Continental Shelf Research 101 (2015) 117–124

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Geochemical evidence for biogenic methane production and consumption in the shallow sediments of the SE Mediterranean shelf (Israel) Michal Sela-Adler a, Barak Herut b,c, Itay Bar-Or a, Gilad Antler d, Efrat Eliani-Russak a, Elan Levy a, Yizhaq Makovsky c, Orit Sivan a,n a

Department of Geological and Environmental Sciences, Ben Gurion University of the Negev, Beer-Sheva 84105, Israel Israel Oceanographic and Limnological Research, National Institute of Oceanography, Haifa 31080, Israel Department of Marine Geosciences, Leon H. Charney School of Marine Sciences, University of Haifa, Haifa, Israel d Department of Earth Sciences, University of Cambridge, Cambridge CB2 3EQ, UK b c

art ic l e i nf o

a b s t r a c t

Article history: Received 25 November 2014 Received in revised form 9 March 2015 Accepted 3 April 2015 Available online 9 April 2015

This study presents geochemical evidence for biogenic methane formation (methanogenesis) in the shallow sediments of the oligotrophic SE Mediterranean continental shelf at water depths between 46 and 88 m. Depth-profiles of methane concentrations and related chemical parameters such as dissolved sulfate, dissolved inorganic carbon (DIC), and the stable carbon isotope composition of DIC and methane (δ13CDIC, δ13CCH4, respectively) were measured in six sediment cores (each 4.2–5.4 m long) in order to characterize the processes that involve methane production and decomposition. All the sediment cores reached the consumption depth of the entire sulfate pool and the in-situ microbial methane production (methanogenesis) zone. Methane concentrations reached saturation levels in one of the cores, but not in the others, probably because the zone of maximum methanogenesis was at a greater depth. Although the sediments exhibit a low TOC content of
1%, the biogenic methane formation indicates a relatively high organic carbon lability capable of sustaining all redox microbial activity potential. Anaerobic oxidation of methane (AOM) was also evident in the sulfate–methane transition zone, showing a distinct isotope signature in diffusion limited conditions. & 2015 Elsevier Ltd. All rights reserved.

Keywords: Methane Methanogenesis AOM SMTZ Stable isotopes Eastern Mediterranean

1. Introduction Methane (CH4) is an important natural gas used as an energy source, but also an effective greenhouse gas whose atmospheric concentration has more than doubled since preindustrial time, to values about 1800 ppbv (Hartmann et al., 2013). It is emitted from both anthropogenic and natural sources, and the emissions are estimated to be about 304–368 and 238–484 Tg CH4 yr  1 (Tg–1  1012 g), respectively as metioned in the IPCC 2013 report (Stocker et al., 2013). The major natural fluxes originate from the terrestrial environment; however, marine sediments are highly important as well, as they contain the largest natural reservoir of methane despite their low emission fluxes (
3%) (Archer, 2007). Natural sources of methane originate from microbial processes that take place in organic rich sediments (methanogenesis, e.g. (Whiticar, 1999)), or from kerogens following thermochemical reactions at high temperatures and pressures (thermogenic n

Corresponding author. E-mail address: [email protected] (O. Sivan).

http://dx.doi.org/10.1016/j.csr.2015.04.001 0278-4343/& 2015 Elsevier Ltd. All rights reserved.

methane, e.g. (Schoell, 1988)). Methanogenesis is the final process of microbial organic matter mineralization in anaerobic environments, after all other electron acceptors (O2, NO3, Mn(IV), Fe(II) and SO4) have been exhausted (Froelich et al., 1979). In freshwater sediments the dominant pathway for methanogenesis is acetate fermentation (acetoclastic methanogenesis, Eq. (1), whereas in marine sediments it is CO2 reduction by hydrogen (hydrogenotrophic methanogenesis, Eq. (2) (Whiticar et al., 1986)

CH3 COOH → CO2 + CH4

(1)

4H2 + CO2 → 2H2 O + CH4

(2)

When the produced methane diffuses into contact with an available electron acceptor it can be consumed by microbial oxidation (methanotrophy). Methanotrophy is the main process that prevents the escape of methane into the atmosphere. In oxic water methanotrophic bacteria are responsible for oxidizing methane into CO2, using O2 as an electron acceptor (Chistoserdova et al., 2005). In marine sediments anaerobic oxidation of methane (AOM), coupled to sulfate reduction as shown in Eq. 3 (Hoehler

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et al., 1994), was found to consume up to 90% of the upward-diffused methane at the sulfate-methane transition zone (SMTZ), thus preventing its release from this huge reservoir into the atmosphere (Valentine, 2002). Frequently, most of the sulfate is reduced by this process (e.g. Niewöhner et al., 1998; Boetius et al., 2000; Aharon and Fu, 2000; Sivan et al., 2007).

CH4 + SO42 − → HS− + HCO3− + H2 O

(3)

The AOM process typically involves a microbial consortium of archaea and bacteria that are affiliated with Methanosarcina-type methanogens and sulfate-reducing bacteria, respectively (Boetius et al., 2000; Orphan et al., 2002; Thauer, 2011). It seems that anaerobic methanotrophic archaea (ANME) oxidize the methane, while the bacterial partner uses the resulting reducing equivalents to reduce sulfate (Thauer and Shima, 2006; Basen et al., 2011). Recently, however, AOM mediated solely by archaea was reported (Milucka et al., 2012), in which the archaea was shown to oxidize the methane and reduce the sulfate to elemental sulfur. Methanogenesis and methanotrophy processes have also been studied using a geochemical isotopic approach. Carbon isotopes provide a good constraint on the depth distribution and location of methanogenesis and methanotrophy and can be used to quantify these processes due to the large carbon isotopic fractionation associated with both methane production and consumption (e.g. Whiticar, 1999; Borowski et al., 2000; Sivan et al., 2007; Adler et al., 2011). During methanogenesis, 12C is strongly partitioned into methane; the δ13C of the methane produced can vary between –50‰ and –100‰. In contrast, the residual dissolved inorganic carbon pool becomes highly enriched in 13C, occasionally by as much as 50–70‰. Oxidizing this methane during AOM, on the other hand, results in 13C-depleted dissolved inorganic carbon (DIC) and in slightly heavier δ13C values of the residual methane, due to both a fractionation of 0–10‰ during the oxidation of methane and to the initial δ13C value of the methane itself (Alperin et al., 1988; Martens et al., 1999). Typically, thermogenic methane is enriched in 13C compared to microbial methane, having δ13CCH4 values that range between approximately –50‰ and –20‰ (Whiticar, 1999). Methane-related processes in the sediments of the Eastern Mediterranean Sea are of great importance, from both the economic and the environmental aspects. Recently, large gas fields from the deep geological layers of the Oligocene–Miocene era were discovered in sediments in the deep water ( 41500 m) of the SE Mediterranean Sea (Levantine basin) (Gardosh and Tannenbaum 2014). Methane seeps were also found in the deep sea sediments (Omoregie et al., 2009,, 2008; Rubin‐Blum et al. (2014). However, methane production in the shallow sediments of the SE Mediterranean continental shelf has not been assumed due to its current oligotrophic conditions (Low nutrient, low chlorophyll; Herut et al., 2000; Kress et al., 2014), which offer a relatively low content of organic carbon in the sediment (o 1%). Schattner et al. (2012) interpreted a band of high amplitude scattered reflectivity

observed in high (
0.3 m) resolution seismic profiles across the continental shelf of northern Israel to reflect the presence of a ‘gas front’ within the seafloor sediments at water depths between 37 and 112 m. They also repeatedly (for over 3 years) observed acoustic reflectivity in the water above the seafloor, which they suggested represents a long term active gas seepage. This study aimed to test whether there is methane formation in the shallow sediments of the oligotrophic SE Mediterranean continental shelf, and to characterize the major sinks and sources of methane in the shelf, including their depth distribution. This was achieved by using a geochemical approach similar to the one used in our work on sediments from estuaries (Antler et al., 2014) and lakes (Adler et al., 2011; Sivan et al., 2011). For the first time in this area, we were able to collect, several sediment cores 5–6 m long from the continental shelf of Israel and to perform depth profiles of methane and its related parameters in the pore water (dissolved sulfate, dissolved inorganic carbon (DIC) and the stable carbon isotope composition of DIC and methane (δ13CDIC, δ13CCH4, respectively)).

2. Materials and methods 2.1. Sampling Six sediment cores were collected from the undisturbed seafloor sediments of the Mediterranean continental shelf of Israel at water depths between 49 and 88 m (Table 1 and Fig. 1), during two cruises on the R.V. Shikmona that took place on August 2013 and February 2014. The sediment cores (each 4.6–5.4 m long) were collected using a Benthos 2175 piston corer. The study site locations were chosen based on the mapping of a sub-bottom depth high amplitude scattered reflectivity layer, interpreted to be related with the presence of free gas bubbles (e.g. (Schattner et al., 2012) ‘Gas Front’), as observed in high (
0.3 m) resolution seismic profiles collected between 2008 and 2013 by Moses Strauss, as well as additional unpublished data). The actual coring locations generally correspond to relatively shallow interpreted gas related reflectivity. The sediment cores were sliced onboard at intervals of 40–50 cm within minutes of extracting the core from the seafloor. About 1.5 ml from the edge of each sediment slice was immediately transferred into N2-flushed crimp bottles containing 5 ml of 1.5 N NaOH for the headspace measurements of CH4 and δ13CCH4 (after Adler et al., 2011). Subsamples of 100 ml from the edge of each sediment slice were stored in vials under anaerobic conditions and pore water was extracted from them on the same day by centrifugation at 4 °C under a N2 atmosphere. The supernatant was filtered through 0.45 μm filters. The sediment was measured for its total organic carbon (TOC) content during the February cruise.

Table 1 The cores sampling dates, positions, water depths and length retrieved. Sampling date

Core name

Water depth (m)

Retrieval (m)

Location

Latitude (N)

Longitude (E)

August 14, 2013

PC–3 PC–5 PC–6

81 87 49

5.1 4.6 4.5

Offshore Haifa Offshore Haifa Offshore Netanya

32°55.29' 32°55.47' 32°17.11'

34°54.14' 34°54.01' 34°44.62'

February 6, 2014

PC–3–14n NRD SG1

82 81 88

4.8 5.35 4.2

Offshore Haifa Offshore Acre Offshore Acre

32°55.30' 32°59.63' 32°57.83'

34°54.14' 34°56.27' 34°55.29'

n

Core PC-3-14 was sampled on February, 2014, in proximity to the sampling position of core PC-3 on August, 2013.

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Fig. 1. Location map of the sites in which the sediment cores were sampled. P-130 and HU are shorter cores drilled previously and their results are presented in Antler et al. (2013).

2.2. Analytical methods Headspace methane concentrations were measured on a gas chromatograph equipped with a FID detector at a precision of 2 μM. Sulfate was measured using a Dionex DX500 high pressure liquid chromatography (HPLC) with an error of 2%. About 0.5 ml of each sample was transferred into a He-flushed vial containing H3PO4 for the headspace measurements of δ13CDIC using a conventional isotopic ratio mass spectrometer (IRMS, DeltaV Advantage, Thermo) with a precision of 70.1‰, and the results are reported versus the VPDB standard. DIC concentration was calculated from the IRMS results according to peak height and to a calibration curve (by standard samples prepared from NaHCO3 with a known DIC concentration) with an error of 70.2 mM. The δ13CCH4 values were measured using an IRMS equipped with a PreCon interface after oxidation to CO2. The error between duplicates of this parameter was less than 0.5‰ and the results are reported versus the VPDB standard. Total organic carbon (TOC) was determined using the potassium dichromate method according to the procedure detailed by (Gaudette et al., 1974) and (Avnimelech, 1989), using NIST 1941b (organics in marine sediment) as reference standard instead of glucose, and is presented in % of dry sediment. The saturation values of methane were calculated based on Duan and Mao (2006) and Mogollón et al. (2009), using pressure, salinity and temperature.

3. Results The continental margin (shelf) of Israel narrows, steepens and deepens from south to north, and is mainly composed of Pliocene– Quaternary Nile-derived sediments (Nir, 1984). The sediments

collected in this study were mostly clayey silts at PC-6 (sediments at water depths of 30–50 m), while the clay fraction dominated at 450 m water depths (Nir, 1984; Sandler and Herut, 2000), at which all other sediment cores were collected (Table 1, Fig. 1). Along the shelf (in
40–50 m water depths) the carbonate fraction increases northward (7–15 wt%) and toward the shore due to the decrease in Nile derived sediments and to higher amounts of local calcium carbonate biogenic fragments (Nir, 1984; Goldsmith et al. 2001). The porosity was approximately 0.5 along the cores. The bottom water temperatures were
17 and 22 °C in February and August, respectively. The chemical and isotopic profiles of the porewater are shown in Fig. 2. Sulfate concentration values within the upper sediments of most cores were about 32 mM, typical to concentrations of the Eastern Mediterranean sea water. The exceptions were NRD, where the upper most part was not sampled (Fig. 2), and SG-1, which showed a lower sulfate concentration of 20 mM, probably due to a disturbance at the top 5 cm. In all collected cores the sulfate concentrations decreased with depth until reaching the SMTZ, the zone where sulfate is depleted and methane concentrations start to increase. The profiles of the sulfate concentrations in PC–3, PC–3–14, PC–5 and PC–6 showed a clear pattern of linear decrease that continued until complete depletion (Fig. 2). The SMTZ in PC–3–14 was located at a 200–350 cm depth below surface, at a 250–370 cm depth in PC–3 and PC– 5, and at a 350–400 cm depth in PC–6. Within the NRD core the SMTZ was located 170–300 cm below the surface. Within core SG–1 sulfate was already depleted at a depth of above 80 cm (it is difficult to determine the exact depth at which depletion occurs due to the low sampling resolution and the probable removal of the top of the core, but it is certainly within the upper 80 cm). The saturation values of the methane were 5 and 10 mM for the

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Fig. 2. Chemical and isotopic depth profiles in pore water from sediment cores collected on August 2013 off the coasts of Haifa (PC-3, PC-5) and Netanya (PC-6) and on February 2014 off the coast of Haifa (PC-3-14, NRD and SG-1). Gray box marks the SMTZ depth in each core based on the measured resolution (from the estimated depth where sulfate concentration is lower than 1 mM to the depth where methane starts to increase). The uppermost part of cores NRD and SG-1 was probably disturbed.

shallow (46 m water depth) and deep (88 m water depth) sediment cores, respectively. Methane concentrations in all cores beside SG–1 did not exceed 2 mM at the bottom of the cores and remained below saturation level at the sampling depths. In core SG–1 bubbles were observed along nearly the entire core, and indeed methane concentrations reached 35 mM. The methane profile was very scattered in this core, possibly due to the bubble interference. The δ13CH4 values ranged between –80 and –100‰

in all the profiles. In some of the profiles (PC–3, PC–3–14 and PC– 5) the first upper samples in which we succeeded in measuring δ13CCH4 (above concentrations of 140 nM) were relatively heavy, and below them there was a sharp decrease of δ13CCH4 followed by a gradual increase with depth (Fig. 2). For example, in PC–3 the values decreased from approximately –82‰ to –85‰ and then increased gradually to –81.5‰. The DIC concentrations (Fig. 2) in the upper sediments of the

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Fig. 2. (continued)

cores were similar or higher than bottom seawater concentrations (
2.3 mM; Álvarez et al., 2013), matching the behavior of the sulfate concentrations. Maximum DIC concentrations were obtained in the SMTZ zone (20–40 mM) in most of the cores (PC–3, PC–3–14, NRD and SG–1). In NRD and SG–1 there was a significant decrease to about 20 mM and 10 mM at the deep sediment, respectively. The δ13CDIC profiles show a gradual downward concave decrease to minimum values of –25 to –40‰ at the SMTZ. Below that depth the values increase gradually up to –5‰ at PC–3–14.

The TOC values (not shown) were measured on February 2014 and are more or less constant with depth, with values of approximately 1% at PC3–14,
0.7% at the NRD core and
0.8% at SG–1.

4. Discussion Previous sediments (cores HU and P–130, Fig. 1) collected in proximity to each other at water depths of 68 and 67 m on the

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continental shelf offshore Haifa, northern Israel, showed similar sulfate concentrations in the seafloor. However, the sulfate concentrations in the
2.5 m deep of these cores decreased only by 50%, and subsequently no methane was detected (Antler et al. 2013). Cores HU and P-130 were extracted in the vicinity where Schattner et al. (2012) showed acoustic reflectivity in the water above the seafloor. It was suggested that this reflectivity represented a long-term active gas seepage, but no sub-surface interpreted gas related reflectivity (Schattner et al. 2012, ‘Gas Front’) was observed. In contrast, the cores of this study were collected where such interpreted gas related reflectivity was observed to be relatively shallow ( o 10 ms ( o8 m) below the seafloor). Indeed, the main observation of this study is that sulfate depletion occurs at a relatively shallow depth and that the transition to methane occurs below that depth. Thus, in all collected sediment cores we reached the SMTZ within the upper 4 m. The SMTZ was particularly shallow in core SG1, within the upper 80 cm. It seems that in the shelf there are cores without methane (Antler et al., 2013), cores with shallow SMTZ and a core with very shallow SMTZ, all of which are in close proximity to one another. This variability probably coincides with the methanohenesis rates and the observed depth variability of the gas reflectivity layer. Indeed, the very shallow SMTZ was observed in SG–1, where the reflective gas layer was extremely shallow, in the tens of centimeters. Higher methanogenesis rates would lead to a greater diffusive flux of methane, and thus to a higher downward diffusive flux of sulfate to the AOM zone, to intensive removal and to a shallower SMTZ (as shown in Sivan et al. (2007)). It is not clear, however, why the shelf contains such a variety of methanogenesis rates with similar TOC levels of 1%. The fact that there is a transition to methanogenesis in accordance with the respiration order from sulfate reduction to methanogenesis, the clear SMTZ in all profiles, and the isotopic values, all clearly suggest that the methane is indeed formed by the biogenic processes of methanogenesis in-situ at these depths and probably below them. The biogenic formation is evident by the carbon isotopic values of methane, which range between –80 and –100‰ in all the profiles, fitting methanogenesis from CO2 reduction pathway (Whiticar, 1999), and by the large increase in δ13CDIC in the deep sediments due to methanogenesis with a large carbon isotopic fractionation. Some advection of methane gas from below cannot be ruled out, and probably occurs in core SG–1; however, the diffusive profiles and the isotopic values fit to in-situ production of methane just below the SMTZ in most of the profiles. The shallow SMTZ and the presence of biogenic methane formation in the shallow sediments indicate that although the eastern Mediterranean shelf is oligotrophic, the TOC substrate is highly labile and is sufficient to sustain all the microbial activity up to methanogenesis. This observation emphasizes the possibility of finding methane in shallow sediment depths at these low TOC levels along the shelf in cases where the organic carbon is highly reactive. Moreover, methane was found above the saturation levels at SG–1, creating bubbles, and it seems that in all the cores collected in this study gas would have appeared if we had been able to collect longer cores and reach the maximum methanogenesis zone. Therefore, the results suggest the presence of methane gas at several locations in the sediments of the shelf. The geochemical profiles also suggest new constraints regarding the AOM process in this system. In the sulfate reduction zone, most of the profiles (PC–3, PC  3–14, PC–5 and PC–6) show a clear linear diffusive curve toward the SMTZ. This type of decrease has been observed in many profiles around the world and indicates that often, most of the sulfate is reduced by methane in the AOM process at the SMTZ rather than by other organic carbon compounds above it (e.g. Niewöhner et al., 1998; Boetius et al., 2000;

Fig. 3. : Schematic profiles of SO42– and CH4 concentrations (a) and δ13C schematic profiles of CH4 and DIC, based on Yoshinaga et al. (2014).

Aharon and Fu, 2000; Sivan et al., 2007; Antler et al., 2014). The AOM process was shown to occur in seeps in the deep eastern Mediterranean (e.g. Rubin‐Blum et al. (2014), but here it was found for the first time at the shelf, in shallow sediments within the SMTZ. The AOM process in these sediments at the SMTZ is revealed by the very light isotopic values of the DIC, which reach as low as –40‰. It is also apparent from the increase in the δ13CCH4 at the upper AOM zone due to enrichment of the residual methane by 0–10‰ (Alperin et al., 1988; Martens et al., 1999). The AOM process releases bicarbonate into the pore water, thus increasing the alkalinity, and can cause carbonate mineral precipitation (e.g. Sivan et al., 2007; Zeebe, 2007; Suess et al., 1999). This can explain the slight decrease in the net DIC observed at the SMTZ in some of the cores. The depletion in δ13CCH4 at the main AOM zone and the gradual increase below it is not obvious, but fits recent observations that showed this phenomenon in the SMTZ zone (Fig. 3, after Yoshinaga et al., 2014). According to Yoshinaga et al. (2014), the upper AOM zone shows the expected enrichment of the residual methane in 13C due to partial oxidation. The bottom of the AOM zone is sulfate-limited and thus affected by the backward reaction of Eq. 3. This back flux leads to the sharp decrease in δ13CCH4 values. Below that zone, in the methanogenesis zone, there is gradual increase in the δ13CCH4 values due to the enrichment of the residual DIC during the methanogenesis process (Fig. 3). The “back reaction” explanation is based on predictions from reactions approaching thermodynamic equilibrium (ΔG-0), and was demonstrated to be 5% back of DIC to methane in AOM (Holler et al., 2011). Yoshinaga et al. (2014) showed, by incubation experiments of methanotrophs grown under different sulfate concentrations, that the residual methane was indeed depleted in 13C when concentrations reached below 0.5 mM of sulfate, whereas in high concentrations of sulfate the AOM showed resulted in the expected 13C methane enrichment. Yoshinaga et al. (2014) also showed that the increase from the δ13CCH4 values found below the depletion depth to the maximum values found at the maximum methanogenesis zone (so called

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Δδ13CCH4) is inversely correlated to the sulfate fluxes toward the SMTZ zone. We calculated the sulfate fluxes in our sediments (PC–3, PC–5, PC–6 and PC–3–14) similarly to Yoshinaga et al. (2014), using the sulfate concentrations gradients with depth. The molecular diffusion coefficient for sulfate in seawater (D0) was taken as 6  10  6 cm2 s  1, the measured porosity was around 0.5, and the assumption was that Ds ¼D0ϕ2 (Berner, 1980) was used for estimating the tortuosity effect. The fluxes ranged between values of 5 mmol cm2 s  1 (PC–5) and 10 mmol cm2 s  1 (PC–3–14). We were not able to calculate Δδ13CCH4, since we did not reach the maximum methanogenesis depth, and therefore the δ13CCH4 values did not level off. However, from the pattern of δ13CCH4 in the methanogenesis zone we can roughly extrapolate our data and estimate that the Δδ13CCH4 in the sediments is approximately 15‰ (an observed increase of
5‰ in δ13CCH4 and a 10‰ extrapolated increase up to the leveling off). These Δδ13CCH4 values fit our calculated sulfate fluxes according to Yoshinaga et al. (2014) and their compiled data set. Thus, it seems that the AOM in the Eastern Mediterranean continental shelf is typical of low sulfate fluxes, similar, for example, to the Gulf of Mexico (Yoshinaga et al., 2014). It should be noted that strangely, our calculated methane upward fluxes toward the AOM zone were not similar to the downward sulfate fluxes, as may be expected from the linear decrease in sulfate. Instead, they were about 1/5 of the sulfate fluxes. This infers that the sulfate was consumed together with the other organic carbon at the AOM zone, through a yet unknown process, in which less reactive organic compounds that were not used according to the diffusive profile of the sulfate are consumed. Another possibility is that methane escaped during the course of the coring and sampling procedure due to pressure release, resulting in a lower estimation of the methane flux. However, the consistent depth pattern of the methane concentration profiles weakens the latter hypothesis. To summarize, our results clearly indicate the microbial production of methane in some shallow low content organic carbon sediments at the eastern Mediterranean shelf. It seems that this organic carbon is highly reactive to methanogens and supports methanogenesis, enabling the consumption of the entire sulfate pool, reaching the SMTZ by upward methane diffusion. The methane reached saturation levels in one of the cores (SG–1), where bubbles appeared. It seems that in the rest of the cores the methane reached its maximum values below the drilling depths. Acknowledgment We would like to thank associate editor Prof. Gary Fones and the anonymous reviewers for their thorough work on the manuscript. The authors thank the captain and crew of the R/V Shikmona and Yaron Gertner’s research assistant, all from the Israel Oceanographic and Limnological Research Institute, for their assistance during field sampling. We thank Ganor’s lab for assisting with the sulphate analyses, and Paradigm for their software support. This work was partly supported by a Grant #212-17-024 named "Methane production and oxidation in the Eastern Mediterranean shelf" from the Israel Ministry of Infrastructure, Energy and Water (to OS and BH Y.M), by the PERSEUS project, by EC Contract #287600 (BH) and by the Ministry of Science and Technology, project # 3-9145 (to O.S. and Y.M).

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