Dynamics of pyrite formation and organic matter sulfurization in organic-rich carbonate sediments

Accepted Manuscript Dynamics of pyrite formation and organic matter sulfurization in organic-rich carbonate sediments Lubna Shawar, Itay Halevy, Ward Said-Ahmad, Shimon Feinstein, Valeria Boyko, Alexey Kamyshny, Alon Amrani PII: DOI: Reference:

S0016-7037(18)30503-9 https://doi.org/10.1016/j.gca.2018.08.048 GCA 10921

To appear in:

Geochimica et Cosmochimica Acta

Received Date: Accepted Date:

30 March 2018 30 August 2018

Please cite this article as: Shawar, L., Halevy, I., Said-Ahmad, W., Feinstein, S., Boyko, V., Kamyshny, A., Amrani, A., Dynamics of pyrite formation and organic matter sulfurization in organic-rich carbonate sediments, Geochimica et Cosmochimica Acta (2018), doi: https://doi.org/10.1016/j.gca.2018.08.048

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Dynamics of pyrite formation and organic matter sulfurization in organicrich carbonate sediments Lubna Shawara, Itay Halevyb, Ward Said-Ahmada, Shimon Feinsteinc, Valeria Boykoc, Alexey Kamyshnyc, Alon Amrania* a

Institute of Earth Sciences, Hebrew University, Jerusalem, Israel

b

Department of Earth and Planetary Sciences, Weizmann Institute of Science, Rehovot, Israel

c

Department of Geological and Environmental Sciences, Ben-Gurion University of the Negev, Beer Sheva, Israel *Corresponding author: [email protected]

Abstract Organic-rich carbonate sediments are deposited in a range of environments today and in the geologic past. A significant part of organic matter (OM) degradation in such sediments often occurs by microbial reduction of seawater sulfate, and the sulfide product may be preserved in pyrite and in organic sulfur (S) compounds. The isotopic composition (δ34S) of these phases can provide valuable information about S cycling in the ocean and in sediment porewaters, but only insofar as the processes governing these δ34S values are understood. To this end, we investigated the pathways, timing and interactions between pyrite and organic S formation during the deposition of organic-rich chalks. As a test case, we studied cores representing the thickest (~350 m) and most complete Late Cretaceous organic-rich sequence along the southern Paleotethyan margin. The organic S and OM contents show an inverse relation with the pyritic S content, which together with the uniform FePy/FeHR ratio (~40%), suggest competition between organic S and pyrite formation. Both kerogen and pyritic S are 34

S-depleted relative to Late Cretaceous marine sulfate (34S~17-20‰), but the kerogen S is 1

consistently and unusually 34S-enriched relative to coexisting pyrite by up to ~38‰. Large S isotope fractionation (~60‰) during microbial sulfate reduction is necessary to reproduce the lowest pyrite 34S values in the core, and relatively invariant 34S values in organic S suggests that this large fractionation was approximately constant during deposition of the chalks in the core. Higher pyrite 34S values observed in the most organic-rich parts of the core may be explained by Fe-limited pyrite formation, perhaps due to the reaction of Fe (e.g., complexation, sorption) with organic compounds. Lesser Fe availability, relative to the OM available for sulfate reduction, limits the ultimate abundance of pyrite, but importantly, it delays the formation of pyrite to deeper below the sediment-water interface, from

34

S-

enriched sulfide produced by Rayleigh distillation of a dwindling sulfate reservoir. Thus, it appears that competing Fe-OM, S-OM and Fe-S reactions can significantly affect the 34S values recorded in pyrite in organic-rich carbonate sediments despite large and relatively constant microbial S isotope fractionation. Keywords Pyrite; Organic sulfur; Reactive iron; kerogen; Sulfur isotopes; organic-Fe complex; Secondary iron sulfides; Competition; Diagenesis 1. Introduction Sedimentary organic sulfur (S) is generally the second largest pool of reduced S after pyrite in rocks and sediments, and in some cases even the largest (Anderson and Pratt, 1995; Werne et al., 2004). The interaction between the sulfur and carbon cycles in the formation of organic-rich rocks can record important information related to redox changes and other biogeochemical events in ancient environments (Sinninghe Damste and De Leeuw, 1990; Amrani, 2014). Moreover, organic S plays an important role in fossil fuel formation and

2

degradation pathways (Tannenbaum and Aizenshtat, 1985; Orr, 1986; Sinninghe Damste, 1990; Baskin and Peters, 1992; de Leeuw and Koopmans et al., 1996; Lewan, 1998). Both sedimentary organic S and pyrite are formed mainly by the abiotic reactions of reduced S species (e.g., HS-, Sx2-) with organic compounds and iron (Fe) species respectively (Goldhaber and Kaplan, 1974; Sinninghe Damste and De Leeuw, 1990; Rickard and Luther, 2007). Reduced S species in anoxic conditions are generated by dissimilatory microbial sulfate reduction (MSR) (Kaplan et al., 1963; Kaplan and Rittenberg, 1964). A large S isotope fractionation (up to ~75‰) is associated with MSR, with profound effects on the isotopic composition of S species in sediments and sedimentary rocks (Wortmann et al., 2001; Brunner and Bernasconi, 2005; Canfield et al., 2010; Sim et al., 2011). Hydrogen sulfide can be reoxidized back to sulfate, either microbially, in which dissolved oxygen or nitrate are used as electron acceptors, or abiotically (e.g., by dissolved oxygen or oxides of Mn(IV) and Fe(III)). This oxidation leads to the formation of a variety of S species of intermediate oxidation-state, both solid (e.g., elemental S) and dissolved (e.g., S x2-, S2O32-, SO32-) (Fry et al., 1988; Zhang and Millero, 1993). Hydrogen sulfide and other reduced S species in sediments react with metals such as iron (Fe) to form 34S depleted Fe sulfides (e.g., FeS, Fe3S4), the most stable and common form being pyrite (FeS 2) (Goldhaber and Kaplan, 1974; Rickard and Luther, 2007). Sulfate anions may precipitate in evaporative environments as gypsum (CaSO4×2H2O) and anhydrite (CaSO4), or in normal marine environments as barite (BaSO4) and carbonate associated sulfate (CAS), and these are usually the most enriched S fraction. The precipitation of S minerals is associated with negligible

34

S-

34

S

fractionations (Price and Shieh, 1979; Raab and Spiro, 1991; Worden et al., 1997; Butler et al.,

2004).

Thus,

the

S-isotopic

composition of

minerals

(mainly

pyrite

and

gypsum/anhydrite/barite) in marine sediments has been suggested to be representative of the 34S of the sulfide and sulfate in sediment pore waters, and by proxy seawater sulfate. This in 3

turn can provide information on the paleo-redox conditions and diagenetic processes in geological sedimentary deposits, and sometimes on the marine S cycle (Claypool et al., 1980; Canfield and Teske, 1996; Johnston et al., 2005; Fike et al., 2006; Gill et al., 2007; Fike and Grotzinger, 2008; Hurtgen et al., 2009; Jones and Fike, 2013) Reduced S species react also with organic compounds to form organic S compounds (OSCs) in sediments during the early stages of diagenesis, and in an euxinic water column during OM deposition, in a process termed sulfurization or vulcanization (Sinninghe Damst'e and De Leeuw, 1990; Raven et al., 2016a). This process is thought to occur when Fe is depleted in the sediment because otherwise it may scavenge reduced S species before they react with organic compounds (Berner and Westrich, 1985; Hartgers et al., 1997). However, other studies suggest simultaneous formation of organic S and Fe sulfides (Urban et al., 1999; Filley et al., 2002). The sulfurization process quenches functionalized organic molecules that are very reactive and would otherwise be mineralized rapidly by microorganisms, thus enhancing organic matter preservation (Sinninghe Damst'e and De Leeuw, 1990; Grice et al., 2003). These OSCs are structurally and isotopically distinct from biosynthetic S compounds and relate to almost every main biological precursor molecule (Sinninghe Damste and De Leeuw, 1990). The isotopic composition preserved in reduced S species in sedimentary rocks exhibits complex dependences on several factors. Among these are the isotopic compositions of the source sulfate, fractionations associated with the microbial and abiotic processes that led to the formation of reduced S and its incorporation into preserved phases, where sulfurization processes take place (i.e., water column and/or sediment), transport within sediment porewater, and exchange with the overlying water column. In addition to these diagenetic processes, after deposition and lithification, the isotopic composition of S species such as pyrite and sulfate, may be further altered by processes such as oxidation, interaction with 4

meteoric fluids or with basinal brines (Fike et al., 2015). The sum of these processes may yield δ34S compositions that are different from that of the original seawater sulfate and the H2S produced from it by MSR, complicating the reconstruction of the ancient S cycle and redox conditions on the basis of isotopic signals observed in sedimentary rocks. In contrast with pyrite, there is little information about how diagenetic or late-stage alterations might change the organic S pool. It is known that despite the common origin of S in sedimentary organic S and Fe sulfides, organic S is typically 34S-enriched relative to coexisting pyrite by up to 30‰ with a global average of about 10‰ (Anderson and Pratt, 1995; Amrani et al., 2005; Raven et al., 2016b). Some of these organic S-pyrite S isotope fractionations originate during early sedimentation and diagenesis, but some may continue to the later stages of diagenesis (Amrani, 2014). This poorly understood phenomenon may have important implications for our understanding of the S cycle in the oceans and in marine sediments. An improved understanding of the relationships between pyrite and organic S and their precursor sulfate may provide another control on reconstructions of the ancient S cycle based on S isotope compositions in sedimentary rocks. This can be especially useful when organic S is the major S phase, as in the Ghareb Formation in Israel, the Monterey Formation in western California, the Great Marsh in Delaware, and Epilimnetic sediments from lakes in southern Norway (Rudd et al., 1986; Cutter and Velinsky, 1988; Zaback and Pratt, 1992; Amrani et al., 2005). To investigate the processes that govern the isotopic composition of pyritic and organic S, we studied a ~350 m-long sequence of organic- and S-rich chalk from the Late Cretaceous in central Israel. In this organic-rich chalk, which we accessed by nine core samples from the Aderet borehole (Fig. 1b), the major S phase is organic. We used a detailed quantification of S and Fe speciation and analyses of S isotope ratios in inorganic and organic S species, to reconstruct the S cycle and sulfurization pathways of organic matter and Fe. The section 5

drilled is an excellent representative of one of the most extensive systems of upwelling deposits, covering large regions of the southern Tethys shelf (Almogi-Labin wt al., 1993; Eshet and Almogi-Labin, 1996; Soudry et al., 2006; Minster, 2009; Meilijson et al., 2014). The Aderet section drilled in the depo-center of the Shefela Basin, Israel is among the thickest sections through the Tethys shelf (Fig. 1a). Additionally, it represents among the longest-lasting continuous accumulation, ~19 m.y. Both its thickness and deposition duration make the Aderet section especially valuable for the study of the upwelling depositional system that prevailed in the Late Cretaceous along the southern margin of the Tethys ocean (Tannenbaum and Aizenshtat, 1985; Gvirtzman et al., 1985; Reiss et al., 1985; Almogi-Labin et al., 1993; Soudry et al., 2006; Meilijson et al., 2014). Thus, the findings shed light on the effects of early and late diagenesis on the major S species’ concentrations and 34S values in organic-rich Cretaceous carbonates, and more generally, on the controls over the S isotope composition of pyrite and organic S in marine sedimentary rocks in one of the most extensive upwelling and high productivity systems known. 2. Methods 2.1. Study site In the current study we analyzed nine core samples of the organic-rich chalk sequence from 260 m to 600 m below the surface in the Aderet borehole, drilled into the depo-center of the Shefela Basin (Fig. 1b). The studied section comprises predominantly the immature organicrich chalk with Type II-S kerogen, (atomic S/C= 0.06-0.16), of the En Zeitim Formation (stratigraphically equivalent to the Ghareb Formation in the Negev, Israel) (Flexer, 1968; Reiss et al., 1985). It represents a high-productivity succession developed at the Late Cretaceous outer shelf belt due to an extensive upwelling system along the southern Tethyan margin (Fig.1a) (Almogi-Labin et al., 1993; Eshet and Almogi-Labin, 1996; Meilijson et al., 2014). During the deposition of these sediments the water column was suggested to be 6

dysoxic to anoxic but non-euxinic (Amit and Bein, 1982; Bein et al., 1990; Alsenz et al., 2015). The sediments accumulated during the Campanian time are characterized, in general, by enrichment in silica (Si), phosphorous (P), organic carbon, carbonates and S relative to the overlying organic-rich chalk accumulated during the Maastrichtian (core samples from 250450m) (Meilijson et al., 2014, 2015). This represents a major change in the dynamics of the upwelling system and a ~3-fold increase in the rate of sediment accumulation at the transition to the Maastrichtian (Fig. 2) (Meilijson et al., 2014, 2015). The transition is also associated with a change in the dominant forms of primary producers in the upper water column, from diatoms during the more siliceous Campanian to calcareous nanoplankton during the Maastrichtian (Almogi-Labin et al., 1993; Eshet et al., 1994; Eshet and Almogi-Labin, 1996; Ashckenazi-Polivoda et al., 2011;Meilijson et al., 2014, 2015). 2.2 Materials AgNO3 ACS reagent (>99%), anhydrous magnesium sulfate (>99.5%), chromium (III) chloride hexa-hydrate (>98%), hydrogen peroxide (30%), tin (II) chloride (>98%) and zinc (20-30 mesh; 99.8%) were purchased from Sigma-Aldrich (St. Louis, USA). Barium chloride dehydrate (>99%) was purchased from Alfa Aesar (Haverhill, USA). Buffer solution (pH 4) was purchased from J.T. Baker (USA). All the analytical grade solvents (dichloromethane, HCl 37%, methanol and pentane) were purchased from Bio-lab Ltd. (Israel). All the solvents were distillated again in the lab. The S isotope reference materials NBS-127 (BaSO4 ; δ34S = 21.1‰), IAEA-S-1 (Ag2S; δ34S = -0.3‰), and IAEA-SO-6 (BaSO4; δ34S = -34.1‰) were purchased from the National Institute of Standards and Technology (NIST, USA) and used for calibration of the IRMS. 2.3 Sample preparation for analysis

7

About 300 g of the upper part of each core was cleaned and smoothed using sandpaper to remove the outer layer of the rock that was exposed to air. The cleaned rock was then manually ground using a ceramic grinder and then stored in a capped vial. 2.4 Elemental analysis and Rock Eval Total organic carbon (TOC) and total sulfur (TS) contents of the core samples were measured by a LECO SC632 elemental analyzer. About 100 mg of the sample, after acidification using H3PO4 (10%) to remove carbonates, was placed in a porcelain crucible and combusted in a pure oxygen atmosphere at 1350°C. Released sulfur dioxide and carbon dioxide were quantified by an IR detector. Rock Eval 6 (Vinci Technologies, France) was used to identify the source rock type, production potential, and maturity of the organic matter. About 20-30 mg of crushed rock was heated in a pyrolysis oven under nitrogen flow followed by combustion under air flow following the methods in Lafargue et al. (1998), and Behar et al. (2001). 2.5 Organic Extraction Bitumen was extracted from 160 g of powdered rock for each core. The powdered rocks were placed into Teflon reactors, and the bitumen was extracted with a mixture of organic solvents (9:1, dichloromethane: methanol) at an abundance of 7 mL/g rock. The samples were extracted using a microwave accelerated solvent extractor (MARS5 Express, CEM Corporation, USA) at 120°C for 30 minutes. The process was repeated twice to ensure complete extraction of the bitumen. The extracted bitumen was then dried with anhydrous magnesium sulfate (activated at 120°C overnight), filtered and then evaporated by a rotary evaporator. Asphaltene separation from bitumen was achieved by dissolution of the bitumen with n-pentane (50:1 bitumen weight) overnight. The asphaltene fraction was washed several times with pentane and then dried and weighed. The maltene fraction was evaporated and weighed. 8

2.6 Sulfur speciation of the residual rock after bitumen extraction The residual rock after bitumen extraction underwent S species separation and quantification following Tuttle et al., (1986) and Acholla and Orr, (1993). Six bulk phases of S were quantified and measured for their 34S values: Bitumen S, water soluble sulfate (SO42-(water)), acid soluble sulfate (SO42-(acid)), acid volatile sulfides (AVS), pyritic S (PS) and kerogen S (KS). The SO42-(water) was dissolved from the residual rock after bitumen extraction, in triple distilled water for 2 hours. Then the acid volatile sulfides (AVS), in the residual rock after bitumen and SO42-(water) extraction, were reduced by hydrochloric acid (6 N). The evolved H2S from the reduction process was precipitated as Ag2 S, filtered, washed, dried and weighed. The SO42-(acid) was precipitated as barium sulfate (BaSO4), from the filtrate of hydrochloric acid (6 N) of AVS treatment, filtered, washed, dried and weighed. The rest of the sulfides in the residual rock after the AVS treatment were reduced by mixture of chromium (II) solution (1 N in 0.1 N solution of HCl) and concentrated hydrochloric acid (Canfield et al., 1986; Fossing and Jørgensen, 1989). This chromium-reduced sulfide (CRS) is predominantly composed of pyrite (hereafter the term pyrite will be used instead of CRS). The evolved H2S from the reduction process was precipitated as Ag2S, filtered, washed, dried and weighed (Fig. 3). Kerogen S was separated and quantified according to a modified method after (Levaggi and Feldstein, 1963). The organic S in kerogen (the residual solid remaining after the removal of the sulfides and sulfates) was oxidized, in a closed system, to sulfate using an oxygen flask. Triple-distilled water (10 mL) was added to the oxygen flask and then purged with oxygen for 15 minutes. Then 20-30 mg of kerogen or bitumen was mixed with sodium peroxide and folded into a paper that was fixed on a platinum flag. The edge of the paper was ignited and immediately introduced into the flask. The capped flask was kept closed for 30 min to ensure

9

complete oxidation of the sulfur and dissolution of the sulfates in the water. The flask was then opened and the water inside boiled to expel dissolved inorganic carbon. Finally, a few drops of concentrated hydrochloric acid (12 N) were added to the solution to get pH of 3-4, and then the sulfate was precipitated as barium sulfate (Fig. 3). The error associated with determining S concentration gravimetrically, in all the previous S extraction methods, ranged between 0-13% (RSD), and averaged 5.4% based on replicate analyses. 2.7 Stable isotope analysis Sulfur isotope (34S) analysis of the S samples (organic and inorganic) were measured by a continuous-flow method using a Flash EA 1112 series elemental analyzer coupled to a Delta Plus mass spectrometer (ThermoFinnigan, Germany) following the methods of (Giesemann et al., 1994). All 34S values are averages of replicate analyses (generally, n=3-5) with precisions better than 0.4‰. The 34S values were calibrated using a calibration curve with the international standardsNBS-127, IAEA-S-1, and IAEA-SO-6 (see details in section 3.2). 34S values are expressed in permil (‰) relative to the Vienna Canyon Diablo Troilite (VCDT) standard according to the following equation: AX= ((R (sample)/R (STD))-1

(1)

Where the X is the chemical symbol for the element (S, O and C) and A is the mass number of the rare isotope (34 or 18 or 13) and R is the isotope ratio of 34S/32S, 18O/16O or 13C/12C. Oxygen isotope (18O) analysis of sulfate was performed with a continuous flow isotope ratio mass spectrometer (20-20 Sercon Ltd., UK) via pyrolysis of BaSO4 to CO in a graphite crucible in a temperature conversion element analyzer (TC/EA, Sercon Ltd., UK) following the procedure by Boschetti and Iacumin (2005). All 18O values are averages of replicate 10

analyses (n=2-3) with precisions better than 0.7‰. The oxygen isotope values of the sample were calibrated using a calibration curve of

18

O/16O ratio of NIST standards and their δ 18O

values; (NBS-127 (BaSO4; δ18O = 8.7‰) and IAEA-SO-6 (BaSO4; δ18O = -10.4‰) (Halas et al., 2007). 18O values are expressed in permil (‰) relative to the Vienna Standard Mean Ocean Water (V-SMOW) standard according to equation 1. Carbon isotope ratios (13C) of kerogen and bitumen were measured by a continuous-flow method using a Flash EA 1112 series elemental analyzer coupled to Delta Plus XL mass spectrometer (ThermoFisher Scientific, Germany). All 13C values are averages of replicate analyses (n=2-3) with precisions better than 0.2‰ except of one sample that had a precision of 0.4‰. The 13C values were calibrated with the international standardSucrose (C12H22O11; 13C = -25.5‰), Glycine (C2H5O2 ; 13C = 31.97‰) and IAEA sucrose C6 (C12H22O11; 13C = -10.8‰). 13C values are expressed in permil (‰) relative to the Vienna Pee Dee Belemnite (V-PDB) standard according to equation 1. 2.8 Iron speciation and quantification Total Fe analysis was performed by ashing approximately 0.6 g sample in the muffle oven for 8 hours at 450°C, followed by heating to near boiling for 24 hours with 10 mL of 6 N hydrochloric acid (Aller et al., 1986). The hydrochloric acid extract was diluted by Milli-Q water. Iron concentrations were measured by a spectrophotometer (LaMotte Smart Spectro spectrophotometer, USA; minimum detection limit (MDL) was 0.07 ppm) at a wavelength of 562 nm using the ferrozine method (Stookey, 1970). A sequential extraction scheme (Table S.1; Poulton and Canfield, 2005) was used to separate highly reactive Fe species, which are characterized by different bioavailability and reactivity towards hydrogen sulfide. The highly reactive Fe fraction was defined as Fe minerals that can react or had reacted with hydrogen sulfide to form iron sulfides (ultimately as pyrite; FeS2). Reactive Fe content was 11

measured in the following Fe pools: 1) Fe carbonates (siderite, ankerite) extracted using sodium acetate (1N); 2) easily reducible Fe (oxy)-hydroxides (ferrihydrite, lepidocrocite) extracted using hydroxylamine-HCl in 25 wt% acetic acid; 3) reducible Fe (oxy)-hydroxides (akaganeite, hematite, goethite) extracted using sodium dithionite (0.29 N); 4) poorly reactive Fe oxides (magnetite) extracted using ammonium oxalate/ oxalic acid. Wet sediment (0.6 g) was placed in a 15 mL falcon tube with 10 mL of an appropriate extractant under constant agitation and oxic conditions. Solutions were diluted by Milli-Q water and measured spectrophotometrically with the same technique as for total Fe content. The Fe content for the same sample was measured in different matrix concentration, most and least concentrated matrix has been used, to estimate the matrix effect. The error (RSD) derived from the matrix effect, for measuring Fe content spectrophotometrically, was better than 1.4%. The organic Fe in the bitumen fraction was dissolved in a mixture of concentrated nitric acid and hydrogen peroxide (30%) following Altundag and Tuzen, (2011). After four hours, this mixture was evaporated and then re-dissolved in nitric acid (1 N). Calibration standards were prepared in the same way using the same matrix of the samples. Iron concentration was measured and quantified using inductively coupled plasma mass spectrometry (ICPMS, Agilent 7500cx, USA). Duplicates analysis of four of our samples showed relative standard error (RSD) that ranged 1-11% and averaged 5% for the Fe speciation method. This is similar error to that reported by Poulton and Canfield (2005). 2.9 Scanning Electron Microscopy (SEM) of pyrite Back-scattered electron (BSE) and secondary electron images (SEI) of thin sections of the studied core samples were obtained using a JEOL 6400 SEM (JEOL Ltd., Japan). The electron energy used was 15 keV. 3. Results

12

3.1 Bulk geochemical parameters Rock Eval data are presented in Table 1. The hydrogen index ranges between 623 and 949 mg/gTOC. The T max ranges between 402 and 415°C, and in the lowermost point of the section (595 m) reaches 422°C. Both the hydrogen index and Tmax are consistent with an immature source rock. The trends of TOC and TS are similar along the section and reach maximum concentrations of 22.4 wt% and 2.9 wt% at depths of 420 m and 393 m, respectively (Fig. 4a). Then, the TOC and TS content gradually decreases to minimum values (7.3 wt% and 0.4 wt%, respectively) at a depth of 595 m (Fig. 4a). Bitumen constitutes up to 0.8 wt% of the core samples at the depth of 528 m. The ratio of asphaltenes to maltenes ranges between 0.9 and 1.7 at depths of 498 m and 528 m, respectively (Table 1). However, for most of the core samples, asphaltenes and maltenes almost equally contribute to the bitumen. 3.2 Sulfur speciation and isotopes The distribution of different S forms is presented in Fig. 4b. The S occurs predominantly as kerogen S (KS), which represents up to 92.8 wt% of the TS at depth of 528 m (Fig. 4b). The second largest reservoir of S in the residual rock is pyrite, up to 28.4 wt% of the TS at the depth of 308 m (Fig. 4b). Water-soluble sulfate (SO42-(water)) is the third largest reservoir of S in the residual rock, up to 19.3 wt% of the TS at the depth of 498 m (Fig. 4c). The other S forms (AVS, acid soluble sulfates (SO42-(acid)) and bitumen S) represent only few percent of the TS (Fig. 4c). The 34S values of the S phases are presented in Fig. 5a and Table 2. The pyritic 34S values ranged between -39.5‰ at the depth of 308 m and -9.8‰ at 420 m. The 34S values of KS ranged between -8.7‰ at the depth of 450 m and 3.4‰ at the depth of 265 m (Fig. 5a). The general trend of KS 34S values is similar to that for the pyritic S (Fig. 5a). Bitumen S, 13

asphaltene S, maltene S and KS generally have similar δ34S values within 1‰ (Table 2). Kerogen S is consistently

34

S-enriched relative to coexisting pyrite, by up to about 38‰ at

308 m. This gap reduces to 11‰ at 420 m (Fig. 5b). Sulfate (SO42-(water) and SO42-(acid)) 34S values have similar trends to each other as well as to pyrite and KS (Fig. 5a). The 34S values of SO42-(water) and SO42-(acid) range between minima of -21‰ and -4‰ at 498 m and maxima of -2‰ and 5.9‰ at depths of 265 m and 595 m, respectively (Fig. 5a). The δ34S values of acid volatile sulfides (AVS) range between -10.6‰ at 498 m and 3.8‰ at 265 m (Fig. 5a). We examined oxidation of pyrite and/or organic S during sampling and lab treatment as an explanation for the relatively depleted δ34S values of the sulfates. Oxidation of the various sulfide forms during handling and storage may lead to the formation of elemental S as well as sulfate that is 34S-depleted relative to marine sulfate. However, to minimize such artifacts, our routine sampling procedure includes removal of the outer part of the core to exclude rocks that were exposed to air. Moreover, over the course of two years, we repeated the sampling process three times and found that the δ34S values of SO42-(water) did not change significantly (0.3-0.6‰), and that replicate concentration analyses of SO42-(water) have RSD of between 0.5 and 13%. Moreover, no elemental S was detected in the samples, which further support our conclusion that the sulfate in our samples did not form by oxidation of reduced S species during handling and storage of the cores. 3.3 Carbon and oxygen isotopes The 13C values of kerogen and bitumen are presented in Fig. 6a. The 13C values range between -29.0‰ at a depth of 420 m and -29.9‰ at a depth of 308 m for kerogen, and between -29.0‰ at a depth of 420 m and -31.7‰ at a depth of 450 m for the bitumen (Fig. 6a).

14

Oxygen isotope ratios were measured in two phases of sulfate, SO42-(water) and SO42-(acid). The general trends of the sulfate phases are similar, with an exception at 420 m, and both show only modest changes in 18O values along the section. The 18O values of the SO42-(water) ranged around -0.5±1.9‰ and those of the SO42-(acid) were around 10.6±2‰ (Fig. 6b). 3.4 Iron speciation Total Fe content (TFe) in the core samples ranged between 0.5 and 2.2 wt% at the depths of 528 m and 341 m, respectively (Fig. 7a). The highly reactive Fe (FeHR, namely Fe carbonates, easily reducible Fe oxyhydroxides, reducible Fe oxyhydroxides and pyritic Fe) comprises 1837 wt% of the total iron (Fig. 7b), and its distribution in the section is presented in Fig. 7c. Generally, the Fepy/FeHR ratio is uniform along the section with an average of 0.42 except for the 308-341 m depth interval, where the Fepy/FeHR reaches maximum values of 0.65 and 0.61, respectively (Fig. 7d). The organic Fe (Fe-Org) content has been measured in the bitumen fraction and normalized to the TOC, with the assumption that the kerogen has similar percentage of Fe as the bitumen (Hirner, 1987; Nwachukwu et al., 2000). The Fe-Org content is uniform along most of the section with an average of 0.04 wt% of the total Fe, and it reaches a maximum value of 0.18 wt% of the total Fe at the depth of 420 m (Fig. 7e). Pyrite in the core is framboidal, as determined by SEM analysis prior to sample treatment (Fig. S1). The pyrite grain diameter of the studied samples ranges from 1.3 to 1.6 m, and the aggregate diameter ranges from 5 to 14m. The grain and aggregate size is not regular or uniform in the same core sample. The aggregate crystal morphology appears in both regular, non-regular packing, spheroidal and flattened (originally spheroidal) framboids (Fig. S1). 4. Discussion

15

The combined S speciation, Fe speciation, and S and O isotope data from the Shefela Basin, Aderet borehole (hereafter “Aderet core”) measured in this study permit distinction between S- and Fe-bearing phases formed during early diagenesis (e.g., pyrite, organic S), and those that formed by late-stage oxidation (e.g., water-soluble sulfate). We start by discussing the evidence for a late-stage oxidation origin of the water-soluble and acid-soluble sulfate, and of the hematite in the core samples and then discuss the patterns observed in the speciation and isotopic composition of phases of a probable primary origin. 4.1 A late-stage oxidation origin of water- and acid-soluble sulfate The concentrations of SO42-(water) and SO42-(acid) in the Aderet core are high, and they represent a significant portion of the total S (up to 19.3 wt%). The 34S values of the sulfates are lower than marine sulfate of the Late Cretaceous (17-20‰ , Claypool et al., 1980) by up to 40.8‰ and 25‰ for SO42-(water) and SO42-(acid), respectively (Fig. 5a). As sulfate is enriched in 34S by the activity of sulfate-reducing microbes in sediments, the 34S-depeleted SO42-(water) and SO42(acid)

in our samples are unlikely to be residual to MSR. Instead, 34S values of SO42-(water) and

SO42-(acid) that are between those of organic S and pyrite (Fig. 5a) suggest an origin of the sulfates by oxidation of these reduced S compounds. Artifacts during sampling and handling are unlikely, as discussed above (Section 4.2), and we suggest that the sulfates were formed by natural oxidation of reduced S species present in the rock. Oxidation of aqueous reduced S compounds (e.g., H2S) during early diagenesis cannot be ruled out, although the low FePy/FeHR ratios, as reflected by the large proportion of Fe(III) oxyhydroxides relative to pyrite in the FeHR fraction (Section 4.4; Fig. 7), and environmental chemical proxies indicating a prolonged anoxic-dysoxic depositional environment (Meilijson et al., 2014), suggest that late-stage, moderate oxidation of pyrite and organic S is a more likely source of the 34S-depeleted sulfates (discussed below).

16

The difference in 34S and 18O values between SO42-(water) and SO42-(acid) (Fig. 8), as well as the different modes of occurrence of these two types of sulfate, may point to two different sources of S and times of sulfate formation. While SO42-(acid) is carbonate-associated sulfate (CAS), which substitutes for carbonate groups in the calcium carbonate matrix of the Aderet core, SO42-(water) occurs as water soluble sulfate minerals, which do not form in normal seawater or in sediment porewater other than in restricted, evaporative basins. We suggest, therefore, that the SO42-(acid) formed during early diagenesis, and was incorporated into authigenic carbonates, and that the SO42-(water) formed during a later oxidation event. The SO42-(acid) is

34

S-enriched relative to the coexisting SO42-(water) by 9-18‰ and is very

similar in isotopic composition to the kerogen S (Fig. 5a). This suggests oxidation of organic S or AVS to form SO42-(acid). The SO42-(water), on the other hand, displays 34S values between those of pyrite and organic S (Fig. 5a), suggesting that this sulfate may represent a mixed source of these two reduced S reservoirs. The fact that 18O values of SO42-(water) are around 0‰ while those of SO42- (acid) are around 10‰ (Fig. 6b) further suggests that these sulfate phases were probably generated at different times, out of solutions with different 18O values, and with different degrees of oxidative S cycling. Compelling evidence for a late oxidation event comes from the presence of hematite (Fe2O3). This Fe oxide is unexpected in an anoxic, organic-rich sediment, with intensive MSR, as it is reactive towards sulfide on relatively short timescales (Berner, 1981; Hu et al., 2006). Therefore, we find it likely that the hematite is a product of pyrite oxidation, which also produces sulfate according to the schematic net reaction: 15O2 (aq) + 4FeS2 (pyrite) + 8H2O  2Fe2O3 (hematite) + 8SO42- + 16H+,

(3)

This reaction generates acidity, and is likely to lead to dissolution rather than precipitation of carbonates, and so the sulfate produced should remain in water-soluble rather than acid17

soluble form. Oxidation of pyrite and organic S are known to be associated with a small S isotope fractionation (~ -0.7‰; Balci et al., 2007; Taylor et al., 1984; Toran and Harris, 1989; Pisapia et al., 2007; Heidel et al., 2011). Thus, the SO42-(water) produced according to this pathway should record 34S values between those of pyrite and organic S, depending on the relative proportions of these end members. We carried out mass balance calculations of 34S in SO42-(water) using the measured concentrations of hematite, SO42-(water) and pyrite. We assumed that hematite was derived solely from pyrite oxidation and that the pyrite S was completely transformed to SO42-(water). We assume a closed system and no S isotope fractionation during pyrite oxidation. The molar difference between hematite (i.e., pyrite oxidation) and SO42-(water) is assumed to represent the organic S oxidation: 34S(SO42-(water)) = X* 34S(Pyrite) + (1-X)*34S(OM),

(4)

where X is the fraction of SO42-(water) derived from oxidized pyrite (constrained by the abundance of hematite iron), 34S(Pyrite) is the S isotope ratio of pyrite, and 34S(OM) is the S isotope ratio of the organic S (kerogen S). Good correlation between the calculated and measured 34S values of SO42-(water) (R2=0.89) supports its formation from the oxidation of pyrite and organic S (Fig. 9). The oxidation of organic S to SO42-(water) dominates in the depth interval between 265 and 341 m. Deeper in the section, a larger fraction of sulfate comes from pyrite oxidation than from organic S oxidation (Fig. 10a). However, throughout the core, the oxidation of organic matter was minor relative to the original content, as deduced also from the low oxygen index (OI) values <40 mg CO2/g TOC (Table 1). Calcite veins observed in the Aderet core, with textures indicative of a late origin, possibly related to flushing of the rock with fresh water, support our suggestion for a post-depositional oxidation event. For the same Aderet core, Burg and Gersman, (2016) reported water that was fresh relative to well water in the Shefela Basin (Nahal Guvrin new and Nahal Guvrin 18

old wells). The fresh water is thought to have flushed this sub-basin (i.e., the Aderet core) during a post-depositional change in its tectonics (Burg and Gersman, 2016), and we suggest that it may have contained dissolved oxygen that oxidized organic matter and various reduced S phases. Several other studies have shown evidence for post-depositional oxidation of organic matter and pyrite in other basins in the same geographic region. Examples for such events come from the Amzaq-Hazra sub-basin in central Jordan, and from the southern Tethyan Paleocene-Eocene Thermal Maximum (PETM) exposures in Egypt (Aubry et al., 2007; März et al., 2016). In these sites, the oxidation events were interpreted to reflect tectonic processes that exposed the rock to chemical remobilization and/or physical reworking (Aubry et al., 2007; März et al., 2016). The analysis of S isotopes in different sulfate phases, as conducted in the present study, can reveal oxidation events that may not be identified otherwise. In some cases the oxidation alters the abundances of pyrite and kerogen S, which can lead to erroneous interpretations of the S cycle in the basin. Using the abundance of hematite and the 34S values and the abundance of water-soluble sulfate to estimate the amount of pyrite and organic S oxidized, under the assumption of a closed system, we reproduced the pyrite and organic S 34S values before the oxidation events in the Aderet core (Fig. 10b). It can be seen that despite these oxidation events and changes in pyrite and (less so) in kerogen S content, their original concentrations did not change significantly. The negligible S isotope fractionations associated with the oxidation suggests that the isotopic profiles of pyrite and organic S remained likewise unaltered. Thus, although recognized throughout the few hundred meters of the studied En Zeitim Formation section, the effects of the recorded post-depositional oxidation on the concentration and S isotopic profiles is minor, lending confidence to interpretations of the isotopic trends.

19

4.2 Pyrite 34S values suggest large microbial fractionation The 34S values of the pyrite in the Aderet core are 34S-depleted by 30-63‰ relative to late Cretaceous marine sulfate (34S 17-20‰, Claypool et al., 1980) (Fig. 5a). As fractionation of S isotopes during sulfide mineral formation is negligible (Berner and Westrich, 1985; Butler et al., 2004; Hartgers et al., 1997), this large apparent sulfate-pyrite fractionation reflects the isotopic composition of porewater sulfide, and is controlled by the processes that govern the production and cycling of sulfide. The production of sulfide is predominantly due to MSR, and indeed the upper end of the range of sulfate-pyrite 34S difference in the Aderet core is similar to some of the largest microbial fractionations associated with MSR, as observed in laboratory cultures and natural environments (Canfield et al., 2010; Sim et al., 2011). However, the large apparent fractionation may be an underestimate of the actual fractionation associated with MSR. If sediment porewater is in poor communication with the overlying water column, as would be the case in rapidly depositing or organic-rich sedimentary environments, then drawdown of porewater sulfate accompanied by isotopic (Rayleigh) distillation, would lead to the formation of sulfide that was fractionated from seawater sulfate by less than the microbial fractionation associated with MSR (Aharon and Fu, 2000; Canfield., 2001). The chalks present in the Aderet core are thought to have deposited slowly (range 1-6.7 cm/ky, average ~4cm/ky; Meilijson et al., 2014), but as previously mentioned, these sedimentary rocks are extraordinarily rich in organic matter. Rapid sulfate drawdown should thus lead to “closed-system” behavior, distillation of S isotopes, and apparent fractionations that were smaller than the microbial fractionation. Indeed, the smallest sulfatepyrite 34S difference in the Aderet core (30-40‰) is observed in the most organic-rich section of the core (depths 398 m and 421 m).

20

Repeated cycles of sulfide oxidation to elemental S, disproportionation of the elemental S to sulfide and sulfate, and microbial reduction of the sulfate to sulfide may lead to progressive depletion of the sulfide in

34

S, leading to ultimate sulfate-sulfide (and sulfate-pyrite)

fractionations that are larger than those associated only with MSR (Canfield and Thamdrup, 1994; Habicht and Canfield, 2001; Sørensen and Canfield, 2004). Therefore, it is possible that the large apparent sulfate-pyrite fractionation observed in some of the Aderet samples reflects more modest microbial fractionation associated with MSR, augmented by the fractionation associated with disproportionation. However, the high organic matter content of the sediments in question is expected to have limited the importance of elemental S disproportionation in two ways. First, the availability of electron acceptors for oxidative cycling of S would have been severely limited. Unlike some organic-rich sediments, in which redox conditions vary on short timescales and lead to cycles of oxidation and reduction (Sørensen and Canfield, 2004), the redox conditions in the relatively deep-water, slowlydepositing, organic-rich Aderet sediments would have been relatively stable and reducing. Second, sulfide-rich conditions are expected to lead to rapid reaction of the elemental S to form polysulfides (Ferdelman et al., 1991; Filley et al., 2002; Heitmann and Blodau, 2006), which are highly reactive towards organic matter (LaLonde et al., 1987; Vairavamurthy and Mopper, 1989; Loch et al., 2002a; Aizenshtat and Amrani, 2004a), and would be scavenged by organic matter sulfurization on relatively short timescales. To summarize, we cannot rule out elemental S disproportionation in the Aderet core, but find it unlikely that it played a major role in generating the large observed sulfate-pyrite apparent fractionation in these sediments. The

34

S depletion in pyrite may be explained by the existence of syngenetic pyrite (i.e.,

formed in the water column), which is often 34S-depleted relative to diagenetic pyrite (Calvert et al., 1996; Lyons, 1997; Raven et al., 2016a). However, the water column of the studied 21

system was non-euxinic (Amit and Bein, 1982; Bein et al., 1990; Alsenz et al., 2015), and it is hard to envision syngentic pyrite formation or OM sulfurization within water column under these conditions. Nevertheless, the observation that all of the pyrite in the studied samples occurs as framboidal aggregates may suggest that pyrite formation occurred within sulfatereducing microbial aggregates or biofilms (Popa et al., 2004; MacLean et al., 2008; Wacey et al., 2015). Therefore, the isotopic composition of the pyrite may reflect that of the instantaneous products of MSR (MacLean et al., 2008; Wacey et al., 2015; Raven et al., 2016b), in a process that is similar in many respects to water-column pyrite formation. The large preserved apparent fractionation, regardless of its exact formation pathway (within biofilms or not), implies that i) microbial fractionation was large, and ii) in some parts of the core pyrite did not preserve the full distillation sequence of increasingly 34S-enriched sulfide produced by MSR. We discuss the latter inference in Section 5.3. The inference of large microbial fractionations associated with MSR is intriguing, given how rich the Aderet sediments are in organic matter (7-22 wt%), and given the expectation that sulfate reduction rates scale with the availability of the organic substrate (Westrich and Berner, 1984; Berner and Westrich, 1985; Aharon and Fu, 2000; Canfield., 2001). Microbial fractionation of S isotopes during MSR is inversely related to specific respiration rates (Aharon and Fu, 2000; Canfield, 2001), and a transition from large fractionations at low rates and small fractionations at high rates is understood to reflect the transition from near-equilibrium conditions (and fractionations) at low rates, to kinetically controlled conditions at high rates (Harrison and Thode, 1958; Wing and Halevy, 2014). Indeed, variation in preserved sulfatepyrite fractionation over Phanerozoic time has been suggested to result from variation in MSR rates (Leavitt et al., 2013). The observational support for large microbial fractionation during MSR in the Aderet core suggests that even under extreme organic-rich conditions, sulfate-reducing microbes may respire slowly and fractionate S isotopes strongly. It is unclear

22

whether this is a general phenomenon, or whether the high concentrations of organic matter in the Aderet sedimentary environment led to reducing conditions that limited the aerobic degradation of large, complex organic molecules to simple organic acids that are accessible to dissimilatory iron and sulfate reducers (Westrich and Berner, 1984; Berner and Westrich, 1985; Aharon and Fu, 2000; Canfield., 2001). A large and relatively invariant S isotope fractionation between sulfate and sulfide (represented by pyrite) may be challenged by the significant variations in the 34S values of pyrite along the Aderet core (Fig 5a). If microbial fractionation itself varied, a similar trend is expected in the 34S values of organic S. However, organic S 34S values vary by only ~12‰, compared to a variation in pyrite 34S values of ~30‰ (Fig. 5a). As we argue below, these observations may be explained by the dynamics of competing organic matter-sulfide, organic matter-iron, and sulfide-iron reactions. 4.3 Organic matter and reactive iron availability control the 34S values recorded in pyrite Reduced S species formed by MSR can react with reactive Fe and organic compounds in depositional environments to form both pyrite and organic S compounds (Goldhaber and Kaplan, 1974; Sinninghe Damste and De Leeuw, 1990; Rickard and Luther, 2007). The preserved S isotope composition of pyrite and organic S depends on the microbial fractionation during MSR and on the suite of processes that modulate that fractionation. As sulfate is reduced by MSR in the sediments, producing

34

S-depleted sulfide, the 34S of the

residual sulfate increases (Fig. 11). The rate of decrease with depth in the sediment of sulfate concentration and the rate of increase in its 34S values depends on the marine sulfate concentration (Gomes and Hutgen 2013; 2015), and the relative rates of MSR and diffusive exchange with the overlying water column. In more “open-system” conditions, for example, when sedimentation is slow, porosity is high, or the organic load is low, sulfate 23

concentrations and 34S values change slowly with depth. The isotopic composition of the sulfide produced at any given depth reflects the isotopic composition of the sulfate at that depth and the microbial fractionation of S isotopes during MSR (Fig. 11, instantaneous sulfide). The isotopic composition of the accumulated sulfide, however, additionally depends on the relative rates of its reaction and diffusion through the sediment (Fig. 11, accumulated sulfide). For example, in sediments where removal of sulfide is slow, perhaps due to limited Fe or organic matter availability, diffusion mixes sulfide that originated from different depths in the sediment and the isotopic composition of sulfide at any given depth departs from a constant offset from that of sulfate. The isotopic composition of porewater sulfide exerts a first-order control on the isotopic composition of pyrite, as fractionation associated with pyrite precipitation is thought to be small (Price and Shieh, 1979; Raab and Spiro, 1991; Worden et al., 1997; Butler et al., 2004). There are several factors, however, that influence the depth within the sediment at which pyrite forms, and therefore, the sulfide that it samples and preserves. Formation of the FeS precursor depends on the availability of Fe(II), and so, the location of pyrite formation and the ultimate amount formed, depend on where in the sediment Fe becomes available for FeS precipitation. Fe(II) in sediment porewater may come from microbial reduction of iron oxides or from chemical reduction by sulfide. These reactions are rapid relative to rates of sediment accumulation. Consequently, in sediments where iron oxides are in a lower abundance than the potential of sulfate to diffuse into the sediment and become microbially reduced, pyrite formation will be limited by Fe availability (Fig. 11, curves a and b). Unless there is a hindrance to microbial or chemical reduction of the iron oxides, pyrite will then form almost exclusively close to the sediment-water interface (immediately below the oxygen penetration depth), before the residual sulfate and instantaneous sulfide produced from it have become very enriched in 34S. As a result, the pyrite will be relatively 34S-depleted, as the 24

sulfide from which it formed, and the apparent fractionation between seawater sulfate and sedimentary pyrite (34SSulfate–Pyrite) will be large. The control of Fe oxide availability on pyrite 34S values, as described above, is expected also in the case of pyrite formation via a ferric hydroxide surface process (Wan et al., 2017), instead of precipitation of a FeS precursor from solution and reaction of the precursor with sulfide or polysulfide (Rickard and Luther, 2007). In sediments where iron oxides are non-limiting, pyrite may form over the entire range of sulfide 34S values produced by isotopic distillation during MSR (Fig. 11, curve c), including those enriched in 34S, and 34SSulfate–Pyrite will be smaller. The large 34SSulfate–Pyrite we observe in the Late Cretaceous organic-rich chalk section studied in the Aderet core may be consistent with Fe limitation, which led to preservation of sulfide produced only in the upper part of the sulfate distillation profile, close to the sediment-water interface. There are several lines of evidence in support of Fe limitation. First, S occurs predominantly in kerogen in these rocks, up to 92.8 wt% of the total S (Fig. 4b). Second, the S content in pyrite and kerogen show an inverse trend (Fig. 4b) and FePy/FeHR ratios are almost uniform throughout the studied section (Fig. 7d). Third, the ratios of total S to total Fe (1.1) and total S to reactive Fe (3.6) suggests an excess of S relative to Fe in the depositional system (Fig. 12a and b). This excess S, which did not react with Fe, apparently reacted with organic compounds, as observed in previous studies (Dean and Arthur, 1989; Zaback and Pratt, 1992). However, there are aspects of the Aderet core, which suggest that factors other than Fe limitation may have contributed to the observed patterns. Mass balance calculations show that if all the reactive Fe reacted with the reduced S available, much more S and Fe would have been precipitated as pyrite; about 6.3 to 44.3 wt% of the TS (FePy/FeHR ~0.9) instead of the current amount of 2.8 to 28 wt% of the TS (FePy/FeHR ~0.4). Assuming that the hematite in the samples was formed by late-stage oxidation of pyrite (Section 5.1, see also

25

Hu et al., 2006), pyrite would still constitute only 4.6 to 32 wt% of the TS (FePy/FeHR ~ 0.50.8). Faster reaction kinetics of reduced S with reactive Fe species than with organic compounds (Gransch and Posthuma, 1974; Hartgers et al., 1997) should result in the preferential formation of Fe sulfides at concentrations of reactive Fe and organic matter commonly observed in marine sediments. However, several studies have shown that during early diagenesis, Fe sulfides may form simultaneously with organic S (Bates et al., 1995; Bruchert and Pratt, 1996; Urban et al., 1999; Filley et al., 2002), and some recent studies suggest that organic S formation can compete with pyrite formation for reduced S species (Werne et al., 2003; Riedinger et al., 2017). Other studies suggest that Fe sulfide may promote organic S formation through the production of reactive S species such as polysulfide anions (Bates et al., 1995; Bruchert and Pratt, 1996; Urban et al., 1999; Filley et al., 2002). Raven et al., 2015 showed that in the Cariaco Basin some specific organic S compounds can be relative to pyrite, while the bulk organic S was

34

34

S-depleted

S-enriched compared to pyrite. This may

point to reaction of some organic compounds with reduced S species shallower within the sediments (Fig. 11) than the reaction of S with Fe species to form Fe sulfides. Formation of pyrite later than organic S may be the result of faster reaction kinetics of some organic compounds with reduced S, very high abundances of organic matter, or some hindrance to Fe sulfide formation. Our observations from the Aderet core suggest that one or more of these mechanisms may have operated in the Late Cretaceous organic-rich chalks, as described below. The negative correlations between pyritic Fe and TOC and between pyritic and organic Fe suggest that the pyrite formation was linked to OM content (Fig. 13a and b). The interaction between functionalized organic molecules and Fe species (both Fe(II) and Fe(III)) is known to form strong organic-Fe complexes such as Fe porphyrins (Oldham and gloyna, 1969; 26

Theiss and singer, 1974; Davis and Leckie, 1978; Davis, 1982; Eckardt et al., 1989; Gu et al., 1995; Baalousha et al., 2008). In this way, organic-Fe complex formation may decrease the concentration of reactive Fe in organic-rich sediments (Rose and Waite, 2003; Colo'n et al., 2008; Li et al., 2009; Aeschbacher et al., 2012). Thus, we suggest that just as the kinetics of OM-reduced S reactions may outcompete those of reactive Fe-reduced S reactions, so may the kinetics of reaction between organic compounds and reactive Fe compete with the reduced S for the reactive Fe. As more OM entered the system, more Fe may become incorporated into organic compounds, form aqueous complexes with dissolved OM, or sorb to particulate OM, and become inaccessible to react with reduced S to form pyrite (Canfield et al., 1992; Tribovillard et al., 2015; Turchyn et al., 2016). The 34S values in pyrite appear to be consistent with such competition, as there is a positive correlation between the 34S values of pyrite and the TOC (Fig. 13c). In this scenario, competing kinetics of Fe reaction with OM and reduced S prevented the formation of pyrite early in the distillation sequence, when sulfide 34S values were at their lowest, and led to formation of pyrite deeper in the sediment, from sulfide that was 34S-enriched relative to the initial sulfide (Fig. 11). A later formation of pyrite may be explained by the decrease in the concentration of functional groups as OM diagenesis progresses (Durand, 1985). Some of these functional groups act as ligands, which complex the Fe species (Rose and Waite, 2003), and the fraction of Fe bound by these complexes depends mainly on the ligand strength and concentration (Hutchins et al., 1999; Martinez et al., 2000; Rose and Waite, 2003; Garg et al., 2007). Therefore, at later stages of diagenesis, when the concentration of functional groups (i.e., ligands) in organic compounds decreases, the fraction of Fe bound by them is expected to decrease as well. Consequently, reactive Fe is gradually released, making it available for reaction with H2S or other sulfide species to form iron sulfides, including pyrite. Thus, in concert with the availability of total reactive Fe, the timing in the distillation sequence at 27

which this Fe became available for reaction with reduced S species probably dictated the 34S of pyrite along the Aderet core, especially in the more organic-rich parts of the section. This process is shown schematically in Fig. 14. 4.4 The relationship between the isotopic composition of kerogen S and iron sulfides Organic S is, on a global average, enriched in

34

S relative to coexisting pyrite by about 10‰

(Anderson and Pratt, 1995), though enrichments of up to ~30‰ are observed in some basins (Raiswell et al., 1993; Anderson and Pratt, 1995; Amrani et al., 2005; Zhu et al., 2013). The δ34S values of the bulk organic S in the Aderet core are consistently enriched relative to coexisting pyrite, with a variable difference (34SKerogen-Pyrite) reaching a maximum of 37.8‰ at 308 m, and then decreasing to 11‰ at 420 m (Fig.5b). These 34SKerogen-Pyrite variations are mainly due to changes in the pyrite δ 34S values, whereas kerogen δ34S values remain relatively constant (Fig. 5a). The largest 34SKerogen-Pyrite values observed in the Aderet core are greater than previously reported values, and this kerogen

34

S enrichment may be

explained by several processes that act individually or in combination, as discussed in (Amrani, 2014). Assimilatory organic S (biosynthetic), which is enriched in

34

S relative to

organic S of a dissimilatory (MSR) origin, can make up to 20-25% of the total sedimentary organic S (Anderson and Pratt, 1995; Bruchert and Pratt, 1996; Canfield et al., 1998). However, assimilatory S probably constitutes a much smaller proportion of total organic S in environments with high rates of MSR and dissimilatory sulfide production (Anderson and Pratt, 1995; Aizenshtat and Amrani, 2004a). In such environments (including the Aderet core), the S content of organic matter can reach 10-12 wt%, which is a factor of ~10 higher than the assimilatory S content of living cells. In addition, biosynthetic S compounds (e.g., proteins) are usually chemically and biologically labile and are expected to rapidly mineralize. Therefore, we do not consider assimilatory S a sufficient explanation for the high organic S δ34S values in the Aderet core. 28

Equilibrium S isotope fractionation during S incorporation into organic compounds can enrich the organic S in

34

S by 4-8‰ (Amrani and Aizenshtat, 2004c; Amrani et al., 2008a).

Furthermore, reduced S species involved in isotopic exchange with the organic S may also affect the isotopic composition of the latter. For example, polysulfide anions (mainly Sn2-) are 34

S-enriched relative to coexisting H2S by 3-6‰ (Amrani et al., 2006). Polysulfide anions are

better nucleophiles than H2S species (mainly HS-), and are therefore suggested to react preferentially with functionalized organic compounds ( Lalonde et al., 1987; Vairavamurthy and Mopper, 1989; Loch et al., 2002; Amrani and Aizenshtat 2004; Wu et al., 2006). Organic S in equilibrium with polysulfide anions will thus be 34S-enriched relative to pyrite by up to 10‰ (Amrani et al., 2008a). Hence, equilibrium isotope fractionation cannot fully explain the large kerogen-pyrite S isotope differences in the Aderet core, if the organic S and pyrite formed from the same reduced S reservoir. We suggest that the observed values of 34SKerogen-Pyrite in the Aderet core (≤~38‰) reflect Fe-limited formation of pyrite in the uppermost sediment (Section 5.3), and formation of organic S throughout the sedimentary column. In the upper portion of the sediment, where both organic matter and reactive Fe are available for reaction with sulfide produced from MSR, both organic S and pyrite will incorporate the early, most

34

However, deeper in the sediment, as isotopic distillation enriches in

34

S-depleted sulfide.

S both the residual

sulfate and the sulfide produced from it during MSR (Fig. 11), reactive Fe runs out and pyrite precipitation ceases, but the high concentrations of organic matter allows its continuous sulfurization. Thus, the organic matter will gradually incorporate more

34

S-enriched reduced

S species and become enriched in 34S relative to the pyrite. Irreversible incorporation of reduced S species into some organic compounds may be associated with a kinetic isotope effect (KIE), which in most cases should lead to more

29

34

S-

depleted organic S compounds than the coexisting pyrite (Brüchert and Pratt, 1996; Raven et al., 2015). However, laboratory experiments could not demonstrate such KIEs other than for sulfurization of alkyl halides, which are not abundant in natural sediments (Amrani and Aizenshtat, 2004c; Amrani et al., 2008a). Moreover, the experimental KIE was in the range of 2-3‰ (Amrani et al., 2008a) and cannot explain the 10-15‰ apparent fractionation between some individual organic S compounds and pyrite observed in the upper part of the Cariaco Basin sediment (Raven et al., 2015) or the ≤38‰ apparent fractionation we observe in the Aderet core. An alternative explanation is that some organic compounds are more reactive than Fe toward reduced S species and will thus record the δ34S of sulfides in the early distillation curve (Fig.11). If uptake of S by these compounds is irreversible they can preserve δ34S values obtained near the sediment-water interface, even as they become progressively buried deeper within the sediment. Irrespective of whether the

34

S-depleted organic S

compounds reflect a KIE upon sulfurization or early and irreversible uptake of reduced S close to the sediment-water interface, the fact that the bulk organic S δ34S values were high relative to pyrite throughout the whole Cariaco sediment core (Werne et al., 2003) suggests that these processes were of minor importance (i.e., small KIE or subsequent exchange between organic S and reduced S species in porewater). Indeed, Raven et al. (2015) showed that the isotopic difference between individual S compounds and pyrite decreases with depth in the sediment, and that organic S δ34S values eventually reach those of pyrite, suggesting that there are other controls on the OSC isotopic values as we discuss below. The isotopic composition of organic S and coexisting porewater sulfide are observed to covary with a slope of ~0.8 (Anderson and Pratt, 1995), suggesting that the majority of organic S is exchangeable with porewater sulfide and polysulfide species on relatively short timescales. Several studies have shown that hydrogen sulfides, polysulfides and acid-volatile sulfide (including iron monosulfide, but excluding pyrite) participate in S isotope exchange 30

with each other and with organic S (Fossing and Jorgensen, 1990; Rickard and Luther, 2007; Méhay et al., 2009). For example, rapid isotope exchange of polysulfide cross-linked polymer (synthetic analogue of protokerogen) with polysulfide and sulfide anions have been demonstrated within hours at ambient to moderated temperatures (Amrani et al., 2006). Thus, even once no more S is added to organic matter, S isotope exchange may lead to continued increase in organic S δ34S values, as δ34S values in residual sulfate and the sulfide produced from it increase down the isotopic distillation sequence (Fig. 11). Such S isotope exchange between reduced S species (other than pyrite) and organic S throughout the diagenetic process may also explain the relatively uniform 34S values of kerogen S along the Aderet core, compared to the variable δ34S values of pyrite, which does not exchange S isotopes with reduced S species. 4.5 Geochemical implications The δ34S values of pyrite and gypsum/anhydrite in marine sediments are thought to represent the 34S values of the sulfide and sulfate in sediment porewaters and the S isotope fractionation associated with MSR. Hence, δ34S measurements (mostly of pyrite) in marine sediments and sedimentary rocks are used to constrain the isotopic composition of seawater sulfate, the fractional pyrite burial and effects on the global paleo-redox budget, the activity of S-metabolizing organisms on a local and global scale, and diagenetic processes in marine depositional environments (Claypool et al., 1980; Canfield and Teske, 1996; Canfield., 2004; Johnston et al., 2005; Fike et al., 2006; Gill et al., 2007; Fike and Grotzinger, 2008; Hurtgen et al., 2009; Jones and Fike, 2013; Leavitt et al., 2013) We have shown in the present study, that the dynamics of interaction between organic matter, reduced S species and Fe is an important additional control on the 34S values preserved in pyrite. In any depositional environment (organic-rich or not), the preserved isotopic

31

composition of pyrite and organic S reflects interplay among reaction, transport and the availability of Fe and organic matter. Specifically, the controlling factors, as discussed here, are the integrated sum of the following processes: 1. the relative importance of reaction and transport, or the system “openness” with respect to sulfate and sulfide, which determines the δ34S profile of reduced S species in the sediment, 2. competition of organic compounds and Fe for the available reduced S species, which determines the depth within the profile of the sulfide that is ultimately sampled and preserved as pyrite and organic S, 3. the total availability of reactive Fe, and the timing at which this Fe becomes available for pyrite formation, 4. S isotope exchange between OSC and the inorganic reduced S pool, and 5. alteration of the original signal during late diagenesis and/or exhumation-related processes (e.g., pyrite oxidation). While some of these processes have been discussed in the literature, the effects of organic matter on the availability of reactive Fe and reduced S for pyrite formation, and therefore, on the 34S values preserved in pyrite and organic S has received much less attention. The mechanisms that were suggested to affect the availability of Fe include anaerobic methane oxidation, microbial Fe reduction, and alteration of Fe-rich minerals, such as smectites (Canfield et al., 1992; Tribovillard et al., 2015; Turchyn et al., 2016). These processes are thought to pace the release of reactive Fe, leading to the formation of secondary 34S-enriched pyrite at a later stage than the primary pyrite that formed during early diagenesis (Chappaz et al., 2014). We suggest a novel, potentially widespread mechanism, wherever a large flux of organic matter enters the depositional environment. Organic matter can trap reactive Fe via complexation or sorption (Fig. 14), and this limits the availability of Fe and prevents it from rapid reaction with reduced S species. At later stages of diagenesis when organic matter functionalities become progressively quenched, some of these complexes break and release reactive Fe back into sediment porewaters, where it reacts with 32

34

S-enriched reduced S

compounds, forms pyrite with relatively high δ34S values. Thus, we suggest that mutual organic-Fe-S competition may control the 34S values of pyrite in organic-rich environments. In such environments these factors should be taken into consideration when attempting to reconstruct the local and global S cycle from S isotope measurements in sedimentary pyrite. 5. Conclusions In this study we investigated S and Fe speciation in organic-rich chalk samples from the upper part of the Late Cretaceous rocks in the Tethyan Levant Basin to establish an improved understanding of the sedimentary S cycle and its link to Fe and organic matter. The mechanisms we suggest and discuss in this paper to explain the relationships between the isotopic compositions of sedimentary pyrite, organic S and seawater sulfate in organic-rich sediments are applicable to a wide range of modern and ancient depositional environments. The main conclusions are: 1. The analysis of sulfate (water and acid soluble) in the studied rocks allowed us to detect exposure to oxidizing fluids and to evaluate its impact on the preserved signals. All sulfate forms measured were 34S-depleted and do not represent the original marine sulfate of the Late Cretaceous. Water-soluble sulfate formed by late-stage oxidation of both pyrite and organic S, probably upon exposure to oxidized meteoric fluids, while acid-soluble sulfate represents the oxidation of reduced S species during relatively early diagenesis and incorporation into carbonate minerals. Despite these diagenetic and exhumation-related oxidation events, and related changes in pyrite and (less so) in kerogen S abundance, the trends of concentrations and δ34S values of pyrite and kerogen S did not change significantly. Therefore, our findings suggest that analysis of coexisting oxidized Fe and S species (in addition to the organic S and pyrite) is

33

necessary to distinguish early diagenetic from late-stage alteration, lending confidence to interpretations of the primary isotopic trends. 2. The 34S values of the pyrite are

34

S depleted by up to 63‰ relative to the source

marine sulfate as a result of the reaction of

34

S-depleted reduced S species with

reactive Fe. The large sulfate-pyrite δ34S difference is likely the result of large, nearequilibrium S isotope fractionation during microbial sulfate reduction. Such large fractionations indicate low rates of sulfate reduction despite exceptionally large abundances of organic matter. 3. The major form of S, by far, is organic S. Even when correcting for a late-stage oxidation event that decreased the abundance of pyrite in the studied section, it appears that some of the reactive Fe did not react to form pyrite. This may be explained by competition between reduced S species and organic matter for reactive Fe, which limited the availability of Fe for pyrite formation. When large fluxes of organic matter entered the sedimentary system, organic matter may have become a significant sink for Fe (through complexation, sorption or Fe incorporation), delaying the formation of pyrite until deeper in the sediment, and farther along the sequence of microbial sulfate utilization and isotopic distillation. 4. There is large

34

S enrichment of kerogen S relative to the coexisting pyrite (up to

38‰). This phenomenon can be explained by the relative timing of pyrite and organic S formation. While the amount of reactive Fe in the Aderet core is limited, organic matter is in large excess, and we suggest that it continued to react with reduced S species long after Fe was depleted. The isotopic composition of organic S represents either the 34S values of the cumulative H2S (with an offset) over a large part of the isotopic distillation sequence, or equilibration between organic S and

34

S-enriched

reduced S species relatively deep within the sediment. The pyrite, on the other hand, 34

appears to have incorporated only H2S produced early in the distillation sequence (34S-depleted), probably due to Fe limitation. Acknowledgments We are grateful to Harold Vinegar and Yuval Bartov from IEI Ltd. for providing the core material and for useful discussions. We appreciate the discussions with Sigal Abramovich, Ari Meilijson and Yoav Oved Rosenberg about the geology and stratigraphy of the studied area and the help with the core sample selection. We thank Alon Angert for the oxygen isotope analysis of sulfates. L.S. is grateful for PhD scholarchips from the Ministry of National Infrastructures, Energy and Water Resources of Israel, and the Harry and Sylvia Hoffman Fund. I.H. acknowledges Starting Grant #337183 from the European Research Council. A.A. Thanks the Israeli Science Foundation grant No. 1738/16 for partial support to this study. References Acholla F. V. and Orr W. L. (1993) Pyrite removal from kerogen without altering organic matter: the chromous chloride method. Energy & Fuels 7, 406–410 . Aeschbacher M., Brunner S.H., Schwarzenbach R.P. and Sander M. (2012) Assessing the effect of humic acid redox state on organic pollutant sorption by combined electrochemical reduction and sorption experiments. Environ. Sci. Technol. 46, 3882– 3890. Aharon P. and Fu B. (2000) Microbial sulfate reduction rates and sulfur and oxygen isotope fractionations at oil and gas seeps in deepwater Gulf of Mexico. Geochim. Cosmochim. Acta 64 (2), 233-246.

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53

Fig. 1. a) Paleogeograhic reconstruction of the Tethyan region in the Late Cretaceous with the major upwelling belts (Ashckenazi-Polivoda et al., 2011) , and b) location map of the Shefela Basin and the Aderet borehole (modified after Minster T., 2009). Fig. 2. The stratigrapic column of the studied Late Cretaceous En Zeitim Formation (stratigraphically equivalent to the Ghareb Formation ) in the Aderet borehole (modified after Meilijson et al., 2014). Fig. 3. Analytical scheme used to separate the various forms of sulfur from the Aderet core samples. Fig. 4. Depth profiles in the Aderet core: a) total organic carbon (TOC) and total S (TS) content, b) distribution of kerogen and pyrite sulfur, presented as wt% of TS, c) distribution of other bulk phases of sulfur. Fig. 5: Depth profiles of δ34S values of different sulfur species in the Aderet core: a) kerogen S (KS), pyrite (PS), water soluble sulfate (SO42-(water)) ,acid soluble sulfate (SO42-(acid)) and acid volatile sulfide (AVS), b) 34S difference between pyrite and kerogen (34S(pyrite-Kerogen)) Fig. 6. Depth profiles of isotopic ratios in the Aderet core a) δ13C values of kerogen and bitumen , b) δ18O values of water soluble sulfate (SO42-(water)) and acid soluble sulfate (SO42(acid)).

Fig. 7. Depth profiles of iron concentration and speciation in the Aderet core: a) Total iron content (TFe), b) highly reducible iron (FeHR) content relative to the total iron content, c) iron content in different highly reducible iron phases, d)FePy/ FeHRratio, e) organic iron content (Fe-Org) as wt% of TFe. Fig.8. Cross plot of δ34S and δ18O values for sulfate phases. Filled green diamonds for water soluble sulfates (SO42-(water)) and open brown diamonds for acid soluble sulfates (SO42-(acid)). Fig. 9. Correlation between the calculated δ 34S values of water soluble sulfates (SO42-(water)) according to Eq. 4, and their measured values.

54

Fig. 10. The effect of secondary oxidation on iron and sulfur speciation in the Aderet Core. a) calculated fraction of sulfate (f-values) comes from pyrite oxidation and kerogen oxidation in the Aderet core samples, b) reconstruction of the pre-oxidation content of pyrite and kerogen sulfur relative to their measured present values in the Aderet core samples. Fig. 11. Schematic illustration of the isotopic composition of porewater sulfate, sulfide, pyrite and organic S below the oxygen penetration depth. Sulfate reduction produces sulfide and leaves residual sulfate that is enriched in

34

34

S-depleted

S. The accumulated pyrite forming

from the sulfide will be fractionated from seawater sulfate by an amount that depends on the availability of iron (a-c). The accumulated organic S will be fractionated from seawater sulfate by an amount that depends on the depth range over which it forms, and on the fractionation associated with reduced S incorporation into organic matter. Fig. 12. Correlation between total sulfur (TS) in the Aderet core and a) total iron (TFe), b) highly reducible iron (FeHR). Fig. 13. Depth profile of a) organic iron (Fe-Org) and pyritic iron, b) pyritic iron and TOC and c) δ34S values of pyrite and TOC of the Aderet core samples. Fig.14. A simplified cartoon demonstrating the effect of a “normal” (a) or “large” (b) organic matter flux, on the availability of Fe(II) for the reaction with sulfide species. At the early stages of diagenesis (upper panel), there are abundant functionalized organic compounds that can complex with Fe(II) and react with reduced S species. At the late stages (bottom panel), functionalized organic compounds are less abundant and Fe(II) is released from the organic complexes and may react with remaining reduced S species to form “late-stage” (34S enriched) pyrite.

Supplementary figure and table caption:

55

Table S1. List of iron fractions and their operational names extracted by treatment with various reagents in sequential extraction procedure according to Poulton and Canfield (2005). Table S2. Concentration (wt% of total S) of water soluble sulfate (SO42-(water)), acid soluble sulfate (SO42-(acid)), acid volatile sulfide (AVS), pyrite, kerogen and bitumen sulfur. Fig. S1. Typical secondary electron microscopy (SEM) photomicrographs of pyrite within the Aderet core samples.

56

Fig.1

(a)

(b) N

Aderet borehole Shefela Basin

Jerusalem

Be′er sheva

20 Km

600

Legend

76.2

79.2

83.6

75.7

550

Fossils Chalk

Chert

Mishash Tongue

En Zeitim

R. A. fructicosa mayaroensis

300

P. palpebra

67.3 69.2

C. G. Plumm. R. calc. havanensis

250

En Zeitim

71.7 72.1

Mount Scopus

400

D. asym G. elevata

500

Maastrichtian

350

Campanian

70.1

. Sant.

450

Zones

Age (Ma.) Stage Group

Depth (m)

Fig.2

Lithology & SR (cm/Ka) 0 2 468

P- Phosphate Marl Slickensides B- Bituminous

Core sample 9:1 DCM:MeOH

Fig.3 Residual rock

Bitumen

TDW

1:50 n-pentane

BaCl2 (5wt%)

Water soluble sulfate

Asphaltenes

Residual rock

Maltenes

HCl (6 N) heat

AgNO3 (5 wt%)

Acid volatile sulfide

Residual rockacid mixture Filtration BaCl2 (5wt%)

Residual rock Cr(II) treatment

Acid soluble sulfate

AgNO3 (5 wt%)

Pyritic S

Residual rockCr(II) mixture Filtration

Residual rock Oxygen flask

Organic S

SR (cm/Ka)

0 2 468

67.3 69.2

25

Kerogen S (wt% of TS) 0 50 100

(a)

Pyrite Kerogen

En Zeitim

0

S species (wt% of TS) 10 20 (c)

(b)

350 400 450 500

En Zeitim

P. palpebra

TOC (wt%) 12.5

Mishash Tongue

Mount Scopus

250

0

300

R. fructicosa

83.6

. Sant.

79.2

Campanian

75.7 76.2

C. G. Plumm. R. calc. havanensis

71.7 72.1

D. asym G. elevata

Maastrichtian

70.1

Depth (m)

Age (Ma.) Stage Group A. mayaroensis Zones

Fig.4

550 600

TOC TS

0

1.5 TS (wt%)

3

0

15 30 Pyritic S (wt% of TS)

SO42-(water) Bitumen

SO42-(acid) AVS

Fig.5

250

δ34S (‰) -10

-40

∆34S(pyrite-kerogen)(‰) 20 10 30 50

(a)

(b)

300

Depth (m)

350 400 450 500 550 600

Kerogen Pyrite AVS SO42-(acid) Original sulfate

∆34S(pyrite-kerogen) SO42-(water)

Fig.6 250

δ13C (‰) -30

-32

-28 -10

(a)

δ18O(‰) 5

20

(b)

300

Depth (m)

350 400 450 500 550 600

Kerogen

Bitumen

SO42-(water)

SO42-(acid)

Fig.7 250

0

TFe (wt%) 2

4 0

(a)

FeHR (wt%) 0.5

Fe species (wt%) 1 0 0.25 0.5 0 (b)

(c)

300

Depth (m)

350 400 450 500 550 600

Poorly crystalline Fe Easily reducible Fe oxyhydroxides Fe Carbonates Reducible Fe oxyhydroxides Pyritic Fe

FePy/FeHR 0.4

Fe-Org (wt% of TFe) 0.8 0 0.1 0.2 (d)

(e)

Fig.8

20

SO42-(acid) SO42-(water)

δ18O (‰)

15

10

5

0

-5

-25

-15

-5 δ34S (‰)

5

15

Fig.9

0

Measured δ34S (‰)

-5

-10

-15

y= 0.8 x - 1.1 R2= 0.9

-20

-25

-25

-20

-15 -10 Calculated δ34S (‰)

-5

0

Fig.10

250

f-values 0.5

0

1

0

Kerogen S (% of TS) 50 100 (b)

(a)

300

Depth (m)

350 400 450 500 550 600 Kerogen

Pyrite

0

20 40 Pyritic S (% of TS) Kerogen after oxidation Kerogen before oxidation Pyrite after oxidation Pyrite before oxidation

Fig.11

Fig.12

3

(a)

(b)

TS (wt%)

2

1 TS/TFe = 1.1

0

0

1

2 TFe (wt%)

TS/FeHR= 3.6

3 0

0.5 FeHR (wt%)

1

Fig.13 Pyritic Fe (wt% of TFe) δ34S (‰) TOC (wt%) -20 12.5 25 -40 0 15 30 0 250 (c) (a) (b)

0

300

Depth (m)

350 400 450 500 550 600

Fe-Org Pyritic Fe

Pyritic Fe TOC

Pyrite TOC

0.25 0.5 0 0 0.1 0.2 0 12.5 25 Fe-Org (wt% of TFe) Pyritic Fe (wt%) TOC (wt%)

Early stages S*

S*

S*

S*

S* S* S*

S*

FeS2

S

S

FeS2

S*

S*

Fe(II) Reactive iron S* Recent S boned

S

S* S

S

S

S*

Fe(II)

Reduced S species

S

S S*

OM

FeS2 S

S*

FeS2 S*

S*

S

S*

S*

S*

FeS2

S*

Fe(II)

FeS2 FeS2

S

S* S*

S*

FeS2 S

S*

S*

Fe(II)

S

S

S* Fe(II)

S*

S

S

S*

Fe(II)

S*

S

S*

S*

S*

Fe(II)

S*

S*

S*

S*

S*

Fe(II)

Fe(II)

S*

Fe(II)

S*

S* S*

(b) High OM flux

S*

Fe(II)

Fe(II)

Late stages

Fig.14

(a) Normal OM flux

Pyrite Late S boned

Functional group ( carbonyl, carboxyl, etc.)

FeS2

Late Pyrite

S

Fe(II)

S

FeS2

Table.1 Depth (m)

Bitumen (wt%)

AS/MA

S2(mg HC/g rock) 58.0

Production index (S1/(S1+S2)) 0.02

S3(mg CO2/g rock) 2.7

Tmax (⁰C)

1.5

S1(mg HC/g rock) 1.4

414.5

Hydrogen Index (mg HC/ g TOC) 803.9

Oxygen Index (mg CO2/g TOC) 38.1

265

0.2

308

0.1

1

1.9

65.1

0.03

2.6

413

762.5

30.2

341

0.3

1.2

1.1

69.5

0.02

2.6

415

949.2

34.6

393

0.7

1.1

2.5

109.4

0.02

3.1

408.7

708.6

20.4

420

0.6

1.1

3

116.7

0.03

3.2

407.3

721.3

19.8

450

0.4

1.1

1.7

87.1

0.02

2.9

411.7

747.4

26.8

498

0.3

0.9

0.9

48.6

0.02

1.9

402

677.5

26.3

528

0.8

1.7

2.3

93.5

0.02

1.9

408.7

853.8

17.3

595

0.1

1.5

0.1

13.8

0.01

1.6

422.3

623.3

73.8

57

Table.2

Depth (m)

Kerogen

Bitumen

Maltenes

Asphaltenes

265

3.4

3.9

4.65

4.05

308

-1.7

-0.15

-0.36

0.03

341

-6.9

-4.8

-4.45

-5.55

393

-0.3

-0.3

0.55

-0.35

420

1.4

1.38

2.15

1.5

450

-8.7

-7.23

-6.55

-6.5

498

-5.73

-5.15

-3.15

-5.3

528

-3.05

-2.35

-0.58

-2.9

595

-0.4

1.7

3.15

1.2

58