Applied Geochemistry 26 (2011) 818–827

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Applied Geochemistry

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Water circulation control on carbonate-d18O records in a low permeability clayformation and surrounding limestones: The Upper Dogger–Oxfordiansequence from the eastern Paris basin, France

Véronique Lavastre a,b,c,⇑, Magali Ader c, Stéphane Buschaert d, Eddy Petit c,1, Marc Javoy c

a Université de Lyon, Université Jean Monnet, F-42023 Saint Etienne, Franceb CNRS, UMR 6524, LMV, F-42023 Saint Etienne, Francec Laboratoire de Géochimie des Isotopes Stables, Institut de Physique du Globe de Paris & Université Paris 7 – UMR CNRS 7154, 4, place Jussieu, 75252 Paris cedex 05, Franced Andra, Parc de la Croix Blanche, 7-8 rue Jean Monnet, 92 298 Châtenay-Malabry Cedex, France

a r t i c l e i n f o a b s t r a c t

Article history:Received 7 July 2010Accepted 3 February 2011Available online 5 March 2011Editorial handling by D.C. Gooddy

0883-2927/$ - see front matter � 2011 Elsevier Ltd. Adoi:10.1016/j.apgeochem.2011.02.003

⇑ Corresponding author at: Université de Lyon, UniSaint Etienne, France. Tel.: +33 (0) 477 481 585; fax:

E-mail address: veronique.lavastre@univ-st-etienn1 Present address: Université Montpellier II, IEM U

Bataillon, 34095 Montpellier Cedex 05, France.

Upper Dogger to Oxfordian Formations in the eastern part of the Paris basin (France) are currently beinginvestigated by the French nuclear waste management agency (Andra), testing the feasibility of long-term deep nuclear waste disposal in the Callovo-Oxfordian claystones. Characterising the hydrogeologicalbehaviour of the Callovo-Oxfordian claystones is, therefore, essential in evaluating its potential as a geo-logical barrier. In order to evaluate and quantify water/rock interactions experienced over geological timeby these Formations, bulk carbonate d13C and d18O were measured and calculations of water–rock ratioswere used to explain carbonate-d18O changes. Meteoric porewater and a maximum temperature reachedof about 40 �C were considered. The Jurassic marine carbonate d13C was preserved in the Callovo-Oxfor-dian claystones and in the overlying limestones (�0.28‰ to 3.39‰/PDB), while the d18O values are lowerby 0–5‰ (�6.25‰ to �1.32‰/PDB). Calculations show that Upper Dogger and Oxfordian Limestone d18Odata:

(i) have random-like distribution through theoretical d18O-W/R curves and(ii) suggest that water/rock ratios (0.08–0.4) needed to explain d18O changes are higher by a factor of

about 2–20 compared to the present-day water/rock ratio.

These features indicate advection in both aquifers. According to the history of the Paris basin, thishydrogeological behaviour could have been effective since Jurassic/Cretaceous transition times. Inversely,the carbonate-d18O content trends observed for the Callovo-Oxfordian data show that changes were con-trolled by post-depositional fluid–rock interaction with water/rock ratio (0.02–0.15) similar to the pres-ent-day porewater/rock ratio. The 130 m thick Callovo-Oxfordian claystone appears remarkablyhomogenous regarding its hydrogeological properties. This study suggests an initial marine porewaterreplacement by meteoric water only after porosity was reduced to its present value, thus demonstratingthat the Callovo-Oxfordian clay has mainly been isolated from advective meteoric water circulation. Onlythe upper 20 m of the Callovo-Oxfordian claystone Formation underwent heterogeneous water–rockexchange (W/R from 0.01 to 0.3), probably as a result of its mineralogical heterogeneity and proximityto the advective Oxfordian Limestone aquifer.

This study of carbonate-d18O confirms the hydrogeological barrier properties of the Callovo-Oxfordianclay and suggests that it has been a natural hydrological barrier since the earliest times of its diagenesis.

� 2011 Elsevier Ltd. All rights reserved.

1. Introduction

During burial diagenesis, chemical and isotopic compositions ofporewater and carbonate minerals are increasingly influenced by

ll rights reserved.

versité Jean Monnet, F-42023+33 (0) 477 485 108.e.fr (V. Lavastre).MR-CNRS 5635, Place Eugène

the effects of water–rock interaction (Morton and Land, 1987)which is controlled by increasing temperature and progressivehydrogeological evolution of sedimentary basins (e.g. Longstaffe,1983, 1987; Land, 1984; Miliken et al., 1981; McBride et al.,1987; Ayalon and Longstaffe, 1988; Dutton and Land, 1988).Carbon and O stable isotopes give valuable insights into diagenetichistories of sediments with carbonate phases (Emery, 1988).

This study investigates the d18O and d13C records of the carbon-ate-containing Upper Bathonian, Callovian and Oxfordian sequencesfrom the eastern part of the Paris basin (France). These Formations

Fig. 1. Eastern Paris basin (France) geological map with location of the studied borehole HTM102 (white circle). URL = Andra Underground Research Laboratory.

V. Lavastre et al. / Applied Geochemistry 26 (2011) 818–827 819

were deposited on the north-western Tethysian shallow-platform,under the influence of both the North Sea and the opening AtlanticOcean (Dercourt et al., 1994; Ziegler, 1990). These marine Forma-tions are currently being studied by Andra (National French Agencyfor radioactive waste management) and CNRS, because of the highconfinement potential of the Callovo-Oxfordian argillites for radio-active-waste storage at depth. Andra has built an UndergroundResearch Laboratory (URL, Fig. 1) at Bure (Meuse/Haute-Marne)about 340 m deep in the Callovo-Oxfordian argillites which areabout 130 m thick and of very low permeability. The Callovo-Oxfordian argillite is a Formation overlying Bathonian limestoneand Lower Callovian (Upper Dogger hereafter) and overlain byOxfordian limestone (Fig. 2).

Fig. 2. Simplified lithostratigraphy of the HTM102 borehole. UB: Upper Bathonian;Call.: Callovian.

Documenting the geochemical effects on minerals of past andrecent hydrological events is essential for understanding and pre-dicting the hydraulic behaviour of the low permeability argillites.Previous studies have shown that limestones, organic matter andclay minerals from the Callovo-Oxfordian argillites were only af-fected by shallow burial and limited diagenesis (Landais and Elie,1999; Rousset and Clauer, 2003; Buschaert et al., 2004; Vincentet al., 2004; Clauer et al., 2007). All these studies support thehydrogeological barrier properties of the Callovo-Oxfordian clayFormation. This work aims to investigate O and C stable isotopecompositions in carbonates which are very sensitive tracers ofwater/rock interaction. It is specifically focused on isotope varia-tions with the objective of characterising and quantifying water/rock interaction experienced by carbonates throughout theirpost-depositional history.

2. Geology

2.1. Overview of the Paris basin

The Paris basin (France) corresponds to the north-western partof the Tethysian Sea. This basin is a late structural depressionwhose tectonic deformation has led both to the exposure of out-crops of middle Jurassic Formations along its eastern margin andto the uplift of bordering Palaeozoic massifs such as the Ardenno-Rhenan Massif in the East, the Armoricain Massif to the Westand the Central Massif in the South. The depression is filled byshallow-water platform sediments consisting of about 2900 m ofinterlayered clay Formations and limestone Formations depositedduring the Mesozoic transgressive event from 240 to 65 Ma ago.During this long event, the Paris basin was first connected to theTethysian Sea and then to both the Tethysian sea and the AtlanticOcean. At the studied site, Mesozoic Formations form a simplemonocline towards the central part of the basin.

During Cenozoic times (from 65 Ma to 1 Ma ago), the subsi-dence rate decreased, and at around 20 Ma a general uplift was ini-tiated in response to the Alpine orogen (Guillocheau et al., 2000).Several local transgressive events from the Atlantic Ocean via theEnglish Channel into the Paris basin allowed the deposition ofabout 250 m of Cenozoic sedimentary rocks (BRGM, 1980;Guillocheau et al., 2000). The present geological configuration ofthe basin evolved during Miocene tectonic events. During Middle

820 V. Lavastre et al. / Applied Geochemistry 26 (2011) 818–827

Pleistocene times (from 0.90 Ma to �0.25 Ma ago) the tectonic up-lift reached its maximum and the eustatic level fell definitively,inducing the general emersion of the Paris basin (BRGM, 1980;Guillocheau et al., 2000).

2.2. Details of the study area

2.2.1. Sampling and geological historyA total of 107 rock samples were collected on the eastern mar-

gin of the Paris basin, from the HTM102 borehole (48�2704000N,05�2200300E), which was located about 3 km SE of the URL site(Fig. 1). The studied stratigraphic sequence, presently found atdepths of 75–490 m, covers the Upper Dogger and Oxfordian stages(Dercourt et al., 1994; Ziegler, 1990). The studied Formations cor-respond to the transgression–regression cycle of the UpperBathonian and Middle Oxfordian (Pellenard et al., 1999). Themaximum burial depth reached by the Callovo-Oxfordian argillitesis estimated as 1000 m (Landais and Elie, 1999) and may haveoccurred during the Upper Cretaceous–Paleogene periods whenmaximum burial in the basin was reached (80–25 Ma).

2.2.2. Lithology and stratigraphyNumerous lithological units form the Upper Dogger to Upper

Oxfordian sequences (Fig. 2). At the base, the Upper Dogger (D4and C1 units) is composed of limestones deposited on a platformenvironment. Overlying these limestones, the middle Callovian tomiddle Oxfordian is composed of the so-called ‘‘Callovo-Oxfordianargillites’’, a shallow-water offshore clay-rich deposit (from med-ian upper (40–80 m) to distal lower >100 m; Collin and Courville,2006). It can be divided into 5 lithological units which are, fromthe base to the top (Fig. 2):

� C2a – silty calcareous argillite.� C2b1 – black bioclastic and pyritic silty calcareous argillite.� C2b2 – black silty calcareous argillite.� C2c – Interbedded black silty calcareous argillite in black marl.� C2d – Interbedded black bioturbated argillite and slightly bio-

turbated calcareous argillite.

Finally, at the beginning of Middle Oxfordian times, a sea-levelregression led to the formation of a carbonate platform (shoreface0–20 m; Collin and Courville, 2006) with a bioclastic deposit cover-ing the argillitic horizons and organized as follows (Fig. 2):

� C3a – dolomite/clayey limestones.� C3d – packstone/boundstone (dolomitic at the base) with tabu-

lar scleratinians.� L1a – bioclastic limestone.� L1b – oolitic limestone.� L2a – reef limestone.� L2b-L2c– oolitic limestone.� K1 – bioclastic/clayey limestone.

These lithological notations will be used in the followingdiscussion.

2.2.3. Mineralogy and hydrogeologyCalcite is the major carbonate mineral found in both Upper

Dogger and Oxfordian limestones. Calcite appears as a series ofearly micro-crystalline cements between oolites and euhedralcrystals of microspar around bioclasts and ooliths. Limestonematrices underwent later recrystallisation as a consequence ofmeteoric fluid circulation, during a period probably related toOligocene (23–35 Ma) extensional activity (Worden and Matray,1995; Buschaert et al., 2004). One of the consequences of this eventis the low permeability �10�11–10�9 m/s measured in both

limestone Formations (Buschaert et al., 2007). In addition, vugsand micro-fractures filled in by carbonates are observed mainlyin Oxfordian limestones and with fewer occurrences in UpperDogger Limestones (Buschaert et al., 2004). The porewater contentof the Upper Dogger is about 3 wt.%. Nevertheless, variations in theporosity (from 1–2% to 16%) have been observed throughout the230 m thick Oxfordian limestone. Seven main zones of relativelyhigh porosity with thickness varying from a few metres up to10 m (Buschaert et al., 2007), have been recognized as zones of rel-atively high water productivity. For both aquifers, groundwaterflows from the outcrops on the eastern part of the basin towardsthe centre of the basin and northwards to the English Channel(e.g. Menjoz and Lambert, 1991; Matray et al., 1994; Buschaertet al., 2007 and references therein). Geochemical and isotopic datashow:

(i) A considerable difference between groundwaters from theOxfordian and the Dogger aquifers. Oxfordian groundwatersare typically meteoric while some groundwater samplesfrom the Dogger aquifer may correspond to a mixture ofmeteoric-brine waters (Buschaert et al., 2007; Lavastreet al., 2010).

(ii) Feeding of both deep-aquifers by meteoric-water infiltration(e.g. Matray et al., 1994; Pinti and Marty, 1995; Giannesiniet al., 2005; Giannesini, 2006) along the outcrops exposedto the SE of the studied site for the Oxfordian, and along tothe SW to NE tableland on the eastern part of the Paris basinfor Dogger (see Lavastre et al., 2010 and reference therein).

(iii) Groundwater residence times exceeding 10 ka for the Oxfor-dian aquifer (from noble gas temperature data; Lavastreet al., 2010) and several hundred thousand years for the Dog-ger aquifer (according to 4He accumulation and the He4/He3

ratio; Marty et al., 1993, 2003; Castro and Jambon, 1998;Barthélemy et al., 2001; ANDRA, 2005; Lavastre et al., 2010).

Only small variations in mineralogical composition occurthroughout the 130 m thick Callovo-Oxfordian argillites, the threemajor constituents being clay minerals (30–40%), calcite (30%) andquartz (20–30%). The clay fraction contains kaolinite, chlorite, illiteand illite–smectite interlayers (Rousset and Clauer, 2003). Argil-lites are cemented by 25–30% of micro-crystalline calcite (Roussetand Clauer, 2003; Buschaert et al., 2004). Calcite is also found ineuhedral crystals, as sparitic secondary cement and as fossil testsessentially in the form of fragments (Pellenard et al., 1999;Buschaert et al., 2004; Gaucher et al., 2004). Argillites display verylow permeability, around 10�14–10�12 m/s (Delay et al., 2006) andno water-producing level was identified in the sedimentary unit.Nevertheless, porewater is an important component of this rock,between 6% and 9% in weight (Gaucher et al., 2004). Geochemicalstudies show that the Callovo-Oxfordian porewater at present isof meteoric origin (Girard et al., 2005; Giannesini et al., 2005;Giannesini, 2006; Buschaert et al., 2007). In the vicinity of the stud-ied site, previous work has shown that the Callovo-Oxfordian claysequence serves as an effective barrier to advective flow and trans-port between the underlying Upper Dogger aquifer and the overly-ing Oxfordian limestone (e.g. Buschaert et al., 2007; Lavastre et al.,2010). Molecular diffusion appears to be the main transfer processfor solutes throughout the Callovo-Oxfordian clay (Lavastre et al.,2005; Buschaert et al., 2007).

3. Methods

For each of the 107 core segments sampled, the outer part wasremoved in order to avoid possible contamination introduced bythe drilling mud.

Table 1Carbonate content, C- and O-isotope (versus PDB) data for bulk rock from the HTM102 borehole.

Lithological level Sample ID Depth, m d13C/PDB ± d18O/PDB ± % calcite

L2c HTM 2502a 75.00 3.10 0.10 �3.10 0.10 99HTM 2509a 86.00 3.02 0.10 �4.81 0.10 99HTM 0225a 87.00 3.39 0.10 �3.06 0.10 99HTM 2522a 111.00 3.02 0.10 �4.45 0.10 99HTL 2523a 112.00 3.05 0.10 �4.19 0.10 99HTM 0274a 116.00 3.07 0.10 �2.49 0.10 99

L2b HTM 2158a 135.00 2.82 0.10 �4.55 0.10 99HTM 0331a 141.00 2.71 0.10 �5.17 0.10 99

L1a HTM 2540a 218.00 2.66 0.10 �5.12 0.10 99HTM 2160a 226.00 2.85 0.10 �6.24 0.10 99HTM 2164a 251.00 3.13 0.10 �5.29 0.10 99HTM 2166a 265.00 2.77 0.10 �5.85 0.10 99

C3b HTM 2171a 289.00 2.10 0.10 �3.88 0.10 99HTM 2593 319.58 2.60 0.10 �3.50 0.11 95HTM 1716 320.84 2.10 0.05 �3.67 0.12 99HTM 2180a 322.00 2.46 0.10 �2.67 0.10 90HTM 2599 324.74 2.24 0.04 �2.79 0.10 85

C3a HTM 2182a 327.00 2.04 0.10 �2.67 0.10 75HTM 2605 327.69 1.65 0.03 �2.57 0.07 75HTM 679 329.41 2.21 0.05 �2.00 0.11 88HTM 80613 330.56 2.40 0.04 �2.75 0.08 92HTM 390 333.62 2.58 0.07 �2.49 0.09 75HTM 2611 336.10 2.48 0.05 �3.28 0.11 66HTM 2608 336.60 2.28 0.05 �2.44 0.13HTM 2089 340.35 2.41 0.03 �2.38 0.16 80HTM 2572 341.37 2.67 0.05 �4.46 0.16 44HTM 80618 342.48 2.14 0.05 �2.13 0.13 37

C2d HTM 739 343.36 2.30 0.15 �4.87 0.25 24HTM 2660 345.69 2.22 0.06 �1.32 0.14 82HTM 2659 346.93 2.48 0.10 �3.79 0.09 30HTM 759 348.96 2.31 0.07 �5.01 0.13 27HTM 781 352.04 2.21 0.06 �3.95 0.17 41HTM 2614 355.50 1.93 0.03 �4.40 0.11 48HTM 0801a 357.00 2.14 0.10 �3.93 0.10 40HTM 80572 358.06 1.97 0.05 �4.34 0.15 32HTM 2629 359.26 1.70 0.04 �5.05 0.15 63HTM 0814 361.16 2.09 0.08 �6.21 0.10 23HTM 2618 363.05 1.90 0.07 �4.69 0.06 42HTM 2348 364.48 2.00 0.09 �3.05 0.18 35

C2C HTM 0841 366.96 1.81 0.11 �3.53 0.14 18HTM 2210 368.35 1.91 0.06 �2.71 0.16 36HTM 80577 369.85 1.63 0.05 �2.44 0.09 58HTM 0854a 371.00 2.24 0.10 �4.04 0.10 59HTM 2100 371.04 1.53 0.03 �2.06 0.09 59HTM 0856 372.24 2.11 0.07 �3.93 0.09 27HTM 0875 375.70 1.79 0.06 �2.99 0.15 30HTM 0881 377.42 2.17 0.10 �3.57 0.13 18HTM 0883 378.31 1.66 0.05 �3.61 0.11 34HTM 0906 382.12 1.74 0.06 �3.11 0.13 27HTM 80628 384.97 1.50 0.05 �2.72 0.15 34HTM 0927 386.20 1.97 0.10 �2.89 0.12 33

C2b2 HTM 0938 388.93 1.46 0.06 �3.13 0.15 27HTM 0942 390.01 1.88 0.05 �3.68 0.11 24HTM 0946 392.00 2.04 0.10 �2.66 0.10 25HTM 0957 394.17 1.88 0.05 �3.69 0.14 27HTM 0966 396.69 1.87 0.05 �3.51 0.08 50HTM 0984 400.09 2.06 0.10 �3.50 0.14 27

C2b1 HTM 0996 402.89 1.93 0.10 �4.60 0.16 17HTM 1001a 405.00 1.89 0.10 �3.67 0.10 24HTM 1004 405.75 1.84 0.17 �3.54 0.18 24HTM 1010 407.63 2.30 0.04 �3.84 0.16 25HTM 2112 409.52 2.00 0.04 �4.01 0.10 21HTM 1030 412.18 1.90 0.07 �3.38 0.11 29HTM 1045 415.12 1.77 0.08 �4.23 0.15 21HTM 1055a 417.00 2.02 0.10 �3.83 0.10 19HTM 1057 417.91 1.63 0.04 �4.33 0.16 19HTM 1066 420.73 1.50 0.05 �4.79 0.14 25HTM 1075a 423.00 1.73 0.10 �2.72 0.10 28HTM 1078 423.21 1.74 0.06 �4.26 0.17 28HTM 1083 424.57 1.72 0.07 �4.67 0.16 17HTM 1086 425.76 1.35 0.05 �3.69 0.13 28

(continued on next page)

V. Lavastre et al. / Applied Geochemistry 26 (2011) 818–827 821

Table 1 (continued)

Lithological level Sample ID Depth, m d13C/PDB ± d18O/PDB ± % calcite

HTM 1094 428.12 1.66 0.05 �3.72 0.11 27HTM 1104 429.47 1.33 0.05 �3.40 0.10 27HTM 1109 431.22 1.54 0.05 �4.07 0.17 28HTM 1120 433.71 1.13 0.08 �2.93 0.19 25HTM 1127 435.11 0.91 0.05 �3.21 0.07 46HTM 1133 436.61 0.99 0.05 �3.29 0.15 30HTM 1143 438.39 0.79 0.06 �2.69 0.10 46HTM 2156a 439.00 0.56 0.10 �2.57 0.10 50HTM 80597 440.22 1.23 0.04 �2.45 0.08 53

C2a HTM 1151 441.23 1.75 0.08 �3.62 0.12 40HTM 1158 443.58 0.69 0.06 �3.48 0.11 33HTM 1161 445.03 1.03 0.02 �3.70 0.11 27HTM 1171 447.47 0.73 0.06 �4.33 0.14 23HTM 1179 449.52 1.23 0.06 �4.58 0.06 27HTM 1185a 451.00 1.17 0.10 �3.47 0.10 25HTM 80633 453.84 1.36 0.08 �4.45 0.13 23HTM 1201 455.81 0.59 0.07 �3.83 0.15 25HTM 1213 458.97 0.95 0.05 �4.66 0.14 26HTM 5282 460.64 1.10 0.04 �4.92 0.10 25HTM 1227a 463.00 1.29 0.10 �3.82 0.10 27HTM 2591 463.49 1.16 0.05 �4.43 0.11 27HTM 80755 465.20 1.22 0.06 �3.79 0.11 35HTM 80741 467.89 1.29 0.07 �3.56 0.13 32HTM 80737 469.64 1.26 0.04 �3.62 0.07 37HTM 2350 471.51 �0.28 0.05 �3.19 0.06 59

C1 HTM 2632 473.93 1.75 0.06 �4.05 0.05 95HTM 2637 477.63 1.86 0.06 �4.31 0.09 97HTM 1279 479.00 1.84 0.05 �3.30 0.13 83HTM 2642 479.97 2.00 0.04 �3.54 0.07 84HTM 2565a 480.00 2.02 0.10 �4.10 0.10 85HTM 1281 481.18 1.84 0.07 �4.18 0.05 87

D4 HTM 2648 484.96 1.53 0.03 �4.50 0.12 96HTM 2186 485.00 1.17 0.10 �4.53 0.10 95HTM 2651 487.27 1.76 0.04 �3.75 0.09 95HTM 2653 488.58 1.65 0.02 �4.13 0.05 98HTM 2654 490.86 1.68 0.07 �4.40 0.11 97

a Samples analysed at the Geosciences Rennes Laboratory, Université de Rennes, France.

822 V. Lavastre et al. / Applied Geochemistry 26 (2011) 818–827

Carbonate d13C and d18O measurements were performed on fi-nely crushed bulk-rock samples. It was found by preliminary pyro-lysis tests on bulk-rock samples at 375 �C under vacuum, that thelow organic matter content (<1%) does not affect the isotopic anal-ysis of bulk carbonates. Subsequently, no organic matter (OM here-after) processing was applied to the rest of the samples.

Seventy nine bulk carbonate samples (see Table 1) were ana-lysed for their C and O isotope compositions at the Laboratoirede Géochimie des Isotopes Stables, Institut de Physique du Globede Paris (France) using an Analytical Precision continuous flowMass Spectrometer (AP2003) equipped with an automated carbon-ate preparation system.

About 1–4 mg of sample powder was placed in a 12 mL Exetain-er� vial, which was subsequently flushed with He. Essentiallyfollowing a McCrea (1950)-type procedure, 200 lL of 100% phos-phoric acid were injected into the flushed vial and the acid reactionperformed at 25 �C for 10 h. After completion of the reaction, asample gas aliquot was introduced into an injection loop. Watervapour was completely removed by a Nafion trap, and isothermalGC separated CO2 from other gases. The same procedure was ap-plied to pure CaCO3 laboratory standards calibrated versus theinternational standard IAEA-NBS19. These laboratory carbonatestandards were incorporated in each set of samples, in order toperform an automatic data-correction from the isotopic fraction-ation occurring during acid attack. Finally, the d13C and d18O resultsare reported relative to the PDB international standard. The car-bonate content was derived from the intensity of the 44 mass/charge signal using a calibration based on varying amounts of lab-oratory standard. The overall reproducibility of sample replicate

measurements is better than ±0.17‰ for d13C and than ±0.26‰

for the d18O.The remaining 24 samples (see Table 1) were analysed at the

Geosciences Laboratory, Université de Rennes (France) againfollowing the McCrea (1950) method. The sample powder(5–10 mg) was reacted with 100% orthophosphoric acid at 25 �Cassuming a(O) = CO2/CaCO3 (extraction) = 1.01025 (Sharma andClayton, 1965). Isotopic analyses were carried out on CO2 gas usinga VG SIRA 10 mass spectrometer and were expressed with the con-ventional d notation versus PDB. Here the average precision for Cand O analyses were, respectively, 0.04‰ and 0.09‰. Uncertaintiesrelated to the PDB scale are estimated to be <0.1‰ for C and ca.0.1‰ for O.

Isotopic measurements from both methods are in perfectagreement.

4. Results

Carbonate contents vary from 84% to 98% in the Upper Doggerlimestone, 17% to 82% in the Callovo-Oxfordian argillites (with anaverage content of about 33%) and from 44% to 99% in the Oxfor-dian limestone (Table 1 and Figs. 3 and 4).

d18O/PDB values vary between �6.24‰ and �1.32‰, without ageneral trend along the profile (Fig. 4). A correlation between car-bonates d18O and carbonate content can be observed for the C2a,C2b1 and C2b2 lithological units of the argillites (Fig. 3): highd18O values usually corresponding to carbonate-rich levels.However, the 10 m thick C2d level shows high amplitude varia-tions for both d18O and carbonate content without any correlation

Fig. 3. d18O versus d13C for carbonates in the Upper Dogger Limestone, Callovo-Oxfordian clay and Oxfordian Limestone bulk rocks. The d18O–d13C area for undisturbedmarine Jurassic calcite (Podlaha et al., 1998) is shown in grey colour. Trend 1 represents marine calcite-d18O evolution during meteoric diagenesis with water/calcite ratiolower than 2.6 (Banner and Hanson, 1990). Trend 2 represents the evolution of d18O and d13C for marine calcite in contact with large meteoric water volume (water/calciteratio >2.6; Banner and Hanson, 1990) and/or for marine calcite affected by early diagenesis (hatched area; e.g. Hudson, 1977; Glasmann et al., 1989; Ader and Javoy, 1998).

Fig. 4. Carbonate content and C- and O-isotope fluctuations versus depth in bulk rock from the HTM102 borehole.

V. Lavastre et al. / Applied Geochemistry 26 (2011) 818–827 823

between the two parameters (Table 1 and Fig. 4). For both UpperDogger and Oxfordian limestones, strong d18O variations are ob-served for constant carbonate content (Table 1 and Fig. 4).

d13C values range from �0.28‰ to 3.39‰ and are neither corre-lated with d18O values nor calcite content (Figs. 3 and 4). Along thedepth profile, d13C increases by about 2‰ from base to top, follow-ing an irregular oscillating trend (Fig. 4).

5. Discussion

18O/16O and 13C/12C ratios measured on bulk Upper Dogger,Callovo-Oxfordian and Oxfordian carbonate rocks from the eastern

Paris basin, correspond to isotopic compositions of Jurassic marinecarbonates worldwide (Veizer et al., 1999). When comparing thebulk carbonates isotope data set obtained for the Upper Dogger–Callovian–Oxfordian sequence with the one obtained for undis-turbed Jurassic belemnite samples from different localities in theworld (Podlaha et al., 1998), the d13C values are in agreement withthese data, while the d18O values are lower by 0–5‰. These obser-vations suggest that water–rock interaction has disturbed the mar-ine d18O without modifying the d13C of the studied carbonatedsequence.

It is generally accepted that during water–rock interac-tion chemical exchanges are controlled by both (i) the relative

824 V. Lavastre et al. / Applied Geochemistry 26 (2011) 818–827

concentrations of the element in the fluid and rock and (ii) water/rock ratio (W/R) (e.g. Banner and Hanson, 1990; Derry, 2010).

If fluids have a low ratio of C/O compared to carbonates, modi-fications of the carbonate-d18O are expected to occur at much low-er W/R ratios than those necessary to significantly shift the d13C(Banner and Hanson, 1990). If fluids have high C/O ratios, withthe C derived from early diagenetic OM mineralisation, or burialdiagenesis OM thermal decomposition, their interaction with car-bonates generally leads to a positively correlated decrease of bothd13C and d18O values (e.g. Hudson, 1977; Glasmann et al., 1989;Ader and Javoy, 1998; Derry, 2010).

The undisturbed d13C values reported in the present data setand the absence of positive correlation between d13C and d18O val-ues, point to water–rock exchange with a C-poor fluid. This infer-ence is consistent with a previous study by Buschaert et al.(2004), which proposed crystallisation of late calcite from vugs,fractures and cemented interparticle porosity in Oxfordian Lime-stone from C-poor fluids and a local C source (by dissolution of hostcarbonates).

It can, therefore, be proposed that the Upper Dogger–Callovian–Oxfordian sequence experienced chemical exchange with fluidswith low C/O, affecting carbonate-d18O but not d13C. The d13Crecord should, therefore, be representative of the environmentalconditions of carbonate precipitation during Upper Dogger,Callovo-Oxfordian and Oxfordian times.

5.1. Evidence for water–rock interactions from the carbonates d18Orecord

The association of positive d18O shifts with carbonate-poor lev-els in argillites (Fig. 4) suggests that isotopic variations are con-trolled by water–rock exchange, and hence by water/rock ratios.Previous studies having identified only minor diagenetic effectson clay minerals and organic matter for the Callovo-Oxfordianargillites (Landais and Elie, 1999; Rousset and Clauer, 2003; Claueret al., 2007). For the present attempt to quantify the effects of dia-genesis isotopically, it is considered that O-isotope exchange oc-curred exclusively between carbonate minerals and porewater.

5.2. Equation and parameters

O-isotope exchange between carbonates and porewater can bequantified by using a simple closed system mass balance equation(e.g. Sheppard et al., 1969; Taylor, 1974):

W=C ¼ dcarbonate f � dcarbonate i

dH2Oi� dcarbonate f � D� � ð1Þ

where W/C is the water–carbonate ratio expressed in terms of oxy-gen molar ratio, dcarbonate f is the measured carbonate-d18O, dcarbonate i

is the initial carbonate-d18O, dH2O i is the initial porewater d18O andD = dcarbonate f � dH2O f.

In order to calculate theoretical curves representing the evolu-tion of dcarbonate during water–rock exchange for a given W/R ratioand varying carbonate contents, the water–carbonate molecularratio (W/C) used to calculate the effect of water–carbonate interac-tion on d18O (dcarbonate f), is converted to water–rock ratio (W/R)using the following relationship:

ðW=RÞmass ¼ ðW=CÞ � \Carbonate content"100� 2:7

ð2Þ

where ‘‘carbonate_content’’ is expressed in % and the 2.7 factor cor-responds to the oxygen molecular mass ratio: Mcalcite/(3 �Mwater).

The resulting carbonate-d18O evolution curves for variouswater–rock ratios are reported in d18O versus carbonate contentdiagrams in Fig. 5.

The parameters used for this calculation are:

– dcarbonate i = �1‰/PDB: mean Jurassic d18O value for undisturbedcarbonates (e.g. Podlaha et al., 1998).

– D � 1000 � ln acalcite-water = 18.03 � (103 � T�1) � 32.42 (Kim andO’Neil, 1997).

– dH2O i = �6‰/SMOW. This value corresponds to the d18O mea-sured on present-day porewater from Callovo-Oxfordian argil-lites (Girard et al., 2005). In addition it is within the �8.5‰ to�5‰ d18O range measured in Oxfordian groundwater (Gianne-sini et al., 2005) and the d18O range estimated for the fluid fromwhich late-forming micritic cement crystallised in the Oxfor-dian limestone (�6.8‰ to �2.5‰; Buschaert et al., 2004); andit is close to the 9.5‰/�8‰ d18O range measured for UpperDogger groundwater (Giannesini et al., 2005).

– At T = 40 �C. This is a reference value established considering thetemperature data obtained for the maximum burial temperaturereached by the Formations during burial up to about 1000 m indepth at temperatures of 32–42 �C recorded by fluid-inclusionmicrothermometry of Oxfordian limestones; 35–48 �C extrapo-lated for Callovo-Oxfordian argillites (Buschaert et al., 2004),and 50 �C for a brief thermal event or 40 �C for an isothermalevent at �160 Ma according to organic-tracer studies (Landaisand Elie, 1999; Clauer et al., 2007 and references therein).

5.3. Uncertainties

During water–rock interaction, both water isotopic compositionand carbonate isotopic composition evolve to reach an intermedi-ate value to initial water d18O and to initial carbonate-d18O. Isoto-pic exchange will be higher with increasing temperature, and/orwith increasing difference between initial carbonate-d18O and ini-tial water d18O and/or with increasing water/rock ratio. Therefore,the chosen values for model parameters described above implysome uncertainties which are discussed below.

Several possibilities could be considered for porewater originand water/rock interaction history (i) isotopic exchange with onlymarine water controlled by the increasing temperature during bur-ial (ii) isotopic exchange with marine water during early stage ofburial and later with meteoric water (iii) isotopic exchange withonly meteoric water. Because the current porewater in the studiedcarbonate rocks is meteoric, the first hypothesis was discarded. Inorder to simplify calculations and to estimate a maximum W/R ra-tio, it is considered that isotopic exchange occurred exclusively inthe presence of meteoric water which has a d18O (<0‰) which isvery different from marine water d18O (close to 0‰) in equilibriumwith marine carbonate at the time of precipitation. To consolidatethis approach, an initial porewater d18O equal to the d18O mea-sured in present-day porewater is assumed. This does not take intoaccount the porewater d18O change during the W/R interactionprocess which predicts an initial porewater d18O lower than the ac-tual value of about �6‰ (Girard et al., 2005). This choice contrib-utes to overestimation of the W/R ratio.

For this study, it is considered that isotopic exchange occurred atmaximum burial (i.e. maximum temperature). This hypothesis im-plies an overestimate of isotopic exchange and so an overestimateof calculated W/R ratios. Considering the chosen temperature ofabout 40 �C and the possible maximum temperature range (32–50 �C) defined by previous experimental work (Buschaert et al.,2004; Landais and Elie, 1999), the distribution of data through theW/R theoretical curves on Fig. 5 has an uncertainty of about ±20%.

5.4. Quantifying water–rock ratios

In this section meteoric water is considered as a fluid, which cansignificantly disturb carbonate isotopic compositions in marine

Fig. 5. Distribution of O-isotope data on d18O versus carbonates content diagram. Theoretical curves representing the evolution of marine carbonate-d18O at a temperature ofabout 40 �C for a given meteoric water/rock ratio and for various rock carbonate contents are represented by black lines.

V. Lavastre et al. / Applied Geochemistry 26 (2011) 818–827 825

826 V. Lavastre et al. / Applied Geochemistry 26 (2011) 818–827

rocks. According to Eq. (1), and considering the parameters given inthe previous section (Section 5.2), theoretical curves representingthe evolution of carbonate-d18O during O-exchange with meteoricwater for a given W/R, varying rock carbonate content, and a tem-perature about 40 �C, are reported on Fig. 5 together with carbon-ate-d18O data for the samples under investigation. The mainfeature observed on Fig. 5 is the Callovo-Oxfordian argillite datadistribution along the 0.02–0.15 W/R ratio curves which points toa control of carbonate-d18O values by water–rock ratios.

In contrast, according to the model, for both Upper Dogger andOxfordian limestones, the water/rock ratios controlling carbonate-d18O values vary between 0.08 and 0.4 and the data do not followW/R curves. These W/R values are higher by a factor of about 2–20than the present-day W/R (see Section 2.2.3.). Two possible scenar-ios can be invoked to explain the higher-water volumes needed toexplain carbonate isotopic changes:

– Most water–rock interaction occurred during the early stage ofdiagenesis in a partially compacted Formation (i.e. with a higherporosity) and in a closed system.

– Isotopic modification of Upper Dogger and Oxfordian carbon-ates occurred in an open system with water replacement duringand/or after the compaction stage.

The latter hypothesis is the one preferred, and supported by:

(i) Meteoric-water infiltration through the Oxfordian Lime-stone, probably occurring during late Oligocene(23–35 Ma) times, and is recorded by the d18O of late calcitecement in vugs and micro-fractures (Buschaert et al., 2004).

(ii) The fact that Oxfordian and Upper Dogger limestones hostdynamic aquifers (e.g. Matray et al., 1994; Lavastre et al.,2010). Their water residence time exceeds 104 ka for theOxfordian (Lavastre et al., 2010) and several hundred millionyears for the Upper Dogger (e.g. Marty et al., 2003; Lavastreet al., 2010).

Callovo-Oxfordian argillite data fit very well with 0.02 and0.15 W/R curves, except for the C2d level for which the data arestrongly scattered between the 0.01 and 0.2 W/R curves (Fig. 5).These features indicate that for most stratigraphic units in theCallovo-Oxfordian (except the C2d level), O-isotope disturbanceof carbonate minerals may have been controlled by a given W/Rratio occurring in a relatively confined system. The fact that theW/R ratios needed to obtain the measured carbonate-d18O are veryclose to the present-day ratio in the Callovo-Oxfordian clay (about0.06 to 0.09; Gaucher et al., 2004), points to meteoric-waterreplacement of marine water in Callovo-Oxfordian pore spacesonly after the initial porosity had been strongly reduced.

Significant diagenetic effects (d18O varying from –6.2 to 1.3‰/SMOW and W/R ranging from 0.01 to 0.3) with variable intensityare observed in the C2d unit (upper 20 m of the Callovo-Oxfordianclay). This feature could be explained by heterogeneous porosityand the variable carbonate abundances in a specific stratigraphicunit (see Fig. 4 for calcite content).

Because (i) no advective transport is observed at the presenttime in the C2d unit and (ii) a W/R ratio higher than 0.1 is notcompatible with present-day W/R in Callovo-Oxfordian claystone,the results show that the strong water/rock interaction undergoneby several layers of C2d unit occurred earlier in the geologicalrecord and thus are products of water paleocirculation. The lateOligocene meteoric-water infiltration observed in the neighbour-ing Oxfordian Limestone (Buschaert et al., 2004) could also beinvoked to explain water–rock interaction recorded throughoutthe C2d unit.

6. Conclusions

This study highlights the contrasting hydrogeological behaviourof the Callovo-Oxfordian argillites compared to its neighbouringstratigraphically lower and upper limestone horizons.

The evaluation of water–rock interactions using O isotope com-positions in the Upper Dogger–Oxfordian sequence shows that thed18O-carbonate values for the Callovo-Oxfordian claystone are con-trolled by low W/R ratios. This suggests that the Formation evolvedas a confined system and that the O isotopic composition of car-bonate minerals underwent only slight changes during theirpost-depositional history. The hydrogeological events responsiblefor the isotopic effects result mainly from a replacement of initialmarine porewater by meteoric water after the present-day poros-ity characteristics had been reached. In such a low permeabilitymedium, only diffusion of meteoric water derived from associatedlimestones can explain the presence of marine porewater flushingthrough the Callovo-Oxfordian Formation (Waber et al., 2005;Clauer et al., 2007 and references therein). The Callovo-Oxfordianclay Formation shows a remarkable homogeneity of water–rockinteraction intensity throughout most of its thickness (between365 m and 472 m in depth). This feature underlines the homoge-neous paleohydrogeological and hydrogeological behaviour of thedifferent lithological units constituting the Callovo-Oxfordian For-mation. Only the upper beds of the Callovo-Oxfordian claystone, atthe C2d level (about 20 m thick), suffered stronger water–rockinteraction. The C2d level probably played a key role as a buffer be-tween the Callovo-Oxfordian Formation and the Oxfordian aquifer.

In contrast, d18O evolution of Upper Dogger and Oxfordian lime-stones occurred in an open system. Both limestones suffered mete-oric water inflow, still active at the present-day as demonstratedby this study and by previous works (e.g. Menjoz and Lambert,1991; Matray et al., 1994; Buschaert et al., 2004; Lavastre et al.,2005, 2010). Meteoric recharge of both Oxfordian and Upper Dog-ger limestones could have occurred from the uppermost Jurassicand Upper Aptian during major periods of basin-scale emersion(Vincent et al., 2004; Clauer et al., 2007 and reference therein).

The carbonate-d18O records shows that the Callovo-Oxfordianclay Formation was clearly isolated from mass water transferoccurring in the Upper Dogger and Oxfordian Limestones throughgeological time. This work thus suggests that the Callovo-Oxfor-dian clay strata have been a natural hydrological barrier since theirearliest diagenesis.

Acknowledgments

The authors would like to acknowledge ANDRA for having madeavailable the rock samples from the eastern Paris basin, France, andfor V. Lavastre’s PhD thesis grant. Peter Bowden and two anony-mous reviewers are gratefully acknowledged for their constructivereviews. Nathalie Jendrzejewski, Jean-Jacques Bourrand andMichel Girard are thanked for their technical assistance during thisstudy. This study was jointly financed by CNRS and ANDRAthrough the GdR FORPRO (Research Action No. 98.IV) and corre-sponds to GdR FORPRO contribution number FORPRO 2007/14.

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