Early diagenetic carbonate bed formation at the sediment–water interface triggered by synsedimentary faults

Chemical Geology 300-301 (2012) 1–13

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Research paper

Early diagenetic carbonate bed formation at the sediment–water interface triggered by synsedimentary faults Nicolas Tribovillard a,⁎, Pierre Sansjofre b, Magali Ader b, Alain Trentesaux a, Olivier Averbuch a, Florent Barbecot c a b c

Université Lille 1 & CNRS, laboratoire Géosystèmes UMR 3298, 59655 Villeneuve d'Ascq cedex, France Equipe de géochimie des isotopes stables, Institut de Physique du Globe de Paris (IPGP), UMR 7154 CNRS, Sorbonne Paris Cité, Univ Paris Diderot, 75005 Paris, France Université Paris-Sud & CNRS, UMR IDES 8148, 91405 Orsay cedex France

a r t i c l e

i n f o

Article history: Received 22 September 2011 Received in revised form 17 January 2012 Accepted 18 January 2012 Available online 24 January 2012 Editor: U. Brand Keywords: Carbonate early diagenesis Sulfate reduction C, O and S isotope geochemistry Black shales Boulonnais Late Jurassic

a b s t r a c t The latest Jurassic of the Boulonnais cliffs (N-France, Strait of Dover) corresponds to a homoclinal ramp-type depositional environment, where sedimentation was clastic-dominated (marlstones-sandstones). The Tithonian marlstone formations exhibit two or three carbonate beds in the vicinity of Wimereux city (3 beds at Wimereux-North and 2 beds at Wimereux-South). The beds are made up with quite pure, finegrained (microspar) carbonate in sharp contrast with the background sedimentation. This paper presents a sedimentological and geochemical study of these limestone beds, aiming to determine whether they are of diagenetic origin and hence whether they could have acted as permeability/migration barriers during hydrocarbon maturation/migration. Mineralogical, chemical, and C and O stable isotope data allow us to infer that the beds formed during synsedimentary early diagenesis at the sediment–water interface (or close to it) as the result of a rise of alkalinity induced by bacterial sulfate reduction. The rise of alkalinity was not counter-balanced by the accumulation of H2S released by sulfate reduction and carbonate ion supersaturation was rapidly reached, causing the formation of laterally-continuous limestone beds. Conditions prone to bacterial sulfate reduction developed episodically at the sediment–water interface as the result of spills of anoxic pore waters onto the seafloor. These spills were probably released by synsedimentary fault movements. Such continuous limestone beds being formed under rather common conditions during the earliest stage of diagenesis of shale deposits must be more frequent in the geological record than hitherto identified. © 2012 Elsevier B.V. All rights reserved.

1. Introduction The geological formations of the Late Jurassic times (Kimmeridgian– Tithonian) crop out for about 25 km along the Boulonnais Cliffs (Strait of Dover, Northern France; Fig. 1A). They represent a proximal, lateral equivalent of the Kimmeridge Clay Formation (KCF) and they accumulated in a clastic-dominated ramp environment. In the Boulonnais, the Late Jurassic is made up of an alternation of marlstone-dominated formations (called “Argiles” in Fig. 2) and sandstone-dominated formations (termed “Grès” in Fig. 2) expressing the range of deposition depths from lower offshore to shoreface settings, respectively (e.g., Proust et al., 1995; Al-Ramadan et al., 2005). The transition between the Argiles de la Crèche and Argiles de Wimereux formations (namely, the so-called Bancs Jumeaux Formation, Wheatleyensis+ Pectinatus ammonite zones, see below) represents a low-energy shelf facies deposited below wave base, but with some storm influence expressed as thin shelly limestone interbeds (Wignall, 1991; Proust et al., 1995;

⁎ Corresponding author. E-mail address: [email protected] (N. Tribovillard). 0009-2541/$ – see front matter © 2012 Elsevier B.V. All rights reserved. doi:10.1016/j.chemgeo.2012.01.014

Deconinck et al., 1996; Wignall and Newton, 2001; Williams et al., 2001). The Bancs Jumeaux Formation owes its name to the occurrence of two limestone beds (“bancs jumeaux” means twin beds) cropping out to the south of the seaside town of Wimereux. Unexpectedly, North of Wimereux, three beds are observed (Fig. 1B and C). Furthermore, these carbonate beds capture the attention because they contrast sharply with the lithology of the Argiles de la Crèche and Argiles de Wimereux formations, in that they are made of quite pure limestone in a depositional setting where the terrigenous fraction dominates the rocks. Consequently, it may be wondered whether these beds result either from an episode of sedimentation more prone to carbonate deposition (as is observed for the KCF; e.g., Lees et al., 2004 or older sedimentary rocks in the Boulonnais, namely, the Calcaires du Moulin Wibert Formation) or from diagenetic processes. In addition, should these limestone beds be of diagenetic origin, the question is then about the timing of their formation (early or late diagenesis). When we speak of early diagenesis, we include the fact that the formation of the beds may occur at the sediment–water interface, in a synsedimentary way. The question of the timing of the limestone bed formation is important because such beds create migration barriers within a succession of formations making up what can be assimilated to a petroleum source–rock

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N. Tribovillard et al. / Chemical Geology 300-301 (2012) 1–13

A

B

Pointe de la Rochette

United Kingdom

outcrop with 3 beds

r

ve

o fD

Folkestone

its

o

ra St

Calais

Wimereux River

Cap Blanc-Nez Wimereux Boulogne sur mer

Wimereux City

France

Cap de la Crèche

C

outcrop with 2 beds)

500 m

Fig. 1. A Location of the outcrops where the Bancs Jumeaux Formation is visible alongshore in the Boulonnais area with either three carbonate beds (B) at Wimereux-North (01°36′ 20.7″ E and 50°46′34.4″ N, or x = 548336 and y = 2642736 Lambert IIe) or two beds (C) at Wimereux–South (01°35′59.8″ E and 50°45′13.2″ N, or x = 547896 and y = 2640214 Lambert IIe). The maps were drawn using two websites, www.viaMichelin.com and www.planiglobe.com. B: the intermediate carbonate bed is continuous in spite of what is suggested by the picture due to the poor quality of the outcrop (marls are creeping).

system. To answer these questions, we studied the lithology and geochemistry (elemental and isotopic composition) of the carbonate beds and compared them to those of the caging marls. 2. A brief overview about early diagenetic carbonate formation Carbonates can form during early diagenesis as the result of prokaryotes metabolic activity (e.g., Folk and Chafetz, 2000; Peckmann and Thiel, 2004 and references therein). A vast range of chemical reactions is achieved by these organisms to meet their vital requirements as far as carbon sources (anabolism or assimilatory metabolism) and energy sources (catabolism or dissimilatory metabolism) are concerned. Such reactions may result either in a rise or in a fall of the pH and/or alkalinity, which may in turn foster or hamper carbonate ion supersaturation in pore waters, hence calcium carbonate precipitation; the interested reader is orientated toward the comprehensive syntheses by Megonigal et al. (2003) and Burdige (2006, notably chap. 7 and 16). Organic-matter degradation under aerobic conditions is considered as favoring carbonate dissolution in aqueous environment, at least preventing carbonate precipitation. The bioinduced or bio-mediated chemical reactions that lead to carbonate

precipitation operate under anaerobic conditions and, among the various categories of anaerobic bacterially-mediated reactions, the dominant ones are those linked to bacterial sulfate-reduction reactions (conducted by sulfate-reducing bacteria) and anaerobic oxidation of methane (conjointly operated by consortia of anaerobic methaneoxidizing archaea and sulfate-reducing bacteria; e.g., Hinrichs et al., 1999; Boetius et al., 2000; Birgel et al., 2011; Holmkvist et al., 2011; Regnier et al., 2011). A large literature is devoted to the two groups of reactions that are strongly linked with sedimentary organic matter decomposition and with many aspects of the sulfur cycle, especially iron sulfides and pyrite formation; see reviews or syntheses by Coleman and Raiswell (1993); Goldhaber (2003); Megonigal et al. (2003); Jørgensen and Nelson (2004); Burdige (2006); Taylor and Macquaker (2001), among so many others. Basically, sulfate reduction and anaerobic methane oxidation reactions may be summarized by the following equations, respectively: 2−

−

SO4 þ 2CH2 O−> H2 S þ 2HCO3 2−

−

−

SO4 þ CH4 −> HCO3 þ HS þ H2 O

ð1Þ ð2Þ

N. Tribovillard et al. / Chemical Geology 300-301 (2012) 1–13

Base of the Argiles de Wimereux Fm.

Bancs Jumeaux Fm.

P2 150

100

P1

Top of the Argiles de la Crèche Fm.

Wimereux North

S (wt%)

Org. C (wt%)

200

50

Base of the Argiles de Wimereux

cm 0 0 200

0

5

10 15 20

5

10 15 20

25 50 75 100

0

0

1

2

3

4

1

2

3

40

2

4

6

8

1,5

3

4,5

6

P2

150

Bancs Jumeaux Fm.

Wimereux South

CaCO 3 (wt%)

Mode (µm)

3

100

Top of the Argiles de la Crèche Fm.

P1

50

cm 0 0

0

25 50 75 100 0

Fig. 2. Stratigraphic distribution of the contents in CaCO3, organic carbon and total sulfur, and of the mode of the grain size distribution of the b 2 mm, decarbonated fraction of the sediment, for the studied two sections. P1 and P2 are two centimetric levels rich in phosphatic nodules and shell fragments, plus sparse centimetric quartz pebbles. The intermediate carbonate bed of the Wimereux-North Section is nodular.

The alkalinity generated by both reactions may favor precipitation of authigenic carbonates in the shallow subsurface of the seafloor (Berner, 1980): −

2þ

2HCO3 þ 2Ca −> CaCO3 þ CO2 þ H2 O

ð3Þ

Note that CH2O is a simplified, conventional, manner to represent sedimentary organic matter as a fraction of C6H12O6, i.e., the most reactive part of organic matter. More complex formulae exist to represent organic matter but they are useless for our purpose. The conditions favoring carbonate precipitation are all the more met that iron reduction is involved (Curtis and Coleman, 1986; Curtis, 1987; Peckmann and Thiel, 2004) as exemplified by Eq. 4: 2−

þ

−

4CH4 þ CH2 O þ 4SO4 þ 4FeOOH þ 3H −> 5HCO3 þ 4FeS þ 10H2 O ð4Þ

In this case, H + ion consumption is combined to HCO3− generation, which strongly impacts carbonate saturation and favors CaCO3 precipitation. To some extent the CaCO3 isotopic signature can be used to identify the formation pathway (Eq. (1) or Eq. (2))(e.g. Bojanowski, 2012). 3. Material The Bancs Jumeaux Formation is located at the boundary between the Argiles de la Crèche Formation and the Argiles de Wimereux Formation (Fig. 2; Proust et al., 1995; Williams et al., 2001). It is bound at its base and top by two sharp erosional surfaces characterized by accumulations of phosphatic fossils and nodules, and quartz pebbles. These lower and upper P-rich levels are named P1 and P2, respectively. The top of the Argiles de la Crèche Formation (i.e., the P1 horizon) is dated as belonging to the Wheatleyensis zone, and the base of the Argiles de Wimereux Formation (P2 horizon) to the

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The sedimentary organic matter of the Bancs Jumeaux Formation was studied by Tribovillard et al. (2004a). The organic content (in the range 0.1–3.0 wt.%) is largely dominated by a brown-colored amorphous organic matter that is usually interpreted as bacterially reworked from marine phytoplankton. The palynofacies also contains minor amounts of gelified and black opaque terrestrial debris, as well as algal debris and algae. Recently, Tribovillard et al. (2008) observed the occurrence of abundant, large-sized pyrite framboids and polyframboids in the marlstones of the Bancs Jumeaux Formation. Based on the geochemical composition (major, minor and redox-proxying trace metals) of bulk sediments and of polyframboids, they concluded that the Bancs Jumeaux Formation experienced dominantly dysoxic depositional conditions (oxygenated water column and redoxcline lying below the sediment–water interface) and that the pyrite polyframboids formed during early diagenesis as the probable results of sulfate reducing reactions occurring in micro-niches (fecal pellets?). The trace metal concentrations published in Tribovillard et al. (2008) will be used in the present paper.

Pallasioides zone (Geyssant et al., 1993; Herbin et al., 1995). The Pectinatus zone has been recently documented in the upper part of the Bancs Jumeaux Formation (Williams et al., 2001). The Bancs Jumeaux Formation has been interpreted as a transgressive systems tract (Proust et al., 1995; Deconinck et al., 1996; Williams et al., 2001). Figs. 1B and 2 illustrates the Bancs Jumeaux Formation at WimereuxNorth (three limestone beds). The three beds can be observed for about 300 m alongshore (01°36′20.7″ E and 50°46′34.4″ N, or x = 548,336 and y = 2,642,736 Lambert IIe). The formation is about 2 m thick (distance separating P1 and P2). The marls are bioturbated, they contain frequent, scattered, wood fragments and bivalve or brachiopod shells (sometimes phosphatic). The lower and upper limestone beds are about 15 cm thick with clear-cut and horizontal boundaries. The intermediate limestone bed is thicker (15–25 cm) and more irregular: although laterally continuous, it has a somewhat nodular structure. The basal two centimeters of the beds are impregnated with the detrital material that makes up the “background” marl sediments, but the tops of the beds are “clean”, that is, show no such impregnation (Fig. 3A). The beds show a light grey color and some shell fragments are scattered throughout. Bioturbation is evidenced by some mottling and, in one sample of the intermediate bed, the presence of a large burrow (Fig. 4A). The burrow displays neat edges and its infilling is rich in shell fragments, quartz grains and glauconite grains compared to the “host” limestone. At Wimereux-South (01°35′59.8″ E and 50°45′13.2″ N, or x = 547896 and y = 2,640,214 Lambert IIe), the Bancs Jumeaux Formation contains only two beds. They are a bit thicker than at Wimereux-North (20 cm) and present sharp boundaries (Fig. 1A and C). They can be observed in the cliff along a distance of about 800 m. The field conditions do not allow us to observe the contact zone between both outcrops (the one with two and the one with three beds), because of the presence of the city Wimereux, which is bordered by two normal faults.

A

4. Methods We characterized the limestone beds and their caging marlstones using petrographical and chemical methods, including the determination of the isotope composition for C, O and S. At both sections the Bancs Jumeaux Formation was sampled every 10 cm, starting below P1 and ending above P2. Thin sections of the limestone beds were observed using an optical microscope, a cathodoluminescenceequipped microscope (Olympus BX41 with a Citl 8200 MK4 coldcathode cathodoluminescence device operating at 20 kV) and a scanning electron microscope (FEI Quanta 200 Environmental) equipped with a backscattered electron device and an energy dispersive spectroscopy-probe (X-Flash 3001 Brucker + Quantax 400 software).

B

C

0.5 mm

Fig. 3. A: Section of the lower bed of Wimereux-North showing that the basal part of the bed is impregnated with clastic sediment. The vertical isotopic transect was carried out on this sample with a sampling step of 2 cm. B: Microscope observation of the microspar matrix typical of the carbonate beds of the Bancs Jumeaux Formation. In the center of the picture, a pyrite framboid is surrounded by a fringe of coarser spar crystals, as if often the case in these carbonate beds. C: SEM observation (back-scattered electron mode) of the surface of a thin section showing the microspar (dark-colored and some pyrite (light colored) crystals. The arrow points to a cluster of tiny pyrite crystals. High resolution photographs of the pictures may be supplied upon request.

N. Tribovillard et al. / Chemical Geology 300-301 (2012) 1–13

5

B

A

C

D Fig. 4. A: Vertical burrow in the intermediate bed of Wimereux-North, containing more abundant shell fragments than the caging carbonate matrix. The mottled fabric of the limestone indicates bioturbation by infauna. B and C: SEM observation of pyrite polyframboids typical of what can be observed in great abundance in the marlstones of the Bancs Jumeaux Formation. D: Example of small synsedimentary faults that affected the study area through Late Jurassic. The arrow points to two well-visible faults. High resolution photographs of the pictures may be supplied upon request.

The grainsize of the HCl-insoluble fraction of the marls and limestone beds was determined using a Malvern Mastersizer Hydro 2000-G apparatus following the protocol described in Trentesaux et al. (2001). The decarbonated fraction of the sediment was sieved at 2 mm prior to grain-size analysis. The carbonate content was determined with a Bernard-type calcimeter (acid digestion followed by CO2 volume determination; accuracy b 5%). The carbon and sulfur concentration of the rocks was determined using a LECO C-S 125 apparatus (accuracy b 5% as shown by in-house intercalibration with a Thermo CHNS analyzer Flash EA 1112 Series). For stable isotope analysis, the samples were grounded in an agate mortar and sieved so as to ensure a grain size lower than 140 μm. CO2 was extracted from calcite by dissolution with 100% H3PO4 (McCrea, 1950) at 25 °C for 12 h in helium flushed Labco Exetainer® vials. Carbon and oxygen isotopic compositions of the evolved CO2 were measured using a gas chromatograph coupled to an isotope ratio mass spectrometer (GC-IRMS)(Analytical Precision 2003, today entitled GV 2003, provided by GV Instruments), helium being the carrier gas. Three internal standards were used to calibrate the δ 13Csample/ref data provided by the GC-IRMS relative to the PDB scale (Table 1). These standards have been calibrated relative to the PDB using two

international standards, NBS19 and IAEACO1 (IAEA catalogue). Results are given in the usual δ-notation relative to the international Standards PDB for the δ 13C and for the δ 18O. The external reproducibility for δ 13C and δ 18O measurements is of 0.1‰ and 0.2‰ respectively (1σ). Each sample was measured twice and the average of two analyses is reported in the tables. For pyrite (poly-) framboid extraction, samples were crushed to centimeter-size lumps that were decarbonated using HCl. Once deflocculated after several rinsing, the suspended clay minerals were removed from the beakers by removing the supernatant (the operation must be repeated several times). The residue was then treated with HF to solubilize quartz grains and residual aluminosilicate grains. After these operations, only clean pyrite particles and Table 1 Carbon isotope of calcite standard. Standard name

δ13C/PDB

δ18O/PDB

Rennes 2 ACROSS Merck

− 9.76 ± 0.03‰ 0.26 ± 0.04‰ − 8.65 ± 0.02‰

− 9.09 ± 0.12‰ 4.19 ± 0.18‰ − 6.89 ± 0.12‰

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N. Tribovillard et al. / Chemical Geology 300-301 (2012) 1–13

some heavy minerals were still present. To determine the sulfur isotope composition (δ 34S), the pyrite samples were oxidized with O2 at 1050 °C to produce SO2 that was analyzed using a VG Sira mass spectrometer. The results were expressed in δ conventional notation, relative to the standard V-CDT (Vienna Canon Diablo Troilite). The standard used is a Ag2Fe MERCK with δ 34S = +3‰ vs. V-CDT. Each sample was measured twice and the average of two analyses is reported. The analytical precision of measurements is ± 0.3‰; the reproducibility is of ± 1‰. 5. Results 5.1. General features of the Bancs Jumeaux Formation The carbonate content varies between 13 and 90% throughout the formation. Carbonate beds contrast sharply with marls, showing CaCO3 content above 80% whereas marls yield values between 13 and 68% (mean 40%; Table 2, Fig. 2). The measured values of total C and CaCO3 contents allow to derive the organic C content that varies between 0.1 and 3.0%, being the lowest in carbonate beds (Table 2).

The sulfur content is high in marls: between 0.1 and 3.3% (mean 1.1%), being again lower in carbonate beds (Table 2). Considering bulk rock composition, limestone beds show a S content (that is considered to be a proxy to pyrite content) much lower than that marls (Table 2). We recalculated the S content on a carbonate-free-basis in order to compare the pyrite content of marls to that of limestone beds when the “diluting” carbonate fraction is removed. The results (Table 2) show that the carbonate beds fall within the range of values of the marls. The C and O isotopic composition also shows contrasting values between limestone beds and marls. Limestone beds show lower δ 13C values than marls but, for δ 18O, the pattern is not so neat, the values of limestone beds being among the most positive (Table 2, Fig. 5A). Uranium and molybdenum enrichments may be assessed through the calculation of their respective enrichment factors (EF) defined as following: X-EF = [(X/Al)sample/(X/Al)Crust], where X and Al represent the weight% concentrations of element X and Al, respectively. Samples were normalized using the Earth's upper crust (Crust) compositions of McLennan (2001). Any EF value larger than 1.0 will point to

Table 2 Geochemical data and mode of the grain-size distribution of the studied samples. CFB stands for carbonate-free basis. The shaded lines represent the limestone beds. Location

Sample #

Height (cm)

CaCO3 (wt.%)

Mode (μm)

Organic C (wt.%)

S (wt.%)

S CFB (wt.%)

Wimereux South

BJ01 BJ02 BJ03 BJ04 BJ05 BJ06 BJ07 BJ08 BJ09 BJ10 BJ11 BJ12 BJ13 BJ14 BJ15 BJ16 BJ17 BJ18 BJ19 BJ20 BJ21 W01 W02 W03 W04 W05 W06 W07 W08 W09 W10 W11 W12 W13 W14 W15 W16 W17 W18 W19 base

0 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 0 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 0 2 4 6 8 10 12

26.9 20.9 27.5 53.5 41.0 40.3 56.9 65.7 64.4 88.9 50.7 45.2 56.5 42.1 43.1 68.4 88.1 90.4 42.9 33.2 29.6 22.9 23.0 27.5 28.6 15.7 13.0 44.9 80.6 27.7 56.0 90.4 47.4 46.2 39.9 55.8 83.6 26.4 46.4 43.3 77.2 79.0 79.3 78.7 78.6 83.8 89.8 89.4 88.4

11.4 12.9 10.9 6.2 5.8 6.0 5.8 5.7 5.9 5.9 6.2 5.4 6.4 6.1 7.1 8.1 7.5 7.6 5.8 4.9 5.0 14.0 13.1 9.8 12.4 12.6 14.1 9.0 7.5 7.4 5.4 6.0 5.4 5.9 5.9 6.5 6.0 5.5 4.9 4.9

1.5 2.7 1.9 1.9 1.6 1.8 0.9 0.3 0.6 0.5 0.8 0.9 0.1 1.4 1.6 0.7 0.3 0.0 0.0 0.3 2.6 2.0 1.5 1.6 1.6 2.4 2.6 0.8 0.7 3.0 0.4 0.0 0.3 0.5 1.3 0.9 0.5 1.2 0.6 0.0

1.9 2.0 2.4 5.9 3.1 3.5 2.8 2.6 1.0 0.2 1.9 2.3 2.9 2.4 2.1 2.1 0.4 0.7 2.5 3.2 4.8 1.6 2.2 1.2 1.1 2.0 1.9 7.2 0.6 4.7 1.7 0.4 3.5 2.3 2.9 2.9 0.9 3.4 5.1 3.5

2.6 2.5 3.2 12.6 5.2 5.8 6.6 7.6 2.7 1.8 3.9 4.1 6.7 4.2 3.7 6.6 3.4 7.8 4.3 4.8 6.8 2.1 2.8 1.7 1.5 2.4 2.2 13.1 3.3 6.6 3.8 4.5 6.6 4.3 4.8 6.5 5.6 4.7 9.6 6.2

Wimereux North

Transect through the lower bed at Wimereux-N.

top Medium bed at inner burrow Wimereux-N carbonate matrix

δ13C ‰

δ18O ‰

0.69

− 5.01

0.85

− 5.78

0.94 0.71 − 1.7 0.55 0.16 − 2.56 0.7 0.63

− 4.45 − 4.37 − 2.92 − 3.58 − 3.61 − 3.24 − 3.07 − 2.87

0.22 − 1.07 1.33 1.32 1.34 − 3.98 − 2.69 − 1.21 − 0.37 − 0.30 − 0.29 − 1.30 − 1.41 − 3.49

− 3.54 − 2.76 − 3.3 − 3.87 − 3.85 − 4.16 − 3.13 − 2.57 − 2.38 − 2.46 − 2.31 − 2.65 − 2.64 − 3.84

N. Tribovillard et al. / Chemical Geology 300-301 (2012) 1–13

enrichment of an element relative to its average crustal abundance. In practical terms, EFs >3 represent a detectable enrichment of an element over average crustal concentrations, and EFs >10 represent a moderate to strong degree of enrichment (Algeo and Tribovillard, 2009). The analytical data published by Tribovillard et al. (2008) all fall in the area of the U-EF vs. Mo-EF diagram of Algeo and Tribovillard (2009) corresponding to oxic to suboxic conditions of deposition (Fig. 6), which corroborates the earlier conclusions of Tribovillard et al. (2008).

A

7

5.2. Pyrite (poly-) framboids in the marls All marl samples contain pyrite framboids and polyframboids that reach unusually large dimensions. For instance, polyframboids may be as large as 300 μm. SEM observation reveals that, within the same marl samples, many pyrite morphologies may be encountered, regardless of the sample stratigraphic position, or of its chemical composition (Fig. 4B and C). A common feature is that most often individual pyrite crystals observed in the (poly-) framboids are pitted,

200

180

160

Stratigraphy (cm)

140

120 25 50

100

100

90

80

80

60 70 40

20

δ 18O (‰) δ 13C (‰) CaCO3 (%)

δ 18O (‰) δ 13C (‰)

-6

-4

-2

1

2

60

0

2

50

0 -6

-4

-2

0

2

δ (‰) vs PDB δ13C

B

-5

-4

-3

-2

-2,0

-1

0

of ed Top wer b lo the

-2,5 -3,0

-4,0 -4,5 -5,0 -5,5 -6,0

B th ase e lo of w er be d

δ18O

-3,5

lower bed intermediary bed upperbed caging marlstones

Fig. 5. Stratigraphic distribution of the carbon and oxygen isotope data for the Wimereux-North Section (A). The shaded zones represent the three limestone beds. B: crossplot of the isotope data showing that the limestone beds strongly differ from the background sediment (marlstones). The figure also illustrates a transect through the lower bed of Wimereux-North. The sampling step is 2 cm. A clear trend is drawn from the base to the top of the bed (except for the sample at the very top of the bed).

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N. Tribovillard et al. / Chemical Geology 300-301 (2012) 1–13

Mo EF

Bancs Jumeaux

0.3 x SW

SW

100

A

1000

Mo-EF

open marine conditions

B

S U)

W

o/ (M

10

100

ic

lfid

W

10

particulate shuttle

su

ic ox

to

ic ox an

)S

U o/

3

0.

M x(

b

su to c ic ox boxi su

U-EF 1 1 1

1

10

U EF

10

100

100

Fig. 6. Crossplot opposing the respective enrichment factors (EF; see Section 5.1) of uranium and molybdenum for the Bancs Jumeaux (A). This diagram allows reconstructing the paleo-redox conditions of deposition (Algeo and Tribovillard, 2009). The insert B corresponds to the situation of deposition under open marine conditions (see discussion in Algeo and Tribovillard, 2009). The diagonal lines represent multiples (0.3 × and 1×) of the Mo:U ratio of present-day seawater: molar ratios of ~ 7.5 for the Pacific and ~ 7.9 for the Atlantic have been converted to an average weight ratio of 3.1 for the purpose of comparison with sediment Mo:U weight ratios.

as if some previously included things had dissolved after pyrite formation (Fig. 4C). In carbonate beds, pyrite framboids are smaller than in marls and no polyframboids are observed. 5.3. Grainsize of the terrigenous fraction of the sediment Only the b 2 mm fraction of the HCl-insoluble part of the sediment is studied when using a laser-equipped grainsizer. In other words, the rather commonly encountered shell or wood fragments and the phosphate or quartz pebbles that characterize the P1 and P2 levels are not taken into consideration. The various parameters of the grain-size analysis may be conveniently represented by the particle size distribution mode (Figs. 2 and 7). The b 2 mm terrigenous fraction is dominated by particles which size varies between clay and fine silt. The main result is the monotonous distribution of the grain size parameters observed for the Bancs Jumeaux Formation at the studied two sections, whatever the lithofacies, i.e. limestone or marl. Some variability between the marl samples is accounted for by the presence of more or less abundant pyrite (poly-) framboids (Fig. 7). 5.4. Specific features of limestone beds 5.4.1. Microfacies observation of limestone beds All 3 + 2 limestone beds show the same microfacies features. The limestone is a homogeneous calcite microspar where common shell fragments and rare pyrite framboids are interspersed (Fig. 3B and C). A fringe of microsparitic crystals is observed around most pyrite framboids (Fig. 3B). It is emphasized that no coccolith fossils (or molds) could be observed, contrary to what is reported for some carbonate beds of the Kimmeridge Clay Formation (e.g., Lees et al., 2004). Polyframboids are a typical feature of marlstones, not of limestone in the case of the Bancs Jumeaux Formation. Bioturbation is

visible through the presence of mottling. When examining the burrow illustrated in Fig. 4A, we observe that the microspar crystals are slightly larger in the burrow infilling than in the caging microspar, and that the infilling contains abundant shell fragments and common quartz and glauconite grains. Cathodoluminescence imaging shows that all observed thin sections are homogeneous from a “color” point of view. No luminescence contrast was observed, indicating that the chemical composition of the carbonate is homogeneous. This is still true when burrowed portions are compared to unbioturbated “matrix” portions.

5.4.2. Isotope compositions An analytical transect was carried out from base to top of the lower bed of Wimereux-North with a sampling step of 2 cm. The base of the bed is more impregnated with the background terrigenous fraction typical of marls than the upper part that is clean, i.e., more and more impoverished in terrigenous fraction. This vertical evolution, observed on the field, is evidenced by the CaCO3 content rising from 77.2% at the base of the bed up to 89,8% at the top. Fig. 5B shows that the δ 13C increases from the base of the bed (−3.98‰) upward, reaching −0.29‰ but the very top of the bed shows a value of −1.30‰. The same evolution is observed with the δ 18O, although less smooth. In addition, the studied sample of the intermediate bed of Wimereux-North shows a rather large burrows whose infilling was compared to the carbonate matrix. The infilling shows higher values for both δ 13C (− 1.41‰) and δ 18O (−2.64‰) compared to the carbonate matrix (−3.49‰ and − 3.84‰, respectively; Table 2). For four samples, the pyrite content was extracted from the carbonate matrix following the procedure described in Section 4 and analyzed to determine the S isotope composition. The four pyrite extracts have δ34S of −36.4‰, -29.1‰, -26.9‰ and −35.2‰, respectively.

N. Tribovillard et al. / Chemical Geology 300-301 (2012) 1–13

9

Particle Size Distribution 7

Volume (%)

6

100

W11

S: 0.2 wt% carbonate beds

80

5 60

4 3

40

2 20 1 0 0.01

0.1

1

10

100

1000

0 3000

5.5 5 4.5

Volume (%)

4

100

W7

S: 3.1 wt% pyrite-rich marls

80

3.5 60

3 2.5

40

2 1.5

20

1 0.5 0 0.01 8

Volume (%)

7 6

0.1

1

10

100

1000

0 3000 100

BJ11 S: 0.8 wt% pyrite-poor marls

80

5

60

4 40

3 2

20

1 0 0.01

0.1

1

10

100

1000

0 3000

Particle Size (µm) Fig. 7. Particle size distribution of the sieved, decarbonated, fraction of the sediment. This fraction consequently does not contain the shell or wood fragments, phosphatic clasts or nodules, an quartz pebbles larger than 2 mm encountered in the tempestites of the uppermost part of the Argiles de la Crèche Fm. or in the P1 and P2 centimetric levels of the Bancs Jumeaux Fm. W11 is representative of the terrigenous fraction of the various carbonate beds studied here. W7 is representative of the marls rich in pyrite (poly-) framboids; BJ11 is representative of the pyrite-poor marls.

6. Interpretations The questions we are facing are: 1) what is the origin of the limestone beds - sedimentary carbonate particle accumulation or diagenetic CaCO3 precipitation? – and, 2) in the case where the beds are of diagenetic origin, at which stage(s) of diagenesis did they form (in the organo-diagenesis zone, i.e. in the sulfate-reduction zone or the anaerobic methane oxidation zone, or in the burial diagenesis zone)? and at which depth (ranging from the sediment water interface down to several kilometer depth?). The limestone beds of the Bancs Jumeaux Formation differ markedly from the general style of the Late Jurassic of the Boulonnais area that is characterized by clastic-dominated sediments deposited on a ramp where limestone beds are rare and generally quite rich in large bivalve shells. The fact that the beds are two or three whereas the distance between the two outcrops is only ca. 2 km does not

advocate for the “sedimentation” hypothesis. However, the Kimmeridge Clay Formation (a distal lateral equivalent of the Late Jurassic of Boulonnais) contains calcite-made limestone beds that result from biogenic particle accumulation (such as the White Stone Band or Freshwater Steps Stone Band; e.g., Lees et al., 2004; Pearson et al., 2004), as well as dolomitic cementstone beds (such as the Washing Ledge, occurring within oil shales) and carbonate concretions of diagenetic origin sometimes aligned along beds (e.g., Astin and Scotchman, 1988; Scotchman, 1991). Thus, contrary to what is observed for the limestone beds of the KCF, the limestone beds of the Bancs Jumeaux Formation do not show any coccolith (fossils or molds) and are made of calcite microspar. The microspar may result from neomorphic replacement of a micritic matrix or from precipitation as cement, or even from a direct precipitation under bacterial control (see discussion in, e.g., Folk and Chafetz, 2000; Munnecke et al., 1997, 2001).

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6.1. Early diagenetic origin of the limestone beds The limestone beds are bioturbated. Most of the bioturbation occurred as the beds were still soft, as evidenced by their mottled fabric. However the clear edges of the burrow illustrated in Fig. 4A show that the bed was already firm when the burrower intruded. It indicates that the carbonate mass formed early and began to solidify at shallow depth below the sediment–water interface. In other words, the diagenetic carbonate precipitated during early diagenesis at a time when it could be easily reached and burrowed by infauna. The (micro-) facies observation suggests that the shell fragments and quartz or glauconite grains were injected by bioturbation from the overlying marls down into a carbonate matrix that seems to have been initially pristine. In addition, the color homogeneity of the various limestone beds observed through cathodoluminescence imaging indicates that the beds were formed in a single phase of carbonate precipitation. This observation reinforces the interpretation of the beds forming rapidly with no evidence for subsequent carbonate growth or replacement. The vertical distribution of δ 13C values in the Bancs Jumeaux Formation at Wimereux North shows that carbonate beds differ significantly from marlstones in that they show relatively depleted δ 13C values. The discrepancy between limestone beds and marlstones indicates contrasting respective origins: δ 13C values of marlstones are consistent with those expected for Late Jurassic seawater carbonates, which range from 0 to 2‰ (Veizer et al., 1999; Prokoph et al., 2008) (Fig. 5A and B; Table 2). It means that the bulk of the marl carbonate fraction presents no detectable influence of diagenetic cement that would impact significantly the isotopic signature. In clear contrast, all the δ 13C values of the three limestone beds are slightly yet significantly lower than the marine carbonate values (Fig. 5A and B; Table 2). It indicates that a significant fraction of the limestone beds carbonate content is of diagenetic origin, which lowers their δ 13C signature. The beds do not show δ 18O values strikingly different from those of the marls but they tend to be less depleted in 18O, or in other words they are closer to marine carbonate values (Fig. 5A and B). The fact that the δ 13C are not strongly negative suggests that limestone beds carbonates resulted from two mixing sources: 1) a light C source such as organic matter remineralization and, 2) a seawater dissolved inorganic carbonate source. Consequently, it is inferred that all limestone beds were impacted by diagenesis (organic-matter remineralization) during deposition (influence of seawater on the isotopic signature of C and O). Thus the limestone bed formation occurred during what we may call synsedimentary early diagenesis. The δ 34S values for the pyrites extracted from the Bancs Jumeaux range between −29.4‰ and −36.4‰, to be compared to the marine value for the Late Jurassic of about +15‰ (Prokoph et al., 2008). The marked sulfur isotope fractionation between the seawater values and pyrite (between 45 and 50‰) strongly suggests that bacterial sulfate reduction operated within a comparatively infinite sulfate reservoir, that is, at or very close to the sediment–water interface among anaerobic processes, sulfate reduction is most commonly involved into diagenetic carbonate precipitation because sulfate ions are abundant in seawater and sediment pore water (Jørgensen, 1982; Scotchman, 1991; Peckmann and Thiel, 2004). Here the high S content observed for marlstones as well as limestones (after recalculation on a carbonate-free basis), together with the direct observation of pyrite and of course the low δ 34S values of pyrite advocates for a prominent role played by sulfate reduction reactions at the sediment–water interface. Consequently, in what follows, we shall speak of sulfate reduction only, but it does not exclude the participation of denitrification or Fe–Mn oxy-hydroxide reduction. This is all the more true that the sedimentation rate must have been largely low during the deposition of the Bancs Jumeaux Formation that corresponds to a transgressive systems tract (see Section 3): under low sedimentation rate conditions, the sediment may be enriched in Feand Mn- (oxyhydr-) oxides.

This conclusion, that sulfate reaction operated at the sediment– water interface or in its closest vicinity is corroborated by the isotopic transect carried out through the basal bed of Wimereux-North. From base to top of the carbonate bed, the δ 13C values are less and less 13C-depleted and thus closer and closer to marls and thus to seawater carbonate values. This observation alone also strongly suggests that the conditions leading to carbonate supersaturation and CaCO3 precipitation were met at the boundary between the sulfate reducing zone and the water column, namely at or close to the sediment water interface. Our interpretation is that the carbonate precipitation started at the sediment–water interface with the incorporation of light C isotopes released by organic-matter degradation and then the bed grew upward, incorporating less and less organic matterderived C, hence more and more seawater-derived C. This scenario accounts for the fact that the base of the bed is impregnated with the terrigenous material where it began to grow, and the carbonate is purer and purer upward. The upper part of the bed is almost devoid of terrigenous material because growth must have been quick and the sedimentation rate was probably low (see above). Within the middle bed at Wimereux-North, the carbonate infilling of the studied burrow and the caging carbonate matrix show contrasting isotopic signature (Table 2), the infilling being closer to marine carbonate values. It suggests that the burrower injected “top-down” a material that was more in contact with seawater than the burrowed carbonate matrix. This view is consistent with the scheme in which the limestone bed formed at the sediment–water interface. 6.2. Sulfate reduction reactions at the sediment water interface In the genetic scheme that we put forward to account for the synsedimentary early-diagenetic formation of the carbonate beds, sulfate-reduction reactions are interpreted to operate at the sediment–water interface or just below it. Previous works concluded that largely the Bancs Jumeaux Formation did not experiment anoxic conditions of deposition (Tribovillard et al., 2004, 2008), which is also the conclusion from Fig. 6. In other words, overall depositional conditions were not prone to the operating of sulfate reduction reactions at the sediment–water interface, and the conditions prone to limestone bed formation were met only in limited occurrences. What could favor sulfate reduction reactions at or just below the sediment– water interface? The “fuel” for bacterially-mediated sulfate reduction reactions during early diagenesis (Eq. 1) are the sulfate ion pools and labile organic matter, in the absence of free, dissolved oxygen. During early (especially earliest) diagenesis, the availability of sulfate is not a limiting factor in marine environments. The necessary absence of dissolved oxygen may result from a strong increase of aerobic degradation of organic matter, which exhausts the free-O2 and raises the redox-cline closer to the sediment–water interface, hence launching sulfate reduction (case 1); it may also result from the advection of O2-poor or even O2-devoid bottom water (case 2). In the case of marked bursts of sulfate-reduction induced by an increase of the organic matter flux to the sediment (case 1), an expected side effect would be the formation of abundant pyrite in the case of non-limiting iron availability. When calculations on a carbonate-free basis are considered, the limestone beds are not richer in pyrite than the marls. In addition, no pulses of organic-matter flux are echoed in the vertical distribution of the organic-carbon content. Although pyrite or organic carbon do not testify to episodic bursts of sulfate reduction, this scenario cannot be discounted; however, a less hypothesis-demanding scenario is envisioned (case 2), which also accounts for the lateral variability of the number of the carbonate beds. In this second scheme, if the sea bottom is bathed by oxygenpoor bottom water preventing aerobic organic-matter degradation, sulfate reduction reactions can produce at the sediment–water interface or at least close to it in the absence of changes in the flux of organic matter. Under these circumstances, H2S released through

N. Tribovillard et al. / Chemical Geology 300-301 (2012) 1–13

Eq. 1 would not accumulate at the sea bottom (or just below the sediment–water interface) but could instead diffuse higher up into the water column. H2S is a weak acid opposing to carbonate supersaturation. D. Burdige (2006) explains that, when H2S is eliminated (for instance, as suggested above by diffusion into the water column, or through reactions with iron species; Boudreau and Canfield, 1988; Canfield and Raiswell, 1991), carbonate supersaturation can be reached quite rapidly even if sulfate reduction rates are not high. In other words, if we envision that the sea bottom had been occasionally bathed by oxygen-poor (-deprived) bottom waters, we suggest that sulfate reduction could begin at the sediment–water interface or immediately below it, fueled by dissolved sulfate and labile organic matter (which is present throughout the Bancs Jumeaux Formation). Owing to the easy upward diffusion of H2S, carbonate supersaturation and precipitation are attained at the sediment–water interface, with no intense sulfate reduction reactions required. Thus, a limestone cement bed could form rapidly and, once formed, get burrowed by bottom dwellers, once normal conditions have been reestablished. It is noteworthy that alternate periods of oxic and anoxic seafloor conditions have been documented for the late Kimmerdigian–early Tithonian Argiles de Châtillon Formation lying under the Argiles de la Crèche Formation (Proust et al., 1995) or for the Kimmeridge Clay Formation (Oschmann, 1988). These alternations were expressed through mass mortality events of juvenile faunas possibly linked to strong seasonality or to variations in marine currents. The fact that the bed formed at the sediment–water interface (or close to it) explains why the bed is laterally continuous and not just an alignment of disconnected concretions as is usually the case when concretions grow during later diagenesis at some distance below the sediment–water interface (e.g., Selles-Martinez, 1996). The intermediate bed at Wimereux-North is more nodular than the other two beds, although continuous. Two causes may be put forward: 1) bioturbation was more intense and favored nodularization during later diagenesis; 2) the intermediate bed formed at shallow depth below the sediment–water interface, whereas the other two beds formed at the sediment–water interface, and this shallow burial induced nodularization. If we adopt the scenario exposed above, then the questions are the following: primarily what caused the presence of oxygen-restricted bottom water? Secondarily, why two or three beds are observed, depending on the site studied? 6.3. The possible role of synsedimentary faults The two questions are most probably tied together and a common explanation may be suggested. The fact that limestone beds must have developed at the sediment–water interface rather than deeper in the sediment, strongly suggests that O2-restricted waters spilled over the sea floor of the depositional setting. However, in what was a ramp environment, there is no compelling explanation for such a spill. During the entire course of the Late Jurassic, the Boulonnais area was affected by synsedimentary faulting in relation with the rifting of the Atlantic Ocean (Mansy et al., 2003). The synsedimentary fault movements is observed on the field (Fig. 4D). The idea put forward here is that synsedimentary fault movements could trigger the release of porewaters through the fault planes into the bottom water mass (Fig. 8). The pore waters thus released will be anoxic because originating from horizons lying below the redox-cline. In addition, these porewaters are enriched in dissolved chemical species, thence denser than normal seawater. In particular the pore waters will also be enriched in dissolved inorganic carbon originating from organic matter degradation, and thus of lower δ 13C than seawater. The pore waters suddenly released will spill over the sea bottom and will not mix immediately to the overlying seawater. This spill of oxygen-lean water will thus promote sulfate reduction reactions coupled with easy upward diffusion of H2S into the free seawater,

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therefore triggering synsedimentary early-diagenetic formation of a limestone bed, but this situation will be limited in duration. Nevertheless, this situation may recur with recurrent fault movements. Consequently, several beds may be observed. If the depositional setting is affected by several synsedimentary faults and thereby compartmented in blocks that may evolve differently according to differential fault movements, it explains how neighboring compartments can undergo two times versus three times such recurrent situations, inducing the formation of two versus three limestone beds (Fig. 8). An additional factor is necessary in the model presented here: the hydrodynamic level must be absolutely low on the sea floor. In case of water mixing on the sea floor (wave action, bottom currents), alkalinity cannot reach the critical threshold allowing carbonate ion supersaturation. In addition, if a diagenetic bed forms at the contact between sediments and seawater, it cannot keep its straight boundaries if the hydrodynamic level is not low. In the present case, as exposed above, the Bancs Jumeaux Formation, comprised between the P1 and P2 levels, is interpreted as a transgressive systems tract, as well as an episode of high water depth in the Boulonnais area (Williams et al., 2001). Such depositional conditions, coupled to the fact that the clastic fraction of the sediment is fine-grained, advocates for a low level of hydrodynamism (below storm wave base). 7. Conclusion Early diagenetic carbonate beds resulting from sulfate reduction reactions within the sediment are rather common; however such beds are usually alignments of carbonate nodules, more or less closely spaced laterally. In this study, we present a rare case of episodic laterally-continuous beds formed by early diagenetic processes at the sediment–water interface. Such diagenetic beds are extremely rarely reported in the literature, in contrast to (totally or partially) diagenetic beds of limestone-marls alternations which result from self-organization processes, i.e., a carbonate redistribution during diagenesis leading to a regular alternation of beds and interbeds. In the latter case, carbonate beds are present in great numbers (e.g., Munnecke et al., 2001). In the present case, the diagenetic carbonate beds are strikingly resembling beds formed by particle settling because they must have formed at (or close to) the sediment–water interface with no strong pressure constraint prior to (even shallow) burial. Their formation was possible because sulfate reduction reactions possibly operated in the contact zone between the base of the water mass and the sediments. The magnitude and duration of the phenomenon were not sufficient to induce the development of euxinic conditions. The sulfate reduction reactions at the sediment–water interface were triggered by spills of anoxic waters on the sea floor. This sudden lack of dissolved O2, coupled to the presence of relatively abundant labile organic matter, stimulated sulfate reduction. Usually, when sulfate reduction operates within sediment pore waters, the increase in alkalinity is counter-balanced by the release of H2S that opposes to carbonate supersaturation. Here, H2S was able to diffuse quickly into the water column and did not neutralize the alkalinity rise. Such situations where sulfate reduction can operate at the sediment–water interface are rather common at the geological scale. In the present study, we conclude that the key factor was spills of anoxic waters released by synsedimentary fault movements; in other occasions, the key factors are bottom-water stagnation and/or intense organic matter degradation exhausting bottom waters from their dissolved oxygen. Nevertheless, whatever common such situations may be at the geological scale, carbonate beds such as those reported in this paper are little (or even not) reported in the literature. Our conclusion is that diagenetic beds such as those of the Bancs Jumeaux Formation must have been largely overlooked and that they are most probably much more frequent than usually

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N. Tribovillard et al. / Chemical Geology 300-301 (2012) 1–13

Sea level Sea floor

1

Initial stage of the ramp, affected by synsedimentary faulting

2

Normal movements of the faults and dense, oxygendeprived pore waters spill over the sea floor

3

Formation of a synsedimentary limestone bed; sedimentation continues

4

Normal fault movements in a sub-compartment; spill of pore waters

5

Formation of a second synsedimentary carbonate bed; sedimentation continues

Normal fault movements; spill of pore waters

6

Formation of a third synsedimentary carbonate bed

7 Wimereux-North

Wimereux-South

Fig. 8. Schematic scenario of the synsedimentary early-diagenetic formation of the limestone beds, accounting for the fact that two beds are observed at Wimereux-South and three beds crop out at Wimereux-North. The model is scaleless and does not consider any role played by the sea level. The precise location of the synsedimentary faults is not known but their presence is demonstrated by the geometry of the sediment bodies and field evidences.

thought, especially further back in time, during transitions from anoxic to oxic bottom waters in the oceans, when the chemocline must have resided at the sediment–water interface quite often. This might be the case in the Late Neoproterozoic times for which strong lateral variations of the δ 13C would be easily accounted for by such carbonate precipitation occurring in the bottom waters (for instance the Ediacarian Doushantuo Formation of South China; Ader et al., 2009).

Acknowledgements We thank the technical staff of the Géosystèmes Lab (Université Lille 1) for their collaboration: Laurence Debeauvais, Lea-Marie Emaille, Philippe Recourt, Sylvie Regnier, Sandra Ventalon. The two undergraduate students Joelle Dimangou and Cyril Berthome are acknowledged for their appreciated and stimulating cooperation. We thank Uwe Brand (Editor-in-Chief of Chemical Geology), Jean-

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