two major zones exist in the studied area: the south and the north. Within each zone the differences in groundwater composition arise from the local intensity of the geochemical processes. These differences may be used to assess the major flow directions. The main assumption underlying this approach is that, except


Fig. 10. Major flow direction derived from modelling results.

where mixing of large volumes occurs, or for the specific case of ion exchange, the concentrations of dissolved species may not decrease along the direction groundwater flow. In the northern part of the aquifer there is a clear trend of salt content from east to west along the line Lalbare`de (well 18) Blagnac (well 17)Beaucaire (well 10). The main flow directions are in agreement with the main gradient directions. From Beaucaire and in the western part, the flow pattern is not clear, both head and salinity showing a strong decrease in the west of Beaucaire. The concentration decrease might be explained by flow of dilute waters from the north, while the higher headgradientmay indicate a lower permeability and

a limitedflow from east to west at this place.

We presently do not have enough data to confirm or refute this hypothesis.

In the southern area, the potentiometric map tends to show a major flow direction from south-east to north-west. Most of the waters are very dilute and mainly of bicarbonate type. In the eastern area some trace elements such as fluoride, strontium or lithium are more concentrated. On the other hand isotopic studies showed that in the western part a significant portion of the sulphates arise from pyrite dissolution. Water composition shows large bands of similar composition oriented in the southnorth direction. Modelling results give several constraints on the flow directions; it is not possible to obtain Nogaro 2 water (well 7) from evolution of the water at Demu (well 9). For this reason, we propose that a geochemical watershed limit exists between the two regions and a flow towards the north occurs in this area (Fig. 10). In the western part, the evolution of 34S isotopes


suggests a trend of pyrite dissolution oriented in the same southnorth direction. All geochemical argu-ments thus converge to indicate a south-to-north flow direction.

6.2. Local flow patterns

6.2.1. Closed structures

Around several structures two properties are observed: (i) the estimated age of the waters are older than in the surrounding aquifer, and (ii) the groundwaters are more mineralized and marked by ion exchange. These phenomena occur around the Garlin and Eugenie-les-Bains structures (Fig. 2). The waters sampled close to the Garlin dome are among the oldest water with an estimated age of 35,000 years while the waters from the very close Lespielle borehole (well 19) are 16,000 years old. Furthermore, the modelling step showed that half of the sodium comes from ion exchange at this location. The same phenomenon occurs at well 13 where waters are considerably older (28,000 year) than the nearby well of Geaune (well 14) (19,000 year). At both sites the importance of ion exchange can be related to the presence of thick layers of interbeded clays within the aquifer. The presence of this non-conductive layer may be the origin of the slow flow and thus older water.

6.2.2. Open structures

In contrast to the previous structures, the waters sampled around open structures show only slight variations in the groundwater composition. At several wells the waters seem to be younger, owing to their 14C estimated age. It appears that these younger waters may arise from mixing with superficial aquifers at places where waters from the Eocene sands aquifer discharge. This process is found at Barbotan (Wells 1 and 2) and Castera-Verduzan (Well 11). At Barbotan, another argument may confirm this process. At that place we showed that all sulphates come from pyrite oxidation. The use of the oxygen isotopes of the sulphate molecule also revealed that this oxidation may partly come from oxygen (Andre et al., 2002). The presence of oxygen in the pyrite oxidation process occurring only at that location might be explained by mixing with oxygen rich superficial waters.

6.2.3. More complex structures

The model developed suggests that the waters from the southern part may discharge to the north. This was suggested particularly at Demu (well 9) where the geochemical arguments preclude flow to the west (i.e. to well 7). The Celt-aquitaine fault zone is not continuous and there might be a potential outflow between faults (Fig. 11). It is quite interesting to notice that the waters sampled at well 8, in the northern area, could be obtained by mixing equal proportions of well 9 and well 12 waters. It seems thus plausible to have global outflow to the north at this place resulting in closer flowlines. This would also explain the lower salt content of well 8 waters compared to points located at the east (wells 4 and 10) and the surprisingly steep head gradients in this area.

6.3. Leakage

Leakage is used with increasing frequency in regional modelling of multi-layer aquifers (Toth, 1999). The presence of such vertical fluxes is fairly important for the management of aquifer pumping at the regional scale.

Eocene sands contain very few minerals and groundwaters are very dilute, in the southern part. In this context, the presence of calcium carbonate waters is expected but the presence of sulphate is difficult to explain except if mass transfers from aquitards are considered.

Several processes can explain this vertical sulphate transfer, the major ones being leakage and diffusion. Leakage consists in a vertical flow through an aquitard, in this case from the molasses, to the Eocene sands. As interstitial molasse waters may be saturated with respect to gypsum, i.e. 100 times more concentrated than the aquifer waters, a small vertical water flux can modify the composition of the groundwater in the Eocene sands. On the other hand, sulphate can simply move by diffusion from a concentrated medium (the molasse) to a dilute one (the Eocene sands).

A detailed theory of the different possibilities of vertical exchanges between aquifers and aquitard has been developed elsewhere (Atteia et al., 2005). The authors showed that (i) diffusion might explain low concentrations of some ions and a slight increase along flow of ions present in aquitard pore waters and (ii) leakage could generate fairly important


Fig. 11. Potential flux through the CeltAquitaine flexure.

concentrations of the same ions where the difference between the land surface and the potentiometric surface of the Eocene sands aquifer is quite sufficient to generate leakage (according to Darcys law). For our case, the diffusion of sulphate from the molasse gypseous layers could explain the observed concen-trations and sulphur isotopic content of the aquifer waters in the southern zone. At contrary, the high concentrations found in the north could only be justified by leakage.

This difference between the south and north of the basin can be explained by geological conditions. First, in the south, the molasse is much thicker than in the north, and second, the molasse lithology might be coarser in the north than in the south. The differences in molasse lithologies could be explained by the sedimentation history in both geological compart-ments (Rey, 1995).

7. Conclusion

The results of this study show that detailed hydrochemical data coupled with geochemical mod-elling can help to elucidate the hydrologic and geologic factors controlling water chemistry on a regional basin.

Waters from the Eocene sands aquifer present mainly calcium bicarbonate facies with an evolution of the geochemical facies at several places. To understand these variations, detailed chemical and isotopic data were used; changes in water chemistry were interpreted and three distinct hydrochemical processes have been identified to be responsible of these evolutions: (1) Waters equilibrium with calcite from recharge areas to discharge places; (2) Cation exchange between waters and the aquifer material; and (3) Pyrite oxidation and gypsum dissolution that


can explain the regular increase of sulphur concen-tration along pathways.

These processes identified, a modelling approach allowed us to quantify those. This step by step modelling used a limited number of mineral phases that were identified in the solid and successfully reproduces the concentration of major ions at each sampling points. Numerous constraints that were imposed on the model and its validation by pH or 13C allowed us to justify and quantify the geochemical processes occurring in solution.

Afterwards, based on the local reconstitution of chemical composition of waters, geochemical model-ling was used to delineate geochemical pathways. The origin of local chemical composition variations has been identified and the role played by deep geological structure has been outlined. It appears that a concept of open and closed structures seems to be applicable in this region. This concept would enlarge the classical discussion of faults as barriers or conducts for flow.

The study also shows that the chemical compo-sition of waters cannot be explain without evoking the influence of aquitards. Considering the poor miner-alogy of the sands, the increase of total dissolved solids along pathways seems to be associated to mass transfers from the gypseous molasses, present at the top of the aquifer. This transfer of ions occurs according two processes: leakage, principally in the northern part of the aquifer, and diffusion, mainly in the south of the study area.

In the context of the Aquitaine Basin, the geochemical modelling has proven its ability to constrain efficiently the groundwater flow in an aquifer were few data are available. However, the developed approach shows that a detailed analysis of each reaction, the use of validation variables and information on mineral phases present in the solid are required. At contrary, the concentration increase for one dissolved species can be attributed to the wrong reaction. Along the paper, we also showed that a geochemical model was able to provide four types of hydrodynamic information: estimates of regional flow directions, existing flow barriers, mixing between aquifer waters and interactions between aquifer and aquitard. However, this information remains qualitat-ive, the only solution to reach more quantitative flow patterns is the use of 3D multilayer models

including hydrodynamics, transport and chemistry. The chemistry may constrain flow but the opposite is also true: the flow pattern governs mixing and residence time.

Acknowledgements

The authors would like to thank TotalFinaElf Stockage Gaz France for supporting this work.

References

Andre L., 2002. Contribution de la geochimie a` la connaissance des ecoulements souterrains profonds. Application a` laquife`re des SablesInfra-Molassiques du BassinAquitain,The`se:Bordeaux3. p. 230.

Andre, L., Franceschi, M., Pouchan, P., Atteia, O., 2002. Origines et evolution du soufre au sein de laquife`re des Sables Infra-Molassiques du Bassin Aquitain. C.R. Geosciences 334, 749756.

Appelo C.A.J., Postma D., 1999. Geochemistry, Groundwater and Pollution, A.A. Balkema/Rotterdam/Brookfield/1999, p. 536.

Atteia O., Andre L., Franceschi M., Dupuy A., 2005. Contributions of diffusion, dissolution, ion exchange and leakage from low permeability layers to confined aquifers (Accepted by Water Resour. Res.).

Back W., 1966. Hydrochemical facies and groundwater flow patterns in northern part of Atlantic Coastal Plain. US Geological Survey Professional Paper 498-A, p. 42.

Blavoux, B., Dray, M., Fehri, A., Olive, P., Groning, M., et al., 1993. Paleoclimatic and hydrodynamic approach to the Aquitaine Basin deep aquifer (France) by means of environmental isotopes and noble gases. Int. Symp. App. Isotope Techn. 1993, 293305.

Boule`gue, J., Popoff, G., 1979. Nouvelles methodes de determi-nation des principales espe`ces ioniques du soufre dans les eaux naturelles. J. Fr. Hydrol. 10 (2) 8390 n 29.

Busby J.F., Plummer L.N., Lee R.W., Hanshaw B.B., 1991. Geochemical evolution of water in the Madison aquifer in parts of Montana, South Dakota, and Wyoming. US Geological Survey Professional Paper 1273-F, p. 89.

Cazal, A., Moussie, B., Feneyrou, G., 1967. Etude de la circulation dune eau minerale a` travers un aquife`re complexe, Chronique dhydrogeologie. BRGM 12, 91112.

Chapelle, F.H., Knobel, L.L., 1983. Aqueous geochemistry and the exchangeable cation composition of glauconite in the Aquia aquifer, Maryland. Ground Water 21 (3), 343352.

Chiodini, G., Frondini, F., Kerrick, D.M., Rogie, J., Parello, F., 1999. Quantification of deep CO2 fluxes from Central Itaky. Examples of carbon balance for regional aquifers and of soil diffuse degassing. Chem. Geol. 159, 205222.


Clark, I., Fritz, P., 1997. Environmental Isotopes in Hydrogeology. Lewis Publishers, New York, p. 328.

Claypool, G.E., Holser, W.T., Kaplan, I.R., Zak, I., 1980. The age curves of sulphur and oxygen isotopes in marine sulphate and their mutual interpretation. Chem. Geol. 28, 199260.

Dazy, J., Rochat, J., Olive, Ph., 1980. Nouvelles donnees geochimiques et isotopiques sur les eaux thermales dUriage-les-Bains (Ise`re), 105e`me Congre`s National des Societes Savantes, Caen. Sciences Fasc. II, 111122.

Dubreuilh, J. 1987. Synthe`se paleogeographique et structurale des depots fluviatiles tertiaires du nord du Bassin Aquitain. Passage aux formations palustres, lacustres et marines, The`se de Doctorat detat, Bordeaux III.

Eberts S.M., George L.L., 2000. Regional groundwater flow and geochemistry in the Midwestern basins and Arches aquifer system in parts of Indiana, Ohio, Michigan, and IL. US Geological Survey Professional Paper 1423-C, p. 103.

Edmunds, W.M., Carrillo-Rivera, J.J., Cardona, A., 2000. Geo-chemical evolution of groundwater beneath Mexico City. J. Hydrol. 258, 124.

Franceschi M., 2005. Radon 222 activities in a deep sandy aquifer and geological factor controlling its occurrence (Accepted by Ground Water).

Freeze, R.A., Cherry, J.A., 1979. Groundwater. Englewood Cliffs, Prentice-Hall, NJ, p. 604.

Garrels, R.M., Christ, C.L., 1965. Solutions, Minerals, and Equilibria. Harper and Row, New York. Appendix 2, p. 450.

Gerla, P.J., 1992. Pathline and geochemical evolution of ground-water in a regional discharge area, Red River Valley, North Dakota. Ground Water 30, 743754.

Gosselin, D.C., Harvey, F.E., Frost, C.D., 2001. Geochemical evolution of groundwater in the great plains (Dakota) aquifer of Nebraska: implications for the management of a regional aquifer system. Ground Water 39 (1), 98108.

Guler, C., Thyne, G.D., 2004. Hydrologic and geologic factors controlling surface and groundwater chemistry in Indian Wells-Owens Valley area, southeastern California, USA. J. Hydrol. 285, 177198.

Hendry, M.J., Schwartz, F.W., 1990. The chemical evolution of groundwater in the Milk River Aquifer, Canada. Ground Water 28, 253261.

Holmes, P.R., Crundwell, F.K., 2000. The kinetics of the oxidation of pyrite by ferric ions and dissolved oxygen: an electrochemical study. Geochim. Cosmochim. Acta 64 (2), 263274.

Institut EGID, 1999, Actualisation des cartes piezometriques regionales du bassin hydrogeologique de lAdour, 1998, p. 12, 5 pl.

Labat N., 1998, Role de particularites sedimentaires et structurales sur le comportement des sables sous-molassiques soumis aux fluctuations induites par les stockages souterrains de gaz. Application a` letude de leur influence sur lhydrodynamisme des emergences locales, The`se: Bordeaux 3. p. 228.

Marini, M., Ottonello, G., Canepa, M., Cipolli, F., 2000. Water-Rock interaction in the Bisagno valley (Genoa,

Italy): application of an inverse approach to model spring water chemistry. Geochim. Cosmochim. Acta. 64 (15), 26172635.

Martinez, D.E., Bocanegra, E.M., 2002. Hydrogeochemistry and cation-exchange processes in the coastal aquifer of Mar del Plata, Argentina. Hydrogeol. J. 10, 393408.

Murray, K.S., 1996. Hydrology and geochemistry of thermal waters in the Upper Napa Valley, California. Ground Water 34 (6), 11151124.

Nordstrom, D.K., Ball, J.W., 1989. Mineral saturation states in natural waters and their sensitivity to thermodynamic and analytic errors. Sci. Geol. Bull. 42 (4), 269280.

Ohayon-Courte`s C., 1992. Levolution des espe`ces reduites du soufre dans les eaux minerales, The`se: Bordeaux II, p. 319.

Parkhurst D.L., Appelo C.A.J., 1999. A computer program for speciation, batch-reaction, one dimensional transport and inverse geochemical calculations. US Geol. Surv. Water Resour. Invest. Rep. 99-4259, p. 312.

Pearson, F.J., Rightmire, C.T., 1980. Sulphur and oxygen isotopes in aqueous sulphur compounds, in: Fritz, P., Fontes, J. Ch. (Eds.), Handbook of Environmental Isotopes. Geochemistry. Elsevier, Amsterdam, pp. 227258. Chapter 6.

Pelissier-Hermitte, G., Franceschi, M., Cheron, J., Dupuy, A. 2000. Contamination azotee dune nappe captive par drainance dune nappe libre-Mecanismes et evolution. Colloque international ESRA-1315 septembre. S2-11-S2-14

Plummer, L.N., Sprinckle, C.L., 2001. Radiocarbon dating of dissolved inorganic carbon in groundwater from confined parts of the Upper Floridan Aquifer, Florida, USA. Hydrogeol. J. 9, 127150.

Rechiniac A., 1962. Etude sedimentologique des principales formations detritiques du Paleoge`ne aquitain, The`se: Bordeaux, p. 69.

Rey, J. 1995. Le Bassin dAquitaine: composition, evolution et structure. In: Du Lias Nord aquitain aux molasses mioce`nes Generalites Livret Guide dexcursion, STRATA, Actes du laboratoire de Geologie Sedimentaire et Paleontologie de lUniversite Paul SabatierToulouse, Serie. 2 Memoire, pp. 7140

Rosen, M., Jones, S., 1998. Controls of the chemical composition of groundwater from alluvial aquifers in the Wanaka and Wakatipu basins, Central Otago, New Zealand. Hydrogeol. J. 6, 264281.

Schoeller, H., 1956. Geochimie des eaux souterraines. Masson, Paris, p. 581.

Serrano, O. 2001. Le Cretace Superieur/Paleoge`ne du bassin davant-pays Nord-Pyreneen (Bassin dAquitaine): sedimento-logie, stratigraphie et evolution geodynamique, The`se: Rennes.

Sracek, O., Hirata, R., 2002. Geochemical and stable isotopic evolution of the Guarani Aquifer System in the state of Sao Paulo, Brazil. Hydrogeol. J. 10 (6), 643655.

Thorstenton, D.C., Fisher, D.W., Croft, M.G., 1979. The geochem-istry of the Fox Hills-Basal Hell Creek aquifer in southwestern North Dakota and northwestern South Dakota. Water Resour. Res. 15 (6), 14791498.

Toth, J., 1999. Groundwater as a geologic agent: an overview of the causes, processes and manifestations. Hydrogeol. J. 7, 114.


Trembaczowski, A., Halas, S. 1984. Oxygen and sulphur isotope ratios in sulphate from atmospheric precipitations. In: Proc. Int. Symp. Isotope Hydrol. Water Resour. Dev., I.A.E.A., Vienna.

819820

Walvoord,M.A.,Pegram, P., Phillips, F.M., Person, M.,

Kieft,T.L.,Fredrickson,J.K.,McKinley,J.P.,

Swenson, J.B.,1999. Groundwater flowand geochemistry

in the southeastern San Juan Basin: implications for microbial transport and activity. Water Resour. Res. 35 (5), 14091424.

Weaver, T.R., Bahr, J.M., 1991. Geochemical evolution in the Cambrian-Ordovician Sandstome aquifer, Eastern Wisconsin. 2. Correlation between flow paths and groundwater chemistry. Ground Water 29, 510515.

Andre Et Al 2005

Documents

Transcript of Andre Et Al 2005