Methane-derived authigenic carbonates along the North Anatolian fault system in the Sea of Marmara (Turkey)

Deep-Sea Research I 66 (2012) 114–130

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Methane-derived authigenic carbonates along the North Anatolian fault system in the Sea of Marmara (Turkey) Antoine Cre´mie re a,n, Catherine Pierre a, Marie-Madeleine Blanc-Valleron b, Tiphaine Zitter c, M. Namik C - a˘gatay d, Pierre Henry c a Laboratoire d’Oce´anographie et du Climat: Expe´rimentation et Approches Nume´riques (LOCEAN), UMR 7159, Universite´ Pierre et Marie Curie, 4 Place Jussieu, 75252, Paris Cedex 05, France b Muse´um National d’Histoire Naturelle, UMR 7207 - CR2P, 43 rue Buffon, 75005 Paris, France c CEREGE, Colle ge de France, Chaire de Ge´odynamique, UMR 6635, Aix-Marseille Universite´, Europˆ ole de l’Arbois, Bat. Trocade´ro, BP80, 13545 Aix-en-Provence Cedex 04, France d Faculty of Mines, Geology Department, Istanbul Technical University, Ayazaga 80626, Istanbul, Turkey

a r t i c l e i n f o

a b s t r a c t

Article history: Received 18 November 2011 Received in revised form 23 March 2012 Accepted 4 April 2012 Available online 14 April 2012

The Marnaut cruise (May–June 2007) investigated the submerged part of the North Anatolian fault system, an active tectonic area in the Sea of Marmara. Already known and new fluid venting sites along the fault system were visited by submersible diving. Cold seeps present a considerable diversity of geochemical background associated with occurrences of authigenic carbonate crusts outcropping at the seafloor. Buried carbonate concretions were also recovered by coring within the sediments of the Tekirda˘g Basin and of the Western-High ridge that separates the Tekirda˘g and Central Basins. Interestingly, numerous of these early diagenetic carbonates were found within the transitional sediments from lacustrine to marine environment deposited after the late glacial maximum. The authigenic carbonates are mainly composed of aragonite, Mg-calcite and minor amounts of dolomite, and are often associated with pyrite and barite. The carbon isotopic compositions of carbonates present a wide range of values from  50.6% to þ 14.2% V-PDB indicating different diagenetic settings and complex mixtures of dissolved inorganic carbon from different sources. The low d13C values of the seafloor crusts and of most buried concretions indicate that the carbon source was a mixture of microbial and thermogenic methane and possibly other hydrocarbons that were oxidized by anaerobic microbial processes. The positive d13C values of a few buried concretions from the Western-High ridge reflect the mineralization of heavy CO2, which is thought to represent the residual by-product of oil biodegradation in a subsurface petroleum reservoir that migrated up with brines. Most of the oxygen isotopic compositions of seafloor carbonates are close to the isotopic equilibrium with the present-day bottom water conditions but a few values as low as  1.9% V-PDB indicate precipitation from brackish waters. In buried carbonate concretions, d18O values as high as þ 4.9% V-PDB reflect the contribution of water enriched in 18O. The results support the hypothesis that after the late glacial/Holocene transition, precipitation of authigenic carbonates, now buried within the sediments of the Western-High mound structures, was promoted due to enhancement of anaerobic oxidation of methane, possibly from massive methane release by gas hydrate dissociation, and by sulfate rich Mediterranean water incursion. & 2012 Elsevier Ltd. All rights reserved.

Keywords: Methane-derived authigenic carbonates Marmara Sea Cold seeps North Anatolian fault Biogenic and thermogenic methane Gas hydrates Carbon and oxygen isotopes

1. Introduction Natural escapes of fluids and gas from seafloor margins are widespread manifestations called ‘‘cold seeps’’ found all over the world’s oceans (Campbell, 2006; Judd and Hovland, 2007; Tyler et al., 2003) in specific geological environments such as convergent plate boundaries (Kulm et al., 1986; Sibuet et al., 1988),

n

Corresponding author. E-mail address: [email protected] (A. Cre´mie re).

0967-0637/$ - see front matter & 2012 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/j.dsr.2012.03.014

passive continental margins (Kennicutt et al., 1985; Paull et al., 1984) and active fault systems (Moore et al., 1990) that allow ascending migration of gas-rich fluids throughout the sediments. Typically, these vents release fluids rich in methane and other hydrocarbons that are produced by the degradation of organic matter by microbial or thermogenic processes (Whiticar, 1999). Seafloor discharge of these fluids enriched in reduced compounds, such as methane and hydrogen sulfide, feeds chemosynthetic benthic communities (Sassen et al., 1993; Sibuet and Olu, 1998). Methane is mostly consumed within sediments by the anaerobic oxidation of methane (AOM) mediated by a microbial consortium

A. Cre´mie re et al. / Deep-Sea Research I 66 (2012) 114–130

that oxidizes methane generally via sulfate reduction and produces hydrogen sulfide (Boetius et al., 2000; Orphan et al., 2001; Reeburgh, 1976; Valentine and Reeburgh, 2000).   CH4 þSO2 4 -HCO3 þHS þ H2 O

This reaction occurs mainly at the sulfate-methane transition zone (SMTZ) between the upward advecting methane-rich fluid and sulfate-rich bottom seawater that either diffuses or is advected through the sediment (Borowski et al., 2000; Borowski et al., 1996; Wallmann et al., 1997). Production of bicarbonate ions increases pore water alkalinity, which combined with dissolved Ca2 þ and other divalent cations promotes precipitation of authigenic carbonates in shallow subsurface sediment according to the reaction: 2þ 2HCO -CaCO3 þ CO2 þH2 O 3 þ Ca

Microbial and thermogenic methane is characterized by low d13C values ranging from  110% to  50%, and  50% to  20%, respectively (Whiticar, 1999). The mineralization of oxidized methane is recorded in carbonates presenting low d13C values (Aloisi et al., 2000; Bahr et al., 2009; Feng et al., 2010; Gontharet et al., 2007; Kulm et al., 1986; Mazzini et al., 2004; Naehr et al., 2007; Paull et al., 1992; Peckmann et al., 2001; Pierre and Fouquet, 2007). Production of a 13C-rich CO2 pool in methanogenic sediments by organic-matter fermentation or by CO2 reduction contributes to the formation of 13C-rich carbonates (Boehme et al., 1996; Claypool and Threlkeld, 1983; Irwin et al., 1977; Kopf et al., 1995). In a low temperature and high-pressure regime, water and methane can combine to form gas hydrates within the sediments. These crystalline structures consist of cages formed by water molecules entrapping molecules of light gases such as methane, ethane, carbon dioxide or hydrogen sulfide (Kvenvolden, 1993; Sloan, 2003). Methane hydrates contained in the sediments of continental margins represent a large carbon reservoir sensitive to fluctuations of temperature and pressure. Massive episodic releases of methane to the atmosphere from dissociation of marine gas hydrates are thought to have played a significant role in the past climatic system as this gas contributes to increase the greenhouse effect (Kennett et al., 2000; Kvenvolden, 1988; MacDonald, 1990). Past events of huge and abrupt methane release due to decomposition of marine gas hydrates were inferred from low d13C values of diagenetic carbonates as well

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as foraminifers in marine sediments (Bohrmann et al., 1998; Dickens et al., 1995; Garidel-Thoron et al., 2004; Katz et al., 1999; Kennett et al., 2000; Matsumoto, 1989; Pierre et al., 2002). One of the main purpose of the Marnaut cruise (2007) in the Marmara Sea was to understand the relationships between active faults, fluid emissions and biogeochemical processes in the deep-sea sediments (Henry, 2007). Authigenic carbonates were recovered by submersible dives and in piston cores along the North Anatolian fault network. In this study, we report petrographical, mineralogical and scanning electron microscopy (SEM) observations coupled with stable isotopes analysis of diagenetic carbonates. This paper discusses about the variability of the processes that controlled the carbonate diagenesis, in the local context of the North Anatolian fault system. For the buried concretions retrieved down to 4.8 m below the seafloor (mbsf) at the Western-High ridge located between the deep Tekirda˘g and Central Basins, we discuss their occurrence in the stratigraphic record, their possible relationship with the presence of gas hydrates, and we propose a link with the paleoceanographic history of the Sea of Marmara during the late glacial/Holocene transition.

2. Background informations 2.1. Geological setting The Sea of Marmara is an intra-continental basin located in the northwestern part of Turkey between the Mediterranean Sea and the Black Sea. Seafloor morphology (Fig. 1(B)) is characterized by three deep basins (from west to east: Tekirda˘g, Central and C - inarcik Basins) reaching a maximum depth of 1273 m, which are separated by two transpressional push-up structures (Western High and Central High). The Sea of Marmara presents a long term history of large earthquakes (e.g., Ambraseys and Finkel, 1991). The North Anatolian fault accommodates about 25 mm/year of strike slip motion between the Anatolian block and the Eurasian plate (Fig. 1(A)) (Armijo et al., 1999; Armijo et al., 2002; McClusky et al., 2000; Reilinger et al., 1997). The northern branch of the North Anatolian fault crosses the Sea of Marmara, where it further divides in main and secondary fault branches (Armijo et al., 2002; Becel et al.,

Fig. 1. (A) Regional setting of the tectonic framework in the eastern Mediterranean modified from Okay et al. (2000). (B) Bathymetric map of the Sea of Marmara with the location of studied dives (red dots) and coring sites (white diamonds) along the main Marmara fault (Bathymetry after Rangin et al. (2001) and land from Landsat Msrid, NASA: https://zulu.ssc.nasa.gov/mrsid/). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

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2010; Carton et al., 2007; Hergert and Heidbach, 2010; Le Pichon et al., 2001; Okay et al., 2000). Methane emissions from the seafloor have been identified in the Sea of Marmara and Gulf of Izmit, and are thought to be intensified by earthquake events (Halbach et al., 2004; Kuscu et al., 2005; Kuscu et al., 2008).

seawater with the anoxic freshwater of the glacial Marmara Lake (Reichel and Halbach, 2007). Starting at  12 kyr/cal. BP, the marine sedimentary sequence contains two sapropel layers deposited at  10 kyr/cal. BP and 4.2 kyr/cal. BP (Abrajano et al., 2002; Aksu et al., 2002; Beck et al., 2007; C - a˘gatay et al., 2009; C - a˘gatay et al., 2000; C - a˘gatay et al., 2003; McHugh et al., 2008).

2.2. Oceanographic and paleoceanographic setting 2.3. Cold seeps exploration during the Marnaut cruise Today the hydrology of the Sea of Marmara is characterized by two water masses separated by a permanent pycnocline. The upper water layer ( 25 m thick) corresponds to the outflow of Black Sea brackish waters coming through the Bosphorus strait (sill depth 35 m), and the lower water layer corresponds to the Mediterranean waters entering through the Dardanelles strait (sill depth  65 m) (Bes- iktepe et al., 1994). Due to the shallow depths of the two sills that connect the Sea of Marmara with the Mediterranean Sea and the Black Sea, global sea level variations during the Pleistocene glacial– interglacial cycles were responsible of major changes in the past hydrology and environment of the Sea of Marmara (e.g., Aksu et al., 1999; C - a˘gatay et al., 2000; McHugh et al., 2008). During the last glacial period, the sea level lowering isolated the Sea of Marmara from the Aegean Sea. With no marine water inflow but with river inputs, the Sea of Marmara turned into a freshwater lake (Aksu et al., 1999; C - a˘gatay et al., 2000; Stanley and Blanpied, 1980). The reinvasion of Mediterranean waters into the Marmara Basin occurred at 14.7 kyr/cal. BP but the conversion to marine conditions was gradually established over 2 kyr (Vidal et al., 2010). Within the sediments of the topographic Western High, a disseminated fine-grained authigenic calcite layer (18–25 cm thick), was found at the lithological boundary between lacustrine and marine sediments and estimated to have been formed between 11.5 and 13 kyr/cal. BP (Reichel and Halbach, 2007). This carbonate precipitation event did not show any carbon isotopic anomaly (d13C¼  1.1 to þ0.4% V-PDB) and has thus no relationship with fluid release from seep sites; it may have been formed as a consequence of the mixing of the Mediterranean Sea

Cold seeps and gas bubble emission sites in the Sea of Marmara are widespread, but tend to be more frequent along the edges of the basins, on topographic highs and along active fault outcrops (Geli et al., 2008; Zitter et al., 2008). Active gas venting sites were identified with an EK60 echosounder during the Marnaut cruise by detecting acoustic anomalies created by gas release in the water column (Geli et al., 2008). Seafloor surveys were carried out using manned submersible Nautile. Dives reveal, in addition to free gas bubbles, the presence of black patch areas (up to several meters square) at the location of reduced sediments contrasting with the surrounding superficially bioturbated light beige sediments. Some sites were colonized by endemic benthic communities of cold seeps such as polychaetes (Fig. 2(D)), chemosynthetic bivalves and white/orange microbial mats (Fig. 2(B) and (D)) (Ritt et al., 2010). The western side of Tekirda˘g Basin is associated with a transpressive fault striking along the base of a cliff. Gas escape above the fault occurs along tension gashes and is associated with microseismically active extension at depth (Tary et al., 2011). These fluids present a 3He/4He signature indicative of a mantle contribution (Burnard et al., 2008). Sites on the southern edge of Tekirda˘g Basin (Fig. 2(B)) and on the northeast edge of the Central Basin are located on active fault scarps and expel brackish interstitial water originating from the buried lacustrine sediments deposited in the basins (Zitter et al., 2008). Carbonate chimneys and crusts formed around the conduits and were subsequently broken by tectonic movements (Fig. 2(A)) or slope instability.

Fig. 2. Seafloor pictures from submersible dives. (A) Release of gas bubbles through a fracture containing a white fluffy material (western part of the Tekirda˘g Basin, dive 1647). (B) Carbonate chimneys expelling brackish water and covered by white/orange bacterial mats (eastern part of the Tekirda˘g Basin, dive 1667). (C) Numerous particles sinking under a vast area of chaotic carbonate rocks (Western High, dive 1648). (D) Outcropping carbonate slabs associated with polychaetes in black patches and white bacterial mats on the rim (C - inarcik Basin, dive 1653). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

A. Cre´mie re et al. / Deep-Sea Research I 66 (2012) 114–130

In C - inarcik Basin, cold seeps were found on scree at the base of the Northern cliff, which relates to the main active fault, on a transtensional fault system along the southern edge and at the front of an active landslide at the Southeast end (Geli et al., 2008). Chemical compositions of pore fluids and gas bubbles sampled at the Southern C - inarcik site indicate brackish water and biogenic methane sources (Bourry et al., 2009; Tryon et al., 2010). By contrast, free gas and gas hydrates from sites on the topographic highs have a thermogenic gas signature (Bourry et al., 2009). On Table 1 Geographic coordinates and water depth of carbonate rocks sampled during Marnaut dives. Site

Sample

Coordinates (in degrees, minutes)

Water depth (m)

40150.0430 40148.1790 40148.1100

27130.2150 27137.7690 27137.4600

1056 1113 1118

40148.7840 40148.7860 40148.7860 40149.0560 40149.0700 40149.0600 40149.0610 40149.0610 40149.0300

27146.0600 27146.0590 27146.0580 27146.7770 27146.8220 27146.8280 27146.8270 27146.8270 27146.7400

661 662 662 654 650 649 653 653 659

40149.6220

27156.5140

1197

1650-BC7 1650-BB1 1650-BB2 1650-BB2 1661-R2 1661-R3 1661-R4 1661-R5 1661-R6 1661-R7 1663-BC7 1663-R1 1663-BB chimney 1665-R1

40151.4890 40151.5150 40151.5070 40151.5070 40151.4980 40151.4810 40151.4920 40151.4930 40151.4230 40151.4210 40151.4910 40151.4890 40147.9410 40151.2720

28109.5290 28109.4440 28109.4340 28109.4340 28109.5260 28109.5460 28109.7050 28109.6980 28110.0490 28110.0490 28109.4930 28109.4930 28106.7430 28110.1850

1154 1155 1157 1157 1156 1159 1143 1141 1139 1139 1151 1155 1179 1111

1664-R1 1664-R2

40151.7050 40151.7110

28135.0120 28135.0060

323 326

1653-R3 1653-R5 1658-R1 1659-R1

40142.3000 40142.2730 40148.8280 401 42.9940

29109.4000 29109.5440 29100.3120 29106.9760

1212 1191 1176 1248

Western High 1648-PC5 1648-PC7 1648-PC8 1662-R1 1662-R2 1662-R3 1662-R4 1662-R5 1669-BB Central Basin Western scarp 1649-R1 Eastern scarp

the Central High, fluids are not expelled through the fault zone but on the broad crest south of the fault (Geli et al., 2008). On the Western High, carbonate crusts were sampled on two mounds north of the fault, associated with gas hydrates, gas and oil bubble emissions, and high salinity interstitial water (up to 60 g/L), indicating an origin of the fluids from the Eocene–Oligocene ¨ Thrace Basin (Bourry et al., 2009; Gurgey et al., 2005; ¨ Hos- gormez and Yalc-in, 2005), within the temperature range of the oil window (Tryon et al., 2010). At this site, barite was also found in sediment cores, in mineralized chimneys and presumably forms localized white deposits observed on the seafloor at high flow seepage sites (Tryon et al., 2010).

3. Material and methods 3.1. Samples collection

Latitude (N) Longitude (W) Tekirdag˘ Basin Western scarp 1647-R1 Eastern scarp 1667-R2 1667-R3

117

Central High

C - inarcik Basin

Observations and sampling were realized at fourteen dive sites over the whole Marmara Sea. Fluid seepage areas present various seafloor carbonate deposits, occurring as porous slabs, crusts, pavements that can extend over several meters, and carbonate chimneys. Piles of accumulated carbonate slabs form small mounds (around 20–50 cm high, Fig. 2(C), (D)). The articulate arm of the submersible was used to grab pieces of rock, and a few samples were recovered in superficial sediments by push core (PC) and blade core (BC) (Table 1). Deeper coring in the sediments was carried out in the Tekirda˘g Basin and on the Western-High ridge, using a ten meter long Kullenberg piston corer at locations selected from geophysical surveys (Table 2). Cores were cut into 1 m sections and stored at 3 1C on board. A total of 96 authigenic carbonate samples were thus collected at cold seep sites associated with active faults from widely variable geochemical environments, regarding methane concentration (dissolved or gas phase), hydrocarbon composition (biogenic/thermogenic) and interstitial water salinity and composition (brackish/seawater/ oil reservoir brine). 3.2. Mineralogy Authigenic carbonates were first washed with distilled water to remove salts and sediments. Dried samples were finely crushed in an agate mortar. Bulk mineralogy was determined by X-ray diffraction (XRD) on a Siemens D-500 diffractometer (Cu Ka, Ni-filtered radiation). When necessary, peak correction was performed by reference to the main quartz peak. Integration of peak areas enables a semi-quantitative estimation of mineral phases. Determination of the total carbonate content in weight % (wt%) was performed by reacting 100 mg of carbonate powder with 8N HCl in a manual carbonate calcimeter; the absolute error is 1%. Estimation of the absolute carbonate minerals content was obtained by a combination of these two measurements. The error

Table 2 Geographic coordinates, water depth and core length of Kullenberg cores. Site

Core

Coordinates (in degrees, minutes) Latitude

Longitude

Water depth (m)

Core length (m)

7.75 1.25 8.30 6.30 (Gas hydrate filled) 7.95

Terkirdag˘ Basin Eastern part Western High

MNT-KS30 MNT-KS31 MNT-KS14 MNT-KS27

40148.2160 40148.1550 40149.0520 40148.8920

27137.7860 27137.7740 27146.7670 27146.6400

1118 1101 655 669

Central High

MNT-KS20

40151.7240

28135.0050

335

A. Cre´mie re et al. / Deep-Sea Research I 66 (2012) 114–130

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of this method is approximately 5%. Distribution of the d104 calcite values (main XRD peak) presents a bimodal distribution (Fig. 3) with low-Mg calcite (lMc) (3.015od104 o3.035) and highMg calcite (hMc) (2.975od104 o3.010). The shift of d104 values was used to estimate the substitution of Ca2 þ by Mg2 þ in calcite (Goldsmith and Graf, 1958). The composition of lMc and hMc ranges from 0 to 6 mol% MgCO3 and from 7 to 21 mol% MgCO3, respectively. The MgCO3 mole% content in dolomite was not determined because incorporation of other cations like Fe2 þ and Mn2 þ cannot be excluded; when present, dolomite occurs generally in minor amounts. Scanning electron microscopy (SEM) coupled with an energy dispersive X-ray spectrometer was performed on fresh sections of selected samples allowing the observation of microfacies and a focused qualitative elemental analysis of minerals.

and NBS-19 references. Analytical precision and reproducibility are better than70.05% for both d13C and d18O values. 3.4. Radiocarbon dating AMS  14C dating was performed on unidentified bivalves fragments and on one wood sample at the LMC 14 (CEA, Saclay). Shells were treated in 2 mL of 0.01M HNO3 for 15 min to remove surface coating. Ages were calculated as 14C years BP (Mook and van der Plicht, 1999), corrected for 13C, and the error is expressed as 71s. Calendar age calibration was calculated with the software Calib 6.1.0 (Stuiver and Reimer, 1993), using a 400 years reservoir age for marine shells (Reimer et al., 2009). Results are presented as conventional 14C ages BP ( 72s) with the reservoir effect subtraction (Table 3).

3.3. Stable isotopes analysis The carbon and oxygen isotopic compositions of carbonates were measured on all bulk samples. Nearly 100 mg of powdered samples was reacted with anhydrous orthophosphoric acid for 12 min at 90 1C in a GV Gilson inlet system connected to a dual-inlet IRMS (GV Isoprime). Isotopic compositions are reported in conventional delta (d) units relative to the Vienna Pee Dee Belemnite reference (V-PDB). Sample bracketing was performed with the laboratory reference (Carrara marble) Marceau (d13C¼ þ2.13% V-PDB and d18O¼ 1.83% V-PDB), an internal reference calibrated with the NBS-18

Fig. 3. Distribution of d104 values for calcite and dolomite in carbonate samples.

Fig. 4. Macrofacies of seafloor carbonate concretions. (A) Irregularly shaped fragment of carbonate cemented sediment (sample 1665-R2) with yellowish/ ocrous upper surface. (B) Irregular indurated aragonitic concretion (sample 1664R2). (C) Platy high-magnesium calcite crust (sample 1661-R6) with abundant cavities at the surface possibly representing dissolution features. (D) Carbonate crust covered by a fine black/brown Mn–Fe oxi-hydroxide layer (sample 1661-R4). (E) Tubular concretion corresponding probably to the cementation of a gas conduit for upward migration of methane within sand (sample 1649-R1). (F) Small isolated carbonate concretions (branching tubes) recovered in superficial sediments of a black patch of reduced sediment (sample 1659-R1). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

Table 3 AMS 14C dating, calibrated ages and estimation of the sedimentation rates. Core

Sample

Depth (mbsf)

Nature

Radiocarbon age (year B.P.)

MNT-KS14

II, II, II, II, II, V,

0.64 0.89 1.03 1.03 1.08 3.58

Shell Shell Shell Shell Shell Wood

5 5 9 8 10 15

390 745 925 730 210 735 895 735 790 740 630 760

5 6 10 9 12 18

763 7117 339 769 025 7128 561 790 214 7165 780 7149

11 14 10 11 9 19

MNT-KS27

I, 48 cm I, 87 cm III, 10 cm V, 32 cm

0.48 0.87 1.83 4.05

Shell Shell Shell Shell

1 2 4 11

985 730 385 730 250 730 170 740

1 2 4 12

545 7108 016 796 340 7105 668 7101

31 43 42 32

0 cm 25 cm 70 cm A 70 cm B 75 cm 35 cm

Calendar age (year B.P.)

Sedimentation rate (cm/kyr)

A. Cre´mie re et al. / Deep-Sea Research I 66 (2012) 114–130

4. Results 4.1. Seafloor authigenic carbonates Widespread authigenic carbonate crusts covering the seafloor and concretions in surface sediments were sampled at different sites as outcropping flat centimetre to decimetre-thick pavements cementing bivalve shell fragments, sand, mud and other detrital materials (Fig. 4). Crusts are more or less porous with various colours: whitish, grey, brown and black. A thin red or dark brown

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Fe–Mn oxide/hydroxide layer often covers pavements. Indurated concretions (Fig. 4(F)) were also recovered by pushcores in black patches. The carbonate mineralogy (Table 4) is dominated by aragonite associated with hMc and lMc in lesser amounts, except in the Central Basin where hMc is dominant, sometimes associated with dolomite (up to 16 wt% of the bulk sample). Calcite crystals display various rhombohedral shapes (Fig. 5(A), (B) and (D)). Sheets of mucus-like organic-matter forming a polygonal honeycomb network were observed associated with hMc and aragonite

Table 4 Carbonate content, mineralogy and isotopic composition of seafloor carbonate crusts collected during Marnaut dives. Area

Sample

Total carbonate content (wt%)

Aragonite (wt%)

1647-R1

88

1667-R1 1667-R2 1667-R3 1667-R3 chimney

Low Mg-calcite

High Mg-calcite

Dolomite

d18O V-PDB

d13C V-PDB

(%)

(%)

Weight (%)

mol % Mg

Weight (%)

mol % Mg

Weight (%)

d104 ˚ (A)

88

–

–

–

–

–

–

2.6

 27.9

89 82 79 68

81 31 48 37

7 4 5 6

2 2 3 2

– 47 26 25

– 16 18 16

– – – –

– – – –

1.3 1.9 2.0 1.3

 37.3  40.4  34.7  29.1

1648-PC5 1648-PC7 1648-PC8 1662-R1 1662-R2 1662-R3 a 1662-R3 b 1662-R3 c 1662-R4 1662-R5 a 1662-R5 b 1669-BB

70 78 76 90 81 84 69 92 82 86 82 86

58 63 70 90 76 78 2 90 68 79 77 82

9 8 3 – 4 2 11 2 2 2 3 3

1 2 1 – 1 2 3 1 2 1 1 3

3 7 3 – 1 5 50 – 12 5 2 –

14 18 16 – 15 14 11 – 15 14 14 –

– – – – – – 6 – – – – –

– – – – – – 2.935 – – – – –

2.5 2.5 2.7 3.1 2.7 3.0 2.5 2.5 3.1 2.9 3.0 3.0

 20.0  29.1  27.7  13.9  22.6  27.2  38.7  44.0  20.1  25.3  26.7  12.8

1649-R1

31

5

7

2

19

9

–

–

0.5

 39.9

1650-BlC7 1650-BB1 1650-BB2 CS 1650-BB2 1661-R2 1661-R3 1661-R4 1661-R5 1661-R6 1661-R7 1663-BIC7 1663-R1 1663-BB chimney 1665-R1 a 1665-R1 b 1665-R2 1665-R3

73 55 77 81 82 78 87 87 86 73 76 75 61

8 2 12 1 53 – 39 30 3 37 3 30 –

– – 12 – 7 – – – – 13 – 11 –

– – 4 – 3 – – – – 3 – 3 –

65 53 53 80 13 78 47 57 82 20 72 17 61

10 7 15 11 17 9 12 10 15 15 8 17 7

– – – – 9 – – – – 4 – 16 –

– – – – 2.939 – – – – 2.933 – 2.937 –

1.3 0.2 2.8 2.0 2.7 1.4 1.3 1.7 3.2 2.0 0.5 2.5 2.2

 43.0  47.6  29.0  35.9  44.2  33.6  43.7  33.6  40.9  40.9  42.0  41.0  41.4

84 90 82 81

64 90 35 14

5 – 5 –

2 – 2 –

14 – 37 53

16 – 21 12

– – 4 12

– – 2.929 2.928

3.0 2.8 2.5 2.9

 47.0  44.2 -47.2  40.1

1664-R1 1664-R2

84 88

71 82

5 2

1 1

8 3

19 16

– –

– –

3.1 2.4

 12.9  20.6

1653-R3 a 1653-R3 b 1653-R3 c 1653-R5 1658-R1 1659-R1

89 47 90 89 88 73

86 27 84 89 82 56

1 11 6 – 4 5

1 1 1 – 2 1

3 8 – – 2 12

14 17 – – 15 17

– – – – – –

– – – – – –

2.2 2.4 1.8 2.5 2.3 2.3

 46.4  39.6  39.4  36.5  46.4  29.6

Tekirdag˘ Basin Western scarp Eastern part

Western High

Central Basin Western scarp Eastern scarp

Central High

Cinarcik Basin

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Fig. 5. SEM photomicrographs of seafloor authigenic carbonate crusts. (A) Rhomboedral crystals of foliated high magnesium calcite (sample 1665-R2). (B) Rhombohedral crystals of high magnesium calcite covered by sheets of organic matter, the mucus-like remains forming a filamentous network (white arrow) (sample 1650-BB1). (C) Prismatic crystals of aragonite with dissolution cavities, some of them including silicate flakes (white arrow) (sample 1662-R3). (D) Rice shape crystals of high magnesium calcite with dissolution cavities (1649-R1). (E) Bladed crystals of barite associated with acicular aragonite (sample 1661-R7). (F) Void in Mg-calcite crust infilled with clusters of globular pyrite (sample 1665-R2).

(Fig. 5(B)). Well-crystallized aragonite occurs as fibres or prismatic crystals (Fig. 5(C) and (E)) often filling voids. Local dissolution cavities were observed on aragonite as well as on calcite crystals (Fig. 5(C) and (D), respectively). On the Western-High, sample 1662-R3 shows dissolution cavities of aragonite crystals filled by clay minerals (Fig. 5(C)). Pyrite occurs as disseminated grains and framboids infilling voids, for example foraminifera tests (Fig. 5(F)). Authigenic barite is locally present as elongated and prismatic crystals (Fig. 5(E)). The carbon isotopic compositions of bulk samples from carbonate crusts display a large variability ranging from  47.6 to 12.8% V-PDB. The most 13C-depleted carbonate samples with values around 47% consist of aragonite (1658-R1), hMc (1650BB1) or mixed hMc/aragonite mineralogy (1665-R2). The highest d13C values (up to  13% V-PDB) are only found in carbonate samples containing aragonite (1662-R1, 1669-BB and 1664-R1). Oxygen isotopic compositions of carbonate crusts vary from þ 0.2 to þ 3.2% V-PDB. High d18O values of carbonates are present at Western High (up to þ3.1% V-PDB, 1662-R1) as well as at Central High (up to þ3.1% V-PDB, 1664-R1). 4.2. Carbonate concretions within the sediment Authigenic carbonates within the sediments form light grey to light yellow–brown, millimetre to centimetre thick concretions, nodules and plates (Fig. 6). Samples are well lithified, rounded, flat or tubular. Some of them look like centimetre-diameter

conduits, chimneys or indurated burrows. In thin section, hMc concretion in sample MNT-KS27-VI, 5 cm (Western-High) appears as a clotted micrite with foraminifers, shell fragments, pyrite, and voids more or less filled by microsparite. Botryoidal aragonite fabric was observed in thin section (sample MNT-KS14-II, 75 cm, Western-High) and with SEM (sample MNT-KS14-III, 44 cm, Fig. 7(A)), filling voids presumably associated with bioturbation features in aragonite micrite. Some levels of unconsolidated sediments contain abundant millimetre size carbonate concretions. The host sediments are composed mainly by detrital silicates (quartz, clay minerals, muscovite, and feldspars). The observation of sieved fractions and smear slides shows that a fraction of the carbonate is biogenic (coccoliths, foraminifers, bivalves, ostracods, sponge spicules). Large (up to 500 mm) euhedral prismatic crystals of hyaline gypsum were observed at different levels in the sediments of Western High (i.e., sample MNT-KS14-II, 25 cm; Fig. 8). In the Tekirda˘g Basin, core MNT-KS30 (Fig. 9) contains numerous turbidite intercalations in the upper 5.3 m of dark grey mud, and a debris flow layer (between 4.1 to 4.4 mbsf). From 5.3 mbsf to the bottom, dark brown mud is intercalated with sand layers. Seven carbonate concretions, mostly composed of hMc, were found in this interval from 5.1 to 7.7 mbsf, which corresponds to the Holocene stratigraphic unit. Core MNT-KS31 is only 1.2 m long, and consists of 1 m of grey mud with several hard carbonate layers at the bottom. Their mineralogy consists in mixtures of aragonite, hMc and dolomite (Fig. 7(D)). Carbonate concretions

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from the Tekirda˘g Basin sediments have low d13C values from  51.6 to 19.8% V-PDB and d18O values ranging from  2 to þ2.8% V-PDB (Table 5). Retrieved from the Central High, core MNT-KS20 contains shallow diagenetic crusts (at 0.07 mbsf) showing a d13C value of  23.0% V-PDB and d18O value of þ 2.3% V-PDB.

Fig. 6. Macrofacies of carbonate concretions from sediment cores. (A) Aragonitic concretion with cemented shell fragments (sample MNT-KS14-II, 75 cm). (B) High magnesium calcite smoothed concretion with micro vesicles at the surface (sample MNT-KS27-III, 5 cm). (C) Small concretions corresponding to cemented burrows (sample MNT-KS20-I, 5 cm). (D) Concretion with irregular surface, smooth, microporous or vuggy (sample MNT-KS27-VI, 5 cm) (E) Cm-thick platy grey carbonate crust with upper surface covered by a discontinuous yellowish layer (sample MNT-KS31-I, 105 cm). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

Fig. 8. (A) Macroscopic and (B) SEM observations of gypsum crystals from MNTKS14-II, 25 cm.

Fig. 7. SEM photographs of carbonate concretions. (A) Botryoidal aragonite with a rosette like shape (sample MNT-KS14-III, 44 cm). (B) Large rhombohedral crystals of high magnesium calcite with traces of barite (white arrow, sample MNT-KS27-VI, 5 cm) (C) Scalenohedral crystals of high-Mg calcite with dissolution features (sample MNTKS14-II, 42 cm). (D) Dolomite rhombs with acicular aragonite draped with a dark organic film (white arrow) (sample MNT-KS31-I, 108 cm). (E) Petals of barite with organic film on acicular aragonite (sample MNT-KS14-II, 75 cm). (F) Euhedral pyrite infilling foraminifer test trapped in a cement of aragonite (sample MNT-KS14-II, 75 cm).

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MNT-KS30

MNT-KS31

three main layers of cemented carbonate or nodules are observed within the marine sediments at about 1.0, 1.5 and 3.0 mbsf. A chaotic layer bearing carbonate concretions and organic matter is found in the uppermost part of the lacustrine deposits (from 4.6 to 4.8 mbsf). The nature of this chaotic layer is unclear, as the coring site is located on a mound of complex topography that most likely corresponds to a mud volcano, which could have provided debris flows. Alternatively, gas-hydrate dissociation may have created both disturbance of the sediment and carbonate precipitation, as gas hydrates were recovered from the lower part of the core (Bourry et al., 2009). The isotopic compositions of authigenic carbonate concretions from these two cores present a large range of d13C values from  22% to þ14% V-PDB and d18O values from þ 2.4% to þ4.9% V-PDB.

5. Discussion 5.1. Carbonate mineralogy

Lithology Mud Silty clay Silty sand Medium-coarse sand Carbonates concretions Laminations Shell fragment Wood fragment

Fig. 9. Lithostratigraphy and location of carbonate concretions sampled in cores from the Tekirda˘g Basin (MNT-KS30, MNT-KS31).

At the Western High, two cores MNT-KS14 and MNT-KS27 were retrieved from distinct mound structures north of the main fault. Sediments are impregnated with oil and produce a strong smell of hydrocarbons. The lacustrine marine transition is inferred at 14.5 kyr BP from radiocarbon dating and a change of the micro- and nannofossil fauna and flora in the hemipelagic sediment (Fig. 10). Coccoliths with a rather unusual assemblage with Braarudosphaera sp. are common in the sediment after the lacustrine–marine transition, whereas fresh (e.g., Stephanodiscus neoastrea) to brackish water diatoms species are found below. Two sedimentary units can be distinguished: the younger deposits (from 0 to  12 kyr BP) correspond to the marine stage overlying the lacustrine deposits. The very low total carbonate content of the lacustrine sediments (o2 wt%) seems not to be typical compared to other sites in the Sea of Marmara that do not display such a strong difference in carbonate content between marine and lacustrine levels (C - a˘gatay et al., 2000; McHugh et al., 2008; Reichel and Halbach, 2007). In both cores of the WesternHigh, the marine sequence contains authigenic carbonate concretion-rich levels that are intercalated with grey to dark brown mud, whereas the lacustrine sediments do not show any evidence of diagenetic carbonates. The MNT-KS14 core contains one thick layer bearing carbonate concretions (from 1 to 1.2 mbsf) found just above the marine–lacustrine transition and another one within the transition (from 1.5 to 2 mbsf). In core MNT-KS27,

5.1.1. General consideration The mineralogy of carbonate results from the competitive kinetic growth of the different minerals that is controlled by several physico-chemical factors such as temperature, sulfate concentration, carbonate alkalinity and the Mg2 þ /Ca2 þ ratio and ionic activity of the fluids. At low sulfate concentration, calcite is promoted over aragonite whereas high Mg2 þ /Ca2 þ ratio is thought to inhibit calcite precipitation (Burton, 1993). In the Sea of Marmara, seafloor samples are primarily composed of aragonite because the formation of calcite and dolomite are inhibited due to bottom sulfate-rich seawater with high Mg2 þ /Ca2 þ ratio (Aloisi et al., 2000; Burton and Walter, 1987; Peckmann et al., 2001; Ritger et al., 1987). In addition, the present-day Marmara Sea bottom water has a temperature of 14.5 1C, which has a strong kinetic effect by increasing the precipitation rate of aragonite relative to that of calcite (Burton and Walter, 1987). Among seafloor crusts, carbonates with a significant hMc content are found in the Tekirda˘g Basin and along the northeastern edge of the Central Basin. These sites are characterized by seepage of brackish waters depleted in sulfate and with low Mg2 þ /Ca2 þ ratio issued from late Pleistocene lacustrine sediments (Tryon et al., 2010; Zitter et al., 2008). Such pore waters compositions seem to promote precipitation of hMc over aragonite. Dolomite occurs mainly as a minor mineral phase. The mechanism for dolomite nucleation is generally favoured in sulfate free solutions allowing to overcome the kinetic barrier (Baker and Kastner, 1981), high pH (Hesse and Schacht, 2011) and possibly involves sulfate reducing bacteria (Vasconcelos et al., 1995; Warthmann et al., 2000). Consequently, dolomite crystallization generally occurs at the base of the SMTZ and below in the methanogenic zone (Baker and Burns, 1985; Meister et al., 2011; Meister et al., 2007; Moore et al., 2004). Furthermore, the formation of carbonate crusts significantly reduces sediment permeability and the seawater sulfate diffusion, thus favouring dolomite formation at depth (Bayon et al., 2009; Gontharet et al., 2007; Luff et al., 2005; Luff et al., 2004). A significant proportion of dolomite (up to 42 wt%), associated with hMc, occurs in buried carbonate crusts from core MNT-KS31. This core was taken in the Tekirda˘g Basin where brackish waters are expelled along the fault (Zitter et al., 2008). This major difference in the pore solution chemistry compared to other sites, would favour dolomite and hMc precipitation over aragonite at depth in the anoxic sediments. 5.1.2. Comparison of the Western-High cores In both cores MNT-KS14 and MNT-KS27 at the Western-High ridge, carbonate mineralogy of the concretions differs from one

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Table 5 Carbonate content, mineralogy and isotopic composition of buried carbonate concretions from Marnaut cores. Area/Core

Sample

Tekirdag˘ Basin MNT-KS30 VI, 60 cm VII, 30 cm VII, 35 cm VIII, 41 cm IX, 28 cm IX, 77 cm IX, 93 cm MNT-KS31

Central High MNT-KS20

Total carbonate content (wt%)

Aragonite (wt%)

Low Mg-calcite

High Mg-calcite

Dolomite

Weight (%)

mol % Mg

Weight (%)

mol % Mg

Weight (%)

˚ (A)

d18O V-PDB (%)

(%)

d13 C V-PDB

0.6 1.3 1.3 1.8 0.2  2.0 1.7

 24.3  44.6  45.5  44.4  37.2  19.8  42.8

d104

5.06 5.56 5.61 6.55 7.3 7.53 7.69

63 86 91 70 67 31 85

– 47 55 – – – –

15 – – 6 – – –

3 – – 3 – – –

48 39 36 64 67 31 85

16 15 15 20 9 8 11

– – – – – – –

– – – – – – –

1 1.05 1.08 1.1 1.15 1.2

79 83 80 90 81 92

35 52 4 56 11 81

8 4 5 3 6 4

4 1 3 2 2 3

16 23 28 20 49 6

20 15 19 17 17 16

20 4 42 10 14 –

2.936 2.934 2.926 2.929 2.931 –

2.8 2.3 2.5 2.4 2.7 2.4

 50.6  42.6  47.1  35.8  47.0  42.1

0.05 0.75 0.98 1.03 1.03 1.08 1.08 1.18 1.23 1.48 1.57 1.66 1.67 1.70 1.83 1.88 1.98

79 88 91 90 90 89 91 92 88 86 88 80 92 76 84 86 nd

– – 77 86 87 85 88 89 85 79 77 73 90 45 69 69 –

– – – 3 3 4 3 2 3 6 5 5 – 6 3 5 –

– – – 2 3 3 4 1 1 2 3 1 – 2 2 2 –

79 88 14 – – – – 2 – 2 6 2 2 25 12 11 70

14 10 9 – – – – 12 – 16 15 15 14 18 17 18 12

– – – – – – – – – – – – – – – – 30

– – – – – – – – – – – – – – – – 2.905

3.5 3.3 2.4 3.2 2.9 3.5 2.8 2.9 3.7 3.7 4.1 3.7 3.8 4.0 4.1 3.9 4.1

2.2 14.2  12.7  9.5  13.6  14.4  18.0  15.4  17.0  16.8  16.6  22.4  15.5  13.7  21.3  16.6  14.4

I, 85–90 cm II, 65 cm II, 87 cm II, 90 cm III, 5 cm a III, 5 cm b III, 6 cm IV, 15 cm IV, 25 cm a IV, 25 cm b IV, 40 cm V, 25 cm V, 28 cm V, 87 cm a V, 87 cm b V, 90 cm a V, 90 cm b V, 90 cm c V, 93 cm VI, 5 cm

0.87 1.55 1.60 1.63 1.78 1.78 1.79 2.88 2.98 2.98 3.13 3.98 4.01 4.60 4.60 4.63 4.63 4.63 4.66 4.78

83 85 78 84 86 82 83 84 79 85 88 73 78 70 83 90 84 80 61 94

– – – – – – – – – – – – – – 7 17 – 6 17 –

9 5 7 5 6 7 5 5 9 5 4 1 1 12 5 4 3 5 2 –

3 4 4 4 3 4 3 3 5 4 5 1 2 4 4 5 4 6 3 –

74 80 71 79 80 75 78 79 70 80 84 61 63 58 71 69 81 69 42 94

11 12 13 13 14 12 12 16 12 17 18 14 18 14 15 16 19 17 14 12

– – – – – – – – – – – 11 13 – – – – – – –

– – – – – – – – – – – 2.943 2.937 – – – – – – –

3.3 3.6 3.4 3.4 3.6 3.5 3.3 4.0 3.3 3.6 4.2 4.5 4.9 3.5 4.7 3.8 4.4 4.6 4.0 3.4

2.5 2.2  1.3  2.7  2.6 2.4 2.2  1.3 4.5  5.1  4.2  15.5  21.4 0.5  18.7  14.2  7.8  17.8  16.8 9.8

I, 5 cm a I, 5 cm b I, 7 cm

0.05 0.05 0.07

76 68 67

57 41 39

10 11 12

2 2 2

9 17 16

15 16 16

– – –

– – –

2.6 2.3 2.3

 23.0  22.8  22.9

I, I, I, I, I, I,

100 cm 105 cm 108 cm 110 cm 115 cm 120 cm

Western High MNT-KS14 I, 5 cm II, 42 cm II, 65 cm II, 70 cm a II, 70 cm b II, 75 cm a II, 75 cm b II, 85 cm III, 0 cm III, 25 cm III, 34 cm III, 43 cm III, 44 cm III, 47 cm III, 60 cm III, 65 cm III, 75 cm MNT-KS27

Depth (mbsf)

nd: not determined.

core to another; it is dominated by aragonite in core MNT-KS14 and by hMc in core MNT-KS27. Pore fluid chemistry indicates that the brine source originates from a natural gas and oil reservoir with a possible contribution from gas hydrate formation/dissociation at a shallow subsurface level. Sulfate is absent in pore waters below the upper 30 cm of sediments in MNT-KS14 and below 50 cm in MNT-KS27, and a peak in sulfide concentration is

observed at this level, indicating that sulfate was consumed by AOM (Tryon et al., 2010). Buried carbonate concretions were found down to 4.8 m depth in MNT-KS27 and down to 2 m in MNT-KS14. Hence, based on their location below the present-day SMTZ, it can be supposed that they derived from a paleo-SMTZ, have grown within the first few tens of centimetres within the sediments and were subsequently buried by hemipelagic

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MNT-KS14

Western-High, where high barium concentration (up to 2 mM) was measured in pore fluids (Tryon et al., 2010), it most likely arises from mica and feldspar weathering and also diagenesis from deep fluid rock-interaction. Pyrite is ubiquitous in anoxic sediments, where alteration of detrital iron minerals or Fe-coatings of detrital particles releasing iron into solution control its precipitation. This dissolved iron reacts with hydrogen sulfide produced by bacterial sulfate reduction and it precipitates as sulfide minerals (greigite and pyrite) (Berner, 1984). The abundance of pyrite in diagenetic carbonates from the Marmara Sea confirms that microbial sulfate reduction was a very active process prior to and during carbonate precipitation. Moreover, the presence of both barite and pyrite encrusted within carbonate matrix confirm that carbonate formation took place at the redox chemical boundary between sulfate-rich and sulfate-poor pore waters.

MNT-KS27 Calibrated ages 1,985

5,763 10,025

2,016

Marine deposits 4,250

Tr n io sit an al rv te in

-Wd- 18,780

12,668

Lacustrine deposits

~ 12 kyr ~ 14.5 kyr

6.30 m

Lithology Mud Silty clay

Shell fragment -Wd- Wood fragment

Silty sand Carbonate concretions 8.30 m Fig. 10. Lithostratigraphy, radiocarbon dating and location of carbonate concretions sampled in cores from the Western-High ridge (MNT-KS27, MNT-KS14).

5.3. Oxidation of methane and hydrogen sulfide SEM microfacies observations of seafloor carbonate crusts as well as buried diagenetic carbonate concretions reveal local dissolution of aragonite and hMc crystals. These features are interpreted to result from in situ processes of local acidification enhanced when fluids became under-saturated with respect to the carbonate mineral phase considered. This can result from aerobic oxidation of dissolved reduced species occurring within the upper centimeter oxic sediments above the SMTZ up to the seafloor (Cai et al., 2006) or within the water column. In case of focused methane flux, either as episodic event of gas release or as gentle bubble stream, methane can escape to the AOM filter and may be oxidized by aerobic methanotrophs producing a weak acid (Hanson and Hanson, 1996; Higgins et al., 1981; Himmler et al., 2011; Matsumoto, 1990): CH4 þ 2O2 -CO2 þ 2H2 O Hydrogen sulfide oxidation leads to the formation of sulfuric acid much more corrosive: H2 Sþ 2O2 -H2 SO4

sedimentation. Although both sites record different histories of carbonate precipitation, similar ambient seawater conditions (i.e., temperature, chemical and isotopic compositions of bottom seawater) may be assumed for both sites, at least for the deepest layer that is found close to the lacustrine–marine transition. Several assumptions can explain the difference of carbonate mineralogy at these two sites, such as chemical composition of fluids that exhibits different Mg2 þ /Ca2 þ ratio (  1.6 for MNTKS14 and  0.8 for MNT-KS27) and/or variability of fluid flow velocity over location and time that might have affected the depth of the SMTZ (Luff and Wallmann, 2003; Tryon and Brown, 2004). Moreover, hMc precipitation in sediments from core MNT-KS27 could be related to gas hydrate destabilization, likely as the micro-crystalline hMc from the Cascadia margin enriched in 18O (Bohrmann et al., 1998). 5.2. Sulfide and sulfate minerals associated with the carbonate matrix Carbonate concretions and crusts contain minor amounts of authigenic barite and pyrite. Barite is frequently associated with cold seeps carbonates and its precipitation occurs at the SMTZ or above where fluids containing high barium concentration encounter downward diffusing seawater sulfate (Aloisi et al., 2004; Castellini et al., 2006; Torres et al., 1996). Barium enrichment at depth in pore-fluids originates, in most cases, from the remobilisation of biogenic barite in organic-rich sediments. However on

Both reaction enhances acidification and can lead to partial dissolution of carbonate minerals. In the sediments, bioturbation and bio-irrigation by benthic organisms can amplify the process by ventilating sub-bottom pore waters. Accordingly, the occurrence of authigenic carbonate dissolution features indicate that carbonate stability in shallow sediments is linked to the oxygen depth penetration in the surficial sediments. In addition, oxidation of pyrite is also thought to produce sulfuric acid that dissolves carbonate minerals: 4FeS2 þ 15O2 þ 14H2 O-8H2 SO4 þ 4FeðOHÞ3 The Ca2 þ ions released by carbonate dissolution can react with sulfate ions incoming from seawater or produced by H2S oxidation and then precipitate as authigenic gypsum crystals within the sediments (Fig. 8) (Pierre et al., In press; Pirlet et al., 2011). More locally at the Western-High site, the gas phase migrating with fluids contains 4% CO2 (Bourry et al., 2009) which possibly indicate that acidic fluids can dissolved carbonate. However, as pore water alkalinity and pH are unknown, saturation of fluids relative to carbonate cannot be evaluated. The thin dark to light brown (Fe and Mn) oxyhydroxydes layers coating the surface of some carbonate crusts outcropping at the seafloor income from reduced fluids containing dissolved Fe2 þ and Mn2 þ that reach seafloor and are oxidized in the bottom water (Bayon et al., 2011; Charlou et al., 2004). Such hydrogenous deposits testify that fluids containing methane and other reduced compounds escape through the sediments.

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5.5. Temperature and oxygen isotopic composition of fluids

O-ric

h flu

ids ?

The oxygen-isotopic composition of carbonates provides information about the isotopic composition of the water from which the carbonates precipitated and about the temperature of precipitation. The oxygen-isotopic fractionation between water and mineral phases is different for each carbonate mineral. Presently, the bottom water of the Sea of Marmara exhibits the same oxygen

C-rich DIC

18

The carbon isotope composition of carbonates may originate from various carbon sources. Sediments contain an initial and variable quantity of detrital and biogenic carbonates (e.g., coccoliths, shell fragments, foraminifera) that are trapped during the growth of diagenetic carbonate cement. The carbonate cement precipitates from dissolved inorganic carbon (DIC) in pore fluids that originates mainly from the microbial oxidation of methane and possibly heavier hydrocarbons. Isotopic compositions of free gas sampled at the seafloor show that the origin of hydrocarbons varies at the different sites between microbial and thermogenic sources (Bourry et al., 2009). In the C - inarcik Basin, microbial methane with a d13C value of  64.1% V-PDB is the main gas component (99.6%) whereas at the Western-High and CentralHigh ridges, thermogenic methane (d13C¼  44.4% V-PDB) is associated with heavier hydrocarbons (from C2 to C5) with d13C values ranging from  28 to  8.9% V-PDB and CO2 with d13C ranging from þ25 to þ29% V-PDB (Bourry et al., 2009). Other sources of DIC, supposed to be minor in cold seeps compared to DIC released by AOM, are seawater (d13C around 0% V-PDB and bottom water concentrations of 2 mM), and mineralization of marine organic matter (d13C   20 to  25% V-PDB). The carbon isotopic compositions of diagenetic carbonates show a wide variability from 50.6 to þ 14.2% V-PDB indicating different diagenetic settings (Figs. 10 and 11). Parameters that possibly affect the carbonate isotopic composition are multiple sources of the DIC pool from which carbonates precipitate, temporal variations of hydrocarbon sources supplied to the sulfate-methane reaction zone (i.e., advection of deep-seated fluids from the petroleum reservoir versus shallow microbial methane formation), the proportion of biogenic and detrital carbonates initially present in the sedimentary matrix and microbial metabolic pathways that influence the carbon isotopic fractionation between reduced hydrocarbons and DIC. In the Tekirda˘g, Central and C - inarcik Basins, authigenic carbonate crusts and buried concretions are depleted in 13C (average d13C¼  40% V-PDB, n¼41) implying that 13C-depleted methane, most likely originated from microbial decomposition of organic substrates in anoxic sediments, was the main source of carbon and that microbial oxidation processes have transformed this reduced form of carbon into DIC. These results support the idea that carbonate precipitation was enhanced by an increase of alkalinity mediated by the microbial anaerobic oxidation of methane. This is in agreement with the biomarker study of seafloor carbonate crusts showing high concentrations of microbial, 13C-depleted lipids (Chevalier, 2010; Chevalier et al., 2011), which confirms the involvement of AOM mediated by methanotrophic archaea and sulfate-reducing bacteria consortia in the formation of diagenetic carbonates covering the seabed. On the ridges (Western and Central Highs), where isotopic evidence for thermogenic gas was found, the carbon isotopic composition of diagenetic carbonates points to more complex carbon sources. The d13C values of seafloor carbonate crusts vary over a wide range (d13C¼ 44 to  12.8% V-PDB). The buried carbonate concretions from Western-High ridge also exhibit an important variability of d13C values (d13C ¼  22.4 to þ14.2% V-PDB) with a range shifted towards high positive values. These 13 C signatures of authigenic carbonates suggest the involvement of sources of carbon other than methane (Fig. 12). At both sites, migration of thermogenic fluids is associated with a notable proportion of hydrocarbons heavier than methane. In the Gulf of Mexico where there are similar seepages of thermogenic hydrocarbons, it was suggested that microbial oxidation of heavy hydrocarbons including crude oil was the source of bicarbonate for carbonate precipitation with d13C values around 25% V-PDB

(Feng et al., 2009; Formolo et al., 2004; Mansour and Sassen, 2011; Naehr et al., 2009). In Western High, crude oil impregnates sediments and carbonates pore spaces that might indicate a similar process; however, microbial oxidation of hydrocarbons cannot explain d13C values as high as þ14% V-PDB. In the methanogenic zone, the partial biogenic reduction of CO2 may be responsible for the d13C increase of the remaining CO2 pool (Claypool and Threlkeld, 1983; Whiticar, 1999). Here, we propose that microbial oxidation of crude oil, combined with methanogenesis at depth could generate CO2 enriched in 13C (Bourry et al., 2009; Etiope et al., 2009; Head et al., 2003; Jones et al., 2008). Decrease of the pore water pH, generated by CO2 production, leads to carbonate dissolution, as mentioned before. However, silicates weathering by CO2 can buffer pH of fluids during their migration through the sedimentary column (Claypool and Threlkeld, 1983; Meister et al., 2011; Wallmann et al., 2008). Migration of these fluids with high alkalinity and positive d13C values of DIC, arising from deep horizons through fractured pathways, would have significantly contributed to the formation of 13C-rich carbonates in the subsurface (compared to typically 13 C-depleted cold seep carbonates). Variable mixtures of oxidized hydrocarbons and fluids containing 13C-rich DIC are thought to be possible sources to explain the wide range of carbon isotopic compositions of carbonate concretions (Fig. 12).

δ O ‰ V-PDB

5.4. Origin of dissolved inorganic carbon

125

Br

ac

kis

h

wa

te

rs

δ O isotopic equilibrium of dolomite δ O isotopic equilibrium of aragonite δ O isotopic equilibrium of hMc

δ13C ‰ V-PDB Seafloor carbonate crusts :

Buried concretions :

Tekirdağ Basin dives (1647, 1667)

Tekirdağ Basin cores (MNT-KS30, MNT-KS31)

Western High dives (1648, 1662, 1669)

Western High cores (MNT-KS14, MNT-KS27)

Central Basin dives (1649, 1650, 1661, 1663, 1665)

Central High core (MNT-KS20)

Central High dive (1664) Çinarcik Basin dives (1653, 1658, 1659)

Fig. 11. Carbon and oxygen isotopic compositions of carbonate crusts and concretions. Dotted and dashed lines represent the d18O values of high Mg-calcite (15 mol% Mg), aragonite and dolomite precipitated in isotopic equilibrium with the present day Marmara bottom sea-water (d18O water¼ þ 1.4% V-SMOW, T¼14.5 1C).

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isotopic composition (d18O¼ þ 1.58% V-SMOW) as the eastern Mediterranean bottom water (Rank et al., 1999). The fractionation factors a at 14.5 1C (the bottom water temperature) between calcite and water (Kim and O’Neil, 1997) corrected for a substitution of 15 mol% MgCO3 (Tarutani et al., 1969), aragonite and water (Grossman and Ku, 1986; Han et al., 2004) and dolomite and water (Zheng, 1999) were calculated to estimate the theoretical d18O values of hMc (þ2.34% ), aragonite (þ2.78% ) and dolomite (þ5.81%) that would precipitate in isotopic equilibrium with the present-day bottom water. Most of the seafloor carbonate crusts exhibit d18O values close to isotopic equilibrium with the present-day bottom seawater (Fig. 11) indicating that most probably they formed close to the seafloor. For a few samples, the d18O values fall outside the isotopic equilibrium values with the present-day bottom water. Values as low as þ0.16% for a carbonate crust and  1.96% for buried concretions from the Tekirda˘g Basin (Eastern part) and þ1.3% in crusts from the Central Basin (Eastern scarp) reflect various contributions of brackish waters. These sites of fluid expulsion from buried lacustrine Pleistocene sediments are located at the intersection of canyon and earthquake rupture zones (Armijo et al., 2005; Zitter et al., 2008). On the Western-High ridge, the Holocene marine deposits contain carbonate concretion-rich levels that exhibit high d18O values (up to þ4.9% V-PDB, Fig. 11). The carbonate mineralogy cannot explain this oxygen-isotopic enrichment by about 2% compared to the d18O value at equilibrium with present-day bottom water calculated for aragonite ( þ2.78% V-PDB). Therefore, lower bottom-water temperature during precipitation and/ or the influence of 18O-rich fluids may have led to the increase of d18O values of these buried carbonate concretions. After the transition from the late glacial maximum to the Holocene when the Marmara freshwater lake was gradually converted to a marine environment, climate warming led to an increase by about 10 1C of the sea surface temperature until the hydrological conditions became stable at 8 kyr (Vidal et al., 2010). At the seafloor, benthic foraminifers recorded an 18Oenrichment of about  2%, which resulted from the combined changes of temperature and isotopic composition of the water (Aksu et al., 2002; C - a˘gatay et al., 2009). Considering the depth of the Dardanelles sill (  65 m) and timing of the Mediterranean water incursion, calculation with a bottom water temperature of 10 1C and an isotopic composition of þ2.2% V-SMOW (compromise between LGM and present-day Mediterranean bottom water values), the late-glacial/Holocene transition equilibrium d18O values would be around þ 4.4% for aragonite and þ3.9% for hMc.

Different processes liberate 18O-rich water such as clay ¨ mineral dehydration (Dahlmann and De Lange, 2003), gashydrate dissociation (Davidson et al., 1983; Hesse and Harrison, 1981; Matsumoto, 1989) and deep-sourced fluids coming from oil and/or gas fields (Sofer and Gat, 1975) or modified by mineralwater interactions (Clayton et al., 1966; Giggenbach, 1992; Holser et al., 1979). Because all these hypotheses are conceivable to explain the increase of the 18O content of diagenetic carbonates, it seems to be evident that the oxygen-isotopic composition of carbonate alone is not sufficient to favour the importance of anyone of the proposed processes. Notwithstanding, the strong overprint of buried diagenetic fluids and the occurrence of shallow gas hydrates have probably both contributed to the oxygen isotopic enrichment of authigenic carbonates. 5.6. Timing of carbonate precipitation The age of the seafloor crusts at most sites is not known, but their association with active fluid emissions suggests that they are relatively recent compared to the carbonates buried within the sediments. Considering the high sedimentation rates estimated for the basin sites (more than 1 mm/yr, Armijo et al., 2005; Beck et al., 2007), it is very likely that the small carbonate concretions sampled immediately below the sediment surface (e.g., dive 1659) correspond to very recent, presumably still active precipitation. Samples collected from cores on the Western High are embedded in hemipelagic sediments. The depth at which the carbonates formed initially is unknown, but it was probably less than the present depth of the AOM (  50 cmbsf). Occurrences of carbonates are found near the lacustrine to marine transition, within it and around 20 cm above in core MNT-KS14, and 40 cm below the transition in core MNT-KS27. These two levels may however correlate stratigraphically, if we assume that both have formed shortly after the establishment of dominantly marine conditions. In this case, the precipitation occurred close to the seafloor for MNT-KS14 and deeper at the location of MNT-KS27. This assumption would be consistent with the carbonate mineralogy (aragonite and high Mg-calcite, respectively). The isotopic signature of these carbonates shows that it is unlikely that they purely formed by the mixing of fresh and marine waters in the water column, as proposed by Reichel and Halbach (2007). Nevertheless, the increase of the carbonate saturation index caused by the mixing of Mediterranean water with the residual lacustrine porewater could have triggered carbonate diagenesis. It should also be considered that, independently of the variations of methane flux, the new input of sulfate into the system during the

Fig. 12. Range of d13C values of authigenic carbonates (Tekirda˘g Basin: 5 samples at seafloor, 13 samples in cores; Central Basin: 18 samples at seafloor; C - inarcik Basin: 6 samples at seafloor; Central High: 2 samples at seafloor, 3 samples in core; Western High: 12 samples at seafloor, 37 samples in cores) and of hydrocarbons, carbon dioxide of free gas and gas hydrates. (a) biogenic methane from the Tekirda˘g Basin, (b) thermogenic hydrocarbons and CO2 from the Western-High ridge (from Bourry et al., 2009).

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lacustrine–marine transition enhanced microbial oxidation of hydrocarbons and promoted authigenic carbonate precipitation. Sulfate concentration in the bottom water and its diffusion in subsurface sediments was certainly an important factor that allowed hydrocarbon consumption mediated by microbial communities and authigenic carbonate formation over the late deglaciation period in the Sea of Marmara sediments. In core MNT-KS27 at 6 mbsf, thermogenic gas hydrates were found at in situ T–P conditions (14.5 1C and 669 water depth) far away from the pure methane hydrate stability field (Bourry et al., 2009). Shallow gas hydrates are very sensitive to temperature and pressure fluctuations. In a recent study of the sediments from a core close to the same site, Me´not and Bard (2010) identified biomarkers specific of methanotrophy at 11 kyr BP and interpreted them by a large methane release event triggered by gashydrate dissociation in response to the temperature increase of the bottom water. Thus, dissociation of methane hydrate during the warming event associated with deglaciation could be another important factor at the sites cored on the Western High carbonate mounds. Several discrete diagenetic carbonate concretions occur inter-bedded in core MNT-KS27 within the marine sediments (but not in core MNT-KS14), and might correspond to episodes of higher methane flux, gas-hydrate dissociation/dissolution and carbonate precipitation at this site.

6. Conclusions In the Sea of Marmara, the North Anatolian fault system channels widespread fluid emissions of various origins including brines to brackish waters that are rich in microbial and/or thermogenic hydrocarbons. Authigenic carbonate crusts close to the seafloor and carbonate concretions within the sediments record the composition of the diagenetic fluids at the seeping sites. Aragonite crusts covering the seabed and high Mg-calcite associated with lacustrine water outflows or observed buried within the sediments are the dominant authigenic carbonate minerals sometimes associated with minor occurrence of dolomite. Sulfide and sulfate minerals (pyrite and barite) are commonly associated with the diagenetic carbonates. The d18O values of seafloor carbonate crusts emphasize that they generally formed close to the isotopic equilibrium with the ambient bottom water or more locally with brackish water outflows. The wide range of d13C values of the carbonates reflects the sources diversity of dissolved inorganic carbon including microbial methane, thermogenic hydrocarbons and deep-seated fluids. The 13C-rich authigenic carbonates may be related to the anaerobic microbial oil biodegradation in a subsurface petroleum reservoir. Specifically at the Western High, the isotopic variability of buried concretions reflects different diagenetic settings due to temporal fluctuations of the fluid compositions supplied to the sulfate-methane reaction zone. The location of carbonate concretions within the sedimentary succession indicates that there is a link between the carbonate diagenesis and the paleoceanographic evolution during the last deglaciation. During the early Holocene, the increase of sulfate concentration arising from the incursion of Mediterranean waters, together with the warming of the Sea of Marmara that probably caused massive release of methane from gas hydrate dissociation, are thought to have favoured conditions for diagenetic carbonate precipitation within the sediments on the Western-High ridge. These findings reveal that early carbonate diagenesis is interrelated with fluids migration through active fault but also with the late past climate and environmental changes in the Sea of Marmara.

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Acknowledgements The Marnaut Project has supported this work. We are indebted to the captain, officers and crew members of the Marnaut cruise (2007) on the RV Atalante and submersible Nautile who allowed the collection of carbonates and sediments samples in good conditions and to the Turkish Coast Guards who allowed safe working conditions in heavy ship traffic in the Sea of Marmara. We also thank Ioanna Bouloubassi and Ce´line Grall for their contribution in core description and sampling. We would like to acknowledge warmfully Omar Boudouma UMR 7193 ISTEP for his guidance during SEM observations. We are grateful to Jean Pascal Dumoulin and Christophe Moreau for the AMS 14C measurements carried out at the UMS 2572 LMC14 of Saclay. We thank Giovanni Aloisi for constructive discussions. Reinard Hess, Patrice Imbert and an anonymous reviewer are acknowledged for their constructive reviews that improve this manuscript.

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