The multiple sulphur isotope fingerprint of a sub-seafloor oxidative sulphur cycle driven by iron

Earth and Planetary Science Letters 536 (2020) 116165

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Earth and Planetary Science Letters www.elsevier.com/locate/epsl

The multiple sulphur isotope fingerprint of a sub-seafloor oxidative sulphur cycle driven by iron Jiarui Liu a,b,1 , André Pellerin c,d,2 , Gareth Izon b,2 , Jiasheng Wang a,∗ , Gilad Antler d,e , Jinqiang Liang f , Pibo Su f , Bo Barker Jørgensen c , Shuhei Ono b a

State Key Laboratory of Biogeology and Environment Geology, College of Marine Science and Technology, School of Earth Sciences, China University of Geosciences, Wuhan 430074, China b Department of Earth, Atmospheric and Planetary Sciences, Massachusetts Institute of Technology, Cambridge, MA 02139, USA c Center for Geomicrobiology, Department of Bioscience, Aarhus University, 8000 Aarhus C, Denmark d Department of Geological and Environmental Sciences, Ben-Gurion University of the Negev, Beersheba 84105, Israel e The Interuniversity Institute for Marine Sciences, Eilat 88103, Israel f Guangzhou Marine Geological Survey, Guangzhou 510760, China

a r t i c l e

i n f o

Article history: Received 17 August 2019 Received in revised form 22 November 2019 Accepted 13 February 2020 Available online xxxx Editor: I. Halevy Keywords: cryptic sulphur cycle deep biosphere sedimentary pyrite South China Sea sulphur disproportionation triple sulphur isotopes

a b s t r a c t Oxidative sulphur cycling is pervasive in marine sediments, replenishing the oxidised sulphur reservoir via re-oxidation of sulphide. An active, yet cryptic, sulphur cycle has been proposed to operate at depth beneath the sulphate-methane transition (SMT), fuelled by simultaneous sulphide oxidation and sulphate reduction under low-sulphate conditions. The existence of a cryptic sulphur cycle, however, is centred on porewater and genetic data that have little, to no, preservation potential, and thus are rarely accessible from the geological record. The absence of a suitable archive has hindered our ability to reconstruct the operation and importance of the cryptic sulphur cycle through space and time. To overcome this obstacle, and to develop a better understanding of the oxidative sulphur cycle in the deep biosphere, we have determined the abundance and triple sulphur isotope composition (33 S and δ 34 S) of both elemental sulphur and pyrite extracted from sediments recovered from the methane prone Taixinan Basin, South China Sea. Here, multiple sulphur isotope systematics of pyrite clearly reveal a tiering, with organoclastic sulphate reduction succumbing to sulphate-driven anaerobic oxidation of methane at depth. Importantly, a negative 33 S-δ 34 S correlation was found at the periphery of the SMT that requires repeated and sustained iron-driven sulphide oxidation with concomitant disproportionation of the elemental sulphur product. We conclude that minor sulphur isotopes may provide a unique lens to resolve the cryptic sulphur cycle, allowing the importance of the deep biosphere to be evaluated over geological timescales. In turn, a better understanding of the cryptic sulphur cycle remains central to testing hypotheses linking major elemental cycles and diverse microbial activities that persist under the energy-limited conditions that typify the deep biosphere. © 2020 Elsevier B.V. All rights reserved.

1. Introduction Modulated by a myriad of metabolic and geochemical processes, sulphur cycling plays a fundamental role in regulating Earth’s surface chemistry. For example, in contemporary anaerobic marine sediments, microbially-mediated sulphate reduction dominates organic matter mineralisation (Jørgensen, 1982). Con-

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Corresponding author. E-mail address: [email protected] (J. Wang). 1 Now at Department of Earth, Planetary and Space Sciences, University of California, Los Angeles, CA 90095, USA. 2 These authors contributed equally: André Pellerin and Gareth Izon. https://doi.org/10.1016/j.epsl.2020.116165 0012-821X/© 2020 Elsevier B.V. All rights reserved.

current with the reduction of sulphate to sulphide, the oxidation of sulphide to sulphur intermediates (e.g., elemental sulphur) and sulphate represents a critical component of the marine sulphur cycle (e.g., Jørgensen and Kasten, 2006). In continental margin sediments, organoclastic sulphate reduction (OSR) utilises porewater sulphate via the degradation of organic matter. Additionally, residual porewater sulphate can be consumed at depth at the sulphate– methane transition (SMT). Here, an upward flux of methane reacts with a descending flux of sulphate, fuelling sulphate-driven anaerobic oxidation of methane (AOM; e.g., Boetius et al., 2000). Following the classical redox-cascade of electron accepting processes (e.g., Berner, 1981), the SMT represents the depth where sulphate is exhausted, and methanogenesis attains dominance over

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organic matter mineralisation (Jørgensen and Kasten, 2006). More recently, however, several studies have shown that sulphur cycling is perhaps not restricted to sediments above SMT (Holmkvist et al., 2011; Treude et al., 2014; Brunner et al., 2016; Pellerin et al., 2018), suggesting that the vertical tiering of metabolisms based on their Gibbs energy yields is incomplete, and requires re-evaluation. Combining porewater measurements and radiotracer experiments, Holmkvist et al. (2011) found low, but detectable, levels of porewater sulphate in methanogenic sediments from Aarhus Bay, providing a potential electron acceptor for sub-SMT sulphate reduction. Given the efficiency of the SMT as a sulphate sink, the presence of this deep sulphate was rationalised via cryptic sulphur cycling, whereby sulphate was generated in-situ via the oxidation of downward-diffusing sulphide through reaction with deeply buried iron (III) species (Holmkvist et al., 2011). Subsequently, more detailed biogeochemical studies have illuminated the operation of the cryptic sulphur cycle and its coupling with iron, manganese, and barite in sediments from the Baltic and Beaufort seas (Treude et al., 2014; Brunner et al., 2016; Pellerin et al., 2018). Moreover, genetic assays (e.g., real-time PCR of 16S rRNA and dsrAB genes) have demonstrated the presence of sulphate reducing microorganisms beneath the SMT, providing supporting evidence for at least the metabolic capability for sulphate reduction in these deeper, sulphate impoverished, sediments (e.g., Leloup et al., 2007; Treude et al., 2014). Interestingly, in culture, Sivan et al. (2014) demonstrated that iron-oxides are capable of stimulating sulphide re-oxidation and accelerating sulphate-driven AOM in methanerich settings. Consequently, exploration of the cryptic sulphur cycle could provide new insight into deep biospheric processes, improving our understanding of microbial activity under energy-limited conditions. Microbially mediated sulphur cycling is typically associated with pronounced isotopic fractionations (e.g., Canfield, 2001). Indeed, isotopic data (δ 34 S) derived from various phases has provided unprecedented insight into the operation of the sulphur cycle in ancient and modern environments alike (e.g., Fike et al., 2015). Excitingly, within the last two decades the isotopic toolkit has been expanded, and the generation of accurate and precise minor sulphur isotope data (33 S and 36 S) has provided an additional means to inform on a variety of microbially mediated processes. Similar to 34 S, sulphur’s minor isotopes (33 S, 36 S) are subject to similar fractionation mechanisms, yet their redistribution at both the cellular and ecosystem level results in unique patterns that can be used to ascertain the relative contribution of different enzymatic/biogeochemical pathways (e.g., Farquhar et al., 2003; Johnston et al., 2005; Ono et al., 2006). Previous experimental studies have shown that specific metabolisms produce characteristic minor isotope patterns. For example, the mass-dependent exponent (λ) for microbial sulphate reduction is smaller than the one expected for low-temperature inorganic chemical equilibrium (∼0.515; Farquhar et al., 2003; Johnston et al., 2005; Sim et al., 2011b; Leavitt et al., 2013), while the λ for microbial sulphur disproportionation is close to equilibrium predictions (Johnston et al., 2005). Given the relative infancy of multiple sulphur isotope studies in contemporary sediments, and compounded by the specialised equipment necessary to conduct SF6 measurements, our understanding of sulphur’s minor isotope systematics has been gleaned from relatively few studies exploiting a range of approaches and substrates. For example, much of the multiple sulphur isotope data used to comprehend biogeochemical processes originates from laboratory-based culture experiments, while most studies pursuing the cryptic sulphur cycle target marine-derived porewater sulphates (e.g., Pellerin et al., 2018). Unfortunately, due to the negligible preservation potential of porewater, coupled with the lability of genetic material, a thorough temporal understanding of the cryptic

sulphur cycle has remained elusive. Surprisingly, despite its ubiquity, there is limited work designed to explore and exploit the minor sulphur isotope systematics of pyrites from modern marine settings where multiple sulphur isotope signatures may reveal diverse biogeochemical processes (e.g., Lin et al., 2017; Gong et al., 2018). As such, multiple sulphur isotope systematics of sedimentary pyrite may provide a lens through which the cryptic sulphur cycle can be resolved. To this end, targeting methane-rich sediments from the Taixinan Basin, South China Sea, we report the abundance and triple-sulphur isotope systematics of elemental sulphur and pyrite from core 973-4. Combining these data with reactive iron distributions (Liu et al., 2018), we evaluate the transfer and preservation potential of an isotopic fingerprint associated with a deep iron-driven oxidative sulphur cycle. 2. Material and methods 2.1. Study site Here we focus on core 973-4, a 14-m-long piston core retrieved from 1666 metres water depth within Taixinan Basin (118◦ 49.0818’ E, 21◦ 54.3247’ N), retrieved as part of a cruise with R/V Ocean VI in 2011 targeting the north-eastern continental slope of the South China Sea (Fig. S1). The Taixinan Basin is methane prone, featuring vigorous cold seeps capable of supporting diverse chemosynthetic ecosystems, large seep-induced carbonate accumulations, and massive gas hydrates in the upper continental slope (Feng et al., 2018). Core 973-4 is dominated by dark-green silty clay, with a siltier horizon at 455–605 cm below seafloor (cmbsf). Except for this coarse horizon, featuring abundant foraminifera and silicate aggregates, the radiocarbon-derived age model suggests that sedimentation rates were broadly constant (32 cm ka−1 ) throughout the c.40-thousand-years (ka) of deposition (Fig. 1a). Deposited during the Last Glacial Maximum (LGM), the coarser layer yields atypically old radiocarbon ages relative to the juxtaposed sediment. This age inflection is consistent with the destabilisation of the continental-slope and delivery of aged sediments via turbidity currents during a glacially-induced sea-level low stand. The total organic carbon (TOC) content throughout core 973-4 is low and more-or-less invariant (Zhang et al., 2015), with an average value of 0.38 wt % (Fig. 1b). 2.2. Sulphur extraction After retrieval, core 973-4 was sectioned and refrigerated (< 4 ◦ C). Immediately after the cruise the sections were split, subsampled, and frozen (−20 ◦ C). Aliquots of these frozen samples were used to determine the abundances and isotopic composition of elemental sulphur and pyrite. Firstly, aliquots of frozen sediment were submerged in 20% zinc acetate dihydrate solution, converting polysulphidic zero-valent sulphur to elemental sulphur (S0 ), which was subsequently extracted via sustained agitation (12–16 h) with methanol (1 g sediment/20 mL MeOH; Zopfi et al., 2004). The S0 content of the supernatant was quantified via reversed-phase high-performance liquid chromatography (HPLC, Agilent 1100) using a UV/VIS detector set at 254 nm. Separation was achieved using a C-18 column with a 95:5 mixed methanol:water eluent. The pump speed was set to 1 mL min−1 , resulting in S0 retention times of ∼12 min. Data were quantified using a multi-point calibration curve constructed from solutions of orthorhombic S8 dissolved in dichloromethane, with S0 contents bracketing those of the unknowns. The detection limit for this method is ∼ 1μM, and the analytical precision was deemed to be better than 0.7% (relative standard deviation; RSD). After separation from the sediment residue, the dissolved S0 content of the supernatant was reduced to H2 S via distillation with

J. Liu et al. / Earth and Planetary Science Letters 536 (2020) 116165

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Fig. 1. Age-model and down core trends in grainsize, TOC and interstitial sulphate and methane concentrations. (a) Stratigraphic distribution of the coarse fraction (> 65 μm) and radiocarbon-derived (AMS 14 C) age model from core 973-4. The horizontal bar in a represents the coarse layer with anomalously old radiocarbon ages. (b) Profile of total organic carbon (TOC) content from core 973-4 (Zhang et al., 2015). The vertical dotted line indicates the average TOC content (0.38 wt %). (c) Profiles of interstitial sulphate (circles) and methane (diamonds) concentrations from core 973-4 and other neighbouring sites. Porewater from core 973-4 was not immediately extracted upon core recovery but was extracted several months later via centrifugation of previously frozen sub-samples. This time-lag potentially compromised the porewater sulphate profiles, and thus the sulphate data in and below the SMT from core 973-4 are not shown. The horizontal bar in c represents the estimated depth of the SMT. Panel a and c are adopted from Liu et al. (2018) and references therein. MIS—marine isotope stage. (For interpretation of the colours in the figure(s), the reader is referred to the web version of this article.)

acidified chromium chloride (CrCl2 ). The recovered solid residue was subjected to a two-step sulphide extraction (Fossing and Jørgensen, 1989), where acid volatile sulphur (AVS, mainly Fe monosulphide) was extracted with 6N HCl, followed by a CrCl2 solution to release the chromium reducible sulphur (CRS, mainly pyrite). Despite being identified previously (Zhang et al., 2014), no AVS was detected in any of our extractions. This discrepancy results from ongoing sample oxidation despite being stored at −20 ◦ C (Liu et al., 2018). The liberated H2 S from the S0 and CRS distillations was trapped as silver sulphide (Ag2 S), which was collected by filtration and dried at 50 ◦ C overnight. Abundances of S0 and CRS were determined gravimetrically. Pyrite abundances were calculated using Ag2 S yields, assuming an ideal pyrite Fe:S stoichiometry (1:2), and expressed as micromoles of FeS2 per gram of sediment. Replicate analyses of an in-house pyrite standard demonstrated recoveries exceeded 90%, with a RSD of better than 4%.

δ 3X S () = [(3X RSample /3X RVCDT ) − 1] × 1000 3X

(1)

3X

Where RSample / RVCDT is the isotopic ratio of a sample (3X RSample = 3X S/32 S and 3X = 33 or 34) relative to Vienna Canyon Diablo Troilite (VCDT). The minor isotopic composition of sulphur species is presented in 33 S notation, which describes the deviation of a sample from a reference fractionation line (Eq. (2)):

 33

33

 S () = δ S − 1000 ×

1+

δ 34 S 1000



0.515 −1

(2)

Analytical uncertainties are estimated from the long-term reproducibility of IAEA (S1, S2, S3) and in-house standards, and are deemed to be 0.2 and 0.006 (1 SD or 1σ , n = 39) for δ 34 S and 33 S, respectively. 2.4. Iron extraction and scanning electron microscopy

2.3. Multiple sulphur isotope analyses Multiple sulphur isotope (33 S/32 S, 34 S/32 S) ratios were measured in the Laboratory of Stable Isotope Geobiology at Massachusetts Institute of Technology following the procedure outlined in Ono et al. (2012). Briefly, ∼2 mg of Ag2 S was reacted with fluorine gas (F2 ) at 300 ◦ C for > 8 h. The reaction product, sulphur hexafluoride (SF6 ), was initially purified cryogenically, and then via gas chromatography (GC). Isotope ratios of the purified SF6 were determined using a Thermo-electron MAT 253 isotope ratio mass spectrometer in dual-inlet mode. Simultaneous collection of 32 SF+ 5, 33 + SF5 and 34 SF+ 5 ion beams (m/z ratios = 127, 128 and 129) allows the sulphur isotope data to be expressed in standard delta-notation (Eq. (1)):

Iron extraction protocols are described in Holmkvist et al. (2011) and Liu et al. (2018). Here we exploited two separate techniques to determine separate reactive iron fractions, each using separate subsamples of freeze-dried material. Firstly, the most readily acid-soluble Fe phases (e.g., amorphous Fe (hydro)oxides and ferrihydrite), were extracted via agitation with oxygen-free 0.5 M HCl for one hour. The Fe(II) content of these extracts was determined by spectrophotometry after complexing with 1, 10phenanthroline; while total Fe contents (i.e., Fe2+ + Fe3+ ) were determined after complexing with 1, 10-phenanthroline and 1% hydroxylamine hydrochloride. Poorly crystalline Fe(III) phases were calculated as the difference between the mixed valence and the Fe(II) fractions—A phase termed HCl reactive Fe(III). More crystalline Fe-oxides (e.g., goethite and hematite), termed dithionite

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Fig. 2. Solid-phase abundance (a. elemental sulphur, S0 ; c. pyrite and acid volatile sulphur, AVS; e. iron-oxides) and triple-sulphur isotope data (b. elemental sulphur, S0 ; d. pyrite) from core 973-4. The colour coded data points denote the location of the samples relative to the SMT. Here, data from within the SMT (blue) are separated from older (green) and younger (yellow and green) samples by those that define the (paleo-)S-front (red). Open and closed symbols in b and d denote 33 S and δ 34 S, respectively. The horizontal bars represent the (paleo-)SMT, which is inclusive of the modern SMT. The vertical bars between the panels position the coarse layer identified in Fig. 1a. Iron-oxide and AVS abundance data are adopted from Liu et al. (2018) and Zhang et al. (2014), respectively. SMT—sulphate methane transition; S-front—sulphidisation front; OSR—organoclastic sulphate reduction.

reactive Fe, were quantified by atomic absorption spectrometry following a two-hour mixed dithionite-citrate-acetic acid treatment (Raiswell et al., 1994). Irrespective of the measurement technique, replicate analyses returned a RSD of better than 5% for each reactive Fe pool. All solid-phase data were expressed relative to the dry mass of the sample. Mineral aggregates were identified and handpicked from the wet-sieve-derived coarse fraction (> 65 μm) using a stereomicroscope. The morphology of these handpicked minerals was investigated using a FEI Quanta 450 FEG scanning electron microscope (SEM) equipped with a secondary electron detector. Elemental maps were generated by energy-dispersive X-ray spectroscopy (EDS). 3. Results The concentrations of elemental sulphur determined by HPLC and gravimetrical methods show similar distribution patterns with two prominent peaks at ∼730 and ∼900 cmbsf (Fig. 2a). Concentrations of elemental sulphur and AVS reach maximums at ∼900 cmbsf (Figs. 2a, 2c). The δ 34 S values of elemental sulphur increase from −10.4 to 36.9 between 727 and 882 cmbsf with an associated decrease in 33 S values (Fig. 2b). Between 882 and 926 cmbsf, however, δ 34 S values of elemental sulphur are observed to decrease to 6.5 in concert with an increase in 33 S values. Both δ 34 S and 33 S values of elemental sulphur are more-or-less invariant below 1000 cmbsf (Fig. 2b). Throughout the upper ∼100 cm and the lower ∼500 cm (i.e., 900–1400 cmbsf; Fig. 2c) of core 973-4, pyrite concentrations range from 1 to 4 μmol FeS2 g−1 , with both δ 34 S and 33 S values around 0 (δ 34 S = −4.0 ± 5.7, 33 S = 0.006 ± 0.024; Figs. 2d, 3b). Between ∼100 and ∼600 cmbsf (Fig. 2c), pyrite concentrations increase gradually, coincident with a decrease in δ 34 S values down to −46.0 and an increase in 33 S values up to 0.12 (Figs. 2d, 3b). At roughly 600 cmbsf (Fig. 2c), pyrite contents increase sharply from ∼ 20 μmol FeS2 g−1 to ∼ 100 μmol FeS2 g−1 , forming a plateau that extends to ∼900 cmbsf before relaxing back to background levels. Superimposed on this plateau, discrete peaks, defining maxima in pyrite concentrations, are observed at

709 and 857 cmbsf (Fig. 2c). Between roughly 680 and 880 cmbsf, the δ 34 S and 33 S values of pyrite approach the isotopic composition of seawater sulphate (δ 34 S = 21, 33 S = 0.05; Figs. 2d, 3b). Yet by 882 cmbsf, the δ 34 S values of pyrite begin to exceed that of seawater sulphate (21). Notably, negative 33 S-δ 34 S correlations preserved in pyrite are seen at 612–671 cmbsf and 857–882 cmbsf (Figs. 2d, 3b). The concentrations of Fe-oxides quantified using the two different extraction methods show broadly consistent trends (Fig. 2e): Concentrations of HCl reactive Fe(III) and dithionite reactive Fe are both lower between roughly 500 and 900 cmbsf, while their respective concentrations increase sharply at ∼900 cmbsf and remain high over the lower ∼500 cm of core 973-4. In detail, however, our Fe extraction techniques fail to discriminate between highly-reactive iron and much less reactive sheet-silicate-iron (e.g., Raiswell et al., 1994). Given that both glauconite and chlorite have been identified in core 973-4 (Liu et al., 2018), the reactive iron contents reported in Fig. 2e represent hypothetical maximums and, in fact, the broad depletion seen at ∼600–900 cmbsf may indicate the complete exhaustion of readily pyritisable iron. Moreover, since AVS oxidises to Fe-oxides during exposure to air, the obvious peak of Fe-oxides at ∼900 cmbsf (Fig. 2e) likely reflects AVS oxidation during sample storage and treatment rather than a primary package of oxidant-rich sediments (Liu et al., 2018). 4. Discussion 4.1. Formation of elemental sulphur The solid-phase distributions of sulphur and iron display distinct variability throughout core 973-4. Excluding the coarse layer at 455–605 cmbsf, the sedimentation rate and the TOC content of core 973-4 are broadly invariant. Therefore, changes in the sedimentation rate or TOC availability cannot be invoked to explain the observed isotopic variability (Figs. 1, S2). Rather, the pronounced pyrite enrichment observed between 600 and 900 cmbsf is most readily explained by sulphate reduction coupled to AOM at the SMT (e.g., Lin et al., 2017; Riedinger et al., 2017).

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Fig. 3. The triple-sulphur isotope (δ 34 S vs. 33 S) systematics of elemental sulphur (a) and pyrite (b) isolated from core 973-4. Data points are colour-coded displaying the depth from which they were derived. Orange dashed arrows illustrate the OSR-driven evolution of sulphide from a seawater sulphate (SW) precursor (grey star). The black dashed arrow in a illustrates the depth-dependant isotopic evolution displayed at the S-front. Grey shaded areas in b include data points at the margins of the paleo-SMT (i.e., paleo-S-front). The black dashed line in b depicts a mixing trajectory between OSR- and AOM-derived end-members. The five square symbols in b with the small circle ornamentation indicate pyrite samples within the coarse layer. The paleo-SMT in b is inclusive of the modern SMT. Two sigma uncertainties in 33 S are illustrated, while δ 34 S uncertainties are encompassed by individual data points. Abbreviations and the colour scheme follow Fig. 2.

The elevated pyrite concentrations seen at the SMT in core 9734 (Fig. 2c) imply that the SMT has been generally stable, focusing AOM at roughly the same depth for perhaps thousands of years (März et al., 2008; Liu et al., 2018). Stabilisation of the SMT at this depth is further supported by the presence of 13 C-depleted carbonates (δ 13 C = −7.0 to −3.2 VPDB; total inorganic carbon) and a broad decrease in magnetic susceptibility seen between 600 and 900 cmbsf (Zhang et al., 2018). Moreover, the inscription of a static SMT is also evident in the Fe record. Here, the sulphide generated from AOM reduces reactive Fe-phases, converting them to pyrite. Sustained AOM, therefore, progressively consumes the most reactive Fe-phases, as recorded in core 973-4 as a broad depletion in Fe-oxides at 600–900 cmbsf (Fig. 2e). As reactive Fe-oxides become more scarce, sulphide produced by sulphate-driven AOM can no longer be retained within the SMT and diffuses towards its boundaries (Riedinger et al., 2017). This process promotes further reductive dissolution of Fe-oxides, broadening the chemostratigraphic fingerprint of AOM, forming the so-called sulphidisation front (Sfront) via the sequestration of sulphur as elemental sulphur and Fe monosulphides at the periphery of the SMT (Eq. (3); Jørgensen et al., 2004; Riedinger et al., 2017).

3H2 S + 2FeOOH → S0 + 2FeS + 4H2 O

(3)

Porewater sulphate profiles from core 973-4 and neighbouring sites are broadly similar, revealing consistent SMTs between ∼700 and ∼880 cmbsf (Fig. 1c). Moreover, core 973-4 features two prominent peaks of elemental sulphur, defining active S-fronts at ∼730 and ∼900 cmbsf (Fig. 2a). This differs from the zone of pyritisation at ∼600–900 cmbsf in core 973-4, suggesting that the SMT currently resides deeper within the sediment pile than it once did, and the zone of pyrite enrichment archives the previous position of the SMT (Liu et al., 2018), which we refer to as the paleo-SMT for clarity. While Fe-oxide-fuelled sulphide oxidation should occur both below and above the SMT, in core 973-4, the upper S-front is extremely narrow and not well developed. The differential development of the S-front most likely reflects the lower availability of reactive Fe-oxides at/around ∼700 cmbsf. Owing to the limited data constraining the upper S-front, we focus the ensuing discussion on the lower S-front, which is defined by pronounced peaks

in acid volatile sulphur (AVS) and elemental sulphur abundances at ∼900 cmbsf (Figs. 2a, 2c). Interestingly, elemental sulphur extracted in proximity to the S-front (882–926 cmbsf) displays distinctive sulphur isotope systematics. Here, with distance from the SMT, δ 34 S values are observed to decrease in concert with an increase in 33 S values (Figs. 2b, 3a). As sulphide diffusing downwards through the sediment column is 34 S-enriched relative to the free sulphide present at that horizon (Jørgensen, 1979; Jørgensen et al., 2004), and Fe-driven sulphide oxidation can only induce small isotope effects (Fry et al., 1986), the observed trend of decreasing δ 34 S cannot be readily explained in terms of a simple process, and requires the operation of an active sulphur cycle. 4.2. The isotopic fingerprint of an oxidative sulphur cycle Potential explanations for the observed negative 33 S-δ 34 S correlation (Figs. 2b, 3a) include sulphate reduction with suppressed isotope fractionation, bacterial sulphide oxidation, and iron-driven sulphide oxidation with simultaneous sulphur disproportionation. Generally, laboratory experiments suggest that faster cell-specific sulphate reduction rates (csSRR) yield smaller sulphur isotope fractionations (e.g., Sim et al., 2011b; Leavitt et al., 2013). In reality, however, most of these experiments are performed at a much higher thermodynamic drive than in natural settings, promoting much higher csSRR than is generally observed in marine sediments (Jørgensen et al., 2019, and references therein). In fact, even under the sulphate impoverished conditions and the low csSRR expected around the SMT, theoretical work predicts that large sulphur isotope fractionation should be expressed (Wing and Halevy, 2014). Consequently, sulphur isotope fractionation between sulphate and sulphide would be expected to approach equilibrium predictions (∼70) and thus variable magnitude of sulphur isotope fractionation is an unlikely explanation of the unique sulphur isotope systematics witnessed at the S-front. Typically, microbial sulphide oxidation is considered to yield a small sulphur isotope fractionation (e.g., < 5; Fry et al., 1986) and would often be precluded. Nevertheless, Pellerin et al. (2019) proposed that bacterial sulphide oxidation could foster pronounced sulphur isotope fractionation via two mechanisms: (1) intracellular disproportionation of sulphur intermediates or (2) via reverse dissimilatory sulphate-reduction. The former is similar to Fe-driven sulphide oxidation with attendant disproportionation, yet it is con-

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Fig. 4. Scanning electron photomicrograph and elemental maps of elemental sulphur aggregates from 729 cmbsf (upper S-front; a–c) and 606 cmbsf (d). The distribution of phosphorus (c) demonstrates that the unique morphology and sulphur distribution within the sulphur aggregate (a and b) is not an artefact of uneven coating. No enrichment of iron or other elements was observed within the sulphur aggregate in a.

fined to the cell and thus yields no extracellular elemental sulphur. Examination of sediments from core 973-4 via SEM reveals crystalline elemental sulphur aggregates around the SMT (Fig. 4). The morphology and elemental distribution of these aggregates differ from intracellular sulphur granules (e.g., Nims et al., 2019), implying a different formation mechanism. Given that microbial sulphide oxidation is operational in surficial sediments and thus spatially separated from the S-front (e.g., Jørgensen et al., 2019), we suggest that these elemental sulphur aggregates were the product of Fe-oxide-fuelled sulphide oxidation (Eq. (3)). Accepting that this sulphide oxidation typically imparts a minor isotope effect (e.g., Fry et al., 1986), the isotopic composition of the elemental sulphur should approximate that of coexisting free sulphide. Accordingly, we assume that the isotopic composition of elemental sulphur at the lower boundary of SMT (i.e., 882 cmbsf) records that of free sulphide descending from the SMT (i.e., δ 34 S = 37). This free-sulphide is oxidised at the Sfront by Fe-oxides, producing elemental sulphur (Zopfi et al., 2004; Riedinger et al., 2017), which successively undergoes disproportionation, producing 34 S-depleted free sulphide and 34 S-enriched sulphate (Thamdrup et al., 1993; Canfield and Thamdrup, 1994). Given the rapidity of sulphide oxidation, and absence of associated isotopic fractionation, we combine elemental sulphur and free sulphide into a single pool, which we term the reduced sulphur pool hereafter. Continued sulphide oxidation with concomitant disproportionation of the elemental sulphur product can be approximated by equation (4):

ln(RR /RR,0 ) = (α − 1) ln(f)

(4)

where RR represents the isotopic composition of the residual reduced sulphur pool and RR,0 is the initial isotopic composition of the reduced sulphur pool. The fraction of 32 S in the reduced sulphur pool, and the fractionation factor between the sulphide and sulphate are denoted f and α , respectively (Mariotti et al., 1981; Ono et al., 2006). Adopting the fractionation factor (34 α = 1.022) from Canfield and Thamdrup (1994) allows us to approximate the hypothetical δ 34 S evolution of the reduced sulphur pool. Here, some part of the 34 S-enriched reduced sulphur pool is lost as sulphate, while the remainder is cycled between free sulphide and elemental sulphur, eventually accumulating in the form of 34 Sdepleted elemental sulphur (Fig. 5). For instance, the blue curve in Fig. 5 predicts that about 74% of the reduced sulphur pool is consumed to generate the lowest δ 34 S (6.5) values observed in elemental sulphur, while the purple curve implies that the δ 34 S of the accumulated sulphate may be roughly 47. Mass-dependent sulphur isotope fractionation is described by a power law, whose exponent (λ = ln33 α / ln34 α ) is approximately 0.515 under typical low-temperature equilibrium conditions (Ono et al., 2006). Likewise, at least in culture, microbially mediated disproportionation is found to approximate equilibrium predictions (λ ≈ 0.515; Johnston et al., 2005). Simplistically, disproportiona-

Fig. 5. A Rayleigh fractionation model showing the isotopic evolution of the residual reduced sulphur pool (blue), as well as the instantaneous (red) and the cumulative sulphate pools (purple). Rather than providing quantitative constraints, this figure is intended to illustrate the hypothetical isotopic evolution of free sulphide and elemental sulphur during oxidation and disproportionation (blue arrow).

tion is a branching process where two reservoirs (i.e., sulphide and sulphate) are produced from a single reservoir (i.e., elemental sulphur and other sulphur intermediates). As λ is roughly 0.515, the products should attain more positive 33 S values relative to their precursors due to a non-linear effect in the δ 34 S vs. 33 S coordinate (Ono et al., 2006). Culture experiments corroborate this, with both reduced and oxidised products displaying 33 S enrichments (i.e., higher 33 S) relative to the parent sulphur intermediates (Johnston et al., 2005; Fig. S3). Therefore, it follows that, combined oxidation and disproportionation will yield a reduced sulphur pool featuring lower δ 34 S values and higher 33 S values. This combination of processes, therefore, is the most parsimonious explanation of the negative 33 S-δ 34 S correlation (Fig. 3a) recorded by the elemental sulphur isolated from the S-front. Unfortunately, tracing the minor sulphur isotopic evolution of dissolved sulphate remains difficult. To date disproportionation has received little laboratory attention, and its associated isotope effect has only been quantified in a single study (Johnston et al., 2005). More importantly, the Rayleigh behaviour (Eq. (4)) describes closed system behaviour, which is strictly invalid for marine sediments that typically remain open to chemical exchange (Jørgensen, 1979; Canfield, 2001). For illustration, in a marine setting, descending 34 S-enriched sulphate/sulphide would increase the δ 34 S of their respective sulphur pools at the S-front causing deviations from their modelled compositions. Similarly, removal processes of unknown importance, such as barite precipitation at the S-front (e.g., Dickens, 2001), will also drive data–model mismatch. Given

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these complications, we stress that we do not quantitatively advocate this approach, and stress that Fig. 5 should be considered as an illustrative approximation. A more accurate reactive transport model, of course, would require explicit parameterisation of the diffusive supply of reactants and competing reactions involved in their consumption (e.g., Jørgensen et al., 2004). The necessary data to drive this model are not available from site 973-4. Predicting the minor sulphur isotope evolution of dissolved sulphate remains an open question and future efforts should centre around experimental characterisation of sulphur disproportionation, as well as a more extensive isotopic survey of coupled porewater and solidphase sulphur species. In continental slope settings, the accumulation of porewater sulphide is typically restricted to a narrow interval in proximity to the SMT. In detail, dissolved sulphide is usually abundant within the SMT (up to 0.7 mM), dropping precipitously at its lower margin (März et al., 2008; Riedinger et al., 2017). In the absence of reliable porewater data, the solid phase records from core 973-4 imply an analogous dissolved sulphide distribution. For example, within core 973-4 elemental sulphur and AVS contents peak at the lower boundary of the SMT (∼882 cmbsf), indicating that free sulphide was sufficiently concentrated to act as a sulphur-donor to Fe monosulphide and elemental sulphur. Contrastingly, the presence of vivianite, an authigenic iron phosphate mineral that is unstable in the presence of sulphide (Berner, 1981), below 920 cmbsf demonstrates that free sulphide concentrations decrease rapidly with depth (Liu et al., 2018). While sulphide replete conditions at the periphery of the SMT would most likely inhibit disproportionation (Thamdrup et al., 1993), its depletion with depth would provide a niche where disproportionation may be energetically favourable. Consequently, oxidative sulphur cycling— recycling sulphide to sulphate via disproportionation—should be more favourable down-core, promoting the concomitant decrease in δ 34 S, and increase in 33 S, values observed within the reduced sulphur pool. We conclude, therefore, that the decrease in the δ 34 S values around the S-front (Fig. 2b), and the broad negative 33 S-δ 34 S correlation manifest in elemental sulphur (Fig. 3a), were inherited from an intensely cycled sulphide precursor. 4.3. Sulphur isotopic signature of pyrite Throughout the upper ∼100 cm and the lower ∼500 cm of core 973-4 (Fig. 2c), pyrite concentrations are low, with δ 34 S and 33 S values around 0 (Figs. 2d, 3b). Such low pyrite contents, in concert with near-primordial δ 34 S and 33 S values, can be reconciled with a predominantly external source of potentially igneous origin (e.g., Ueno et al., 2008). The recent discovery of hydrothermal vents in the South China Sea provides a potential proximal igneous source of pyrite (Wang and Jian, 2019). Alternatively, storm reworking of sediments and subsequent rapid transport of isotopically heavy pyrite from the shelf via bottom currents (e.g., Fike et al., 2015) may serve as an additional source of externally-derived pyrite to the sediment column. Although these externally-sourced pyrites are likely to be omnipresent throughout the core, their relative importance clearly changes down-core. For example, between roughly 100 and 600 cmbsf (Fig. 2c), pyrite contents increase progressively with depth, coincident with a decrease in δ 34 S values, reflecting the addition of 34 S-depleted pyrite (Figs. 2d, 3b). Organic matter mineralisation via microbial sulphate reduction is vigorous in the upper layers of the sediment-pile, following the depletion of other, more energetically favourable, electron acceptors (e.g., Jørgensen and Kasten, 2006). In deep-sea sediments with low organic contents (Fig. 1b), depletion of these oxidants (i.e., oxygen, nitrate, manganese- and iron-oxides) occurs slowly, restricting the main sulphate reduction zone to a few meters below the sediment surface (Jørgensen and

7

Kasten, 2006). Thus, the gradual decrease in pyrite δ 34 S values advocates the onset, and increasing importance, of organoclastic sulphate reduction (e.g., Brunner and Bernasconi, 2005; Sim et al., 2011a), acting as a sink for 34 S-depleted sulphur. Culture experiments have shown that λ associated with OSR is lower than equilibrium predictions (0.515; Farquhar et al., 2003; Johnston et al., 2005; Sim et al., 2011b; Leavitt et al., 2013), resulting in sulphide, and ultimately pyrite, with 33 S values that exceed those of the precursor sulphate (Johnston et al., 2008; Lin et al., 2017). This explains the elevated 33 S values (∼0.12; Figs. 2d, 3b), relative to seawater sulphate (0.05; Ono et al., 2006; Tostevin et al., 2014) that are observed between ∼100 and ∼600 cmbsf. Even within the coarse layer of presumed turbidite origin (455–605 cmbsf; Fig. 1a), the δ 34 S and 33 S values approximate those seen between 100 and 600 cmbsf, requiring that turbidite-derived pyrite was minor and subsequently overprinted by OSR. In marine sediments, sulphate reduction via AOM consumes downward-diffusing sulphate, preferentially removing 32 S, enriching the residual sulphate, and any subsequent sulphide, in 34 S (Jørgensen et al., 2004; Antler et al., 2015). The differential diffusion of 32 S, 33 S and 34 S even causes the reduced sulphur pool within the SMT to approach the isotopic composition of seawater sulphate (Figs. 2, 3; Jørgensen, 1979; Masterson, 2016). Deusner et al. (2014) suggested that sulphate-driven AOM is likely to induce sulphur isotope fractionations exceeding 60 between sulphate and sulphide at the SMT, supporting the formation of 34 S-enriched pyrite with a large sulphur isotope fractionation in core 973-4 (see section 4.2; Wing and Halevy, 2014; Pellerin et al., 2018; Jørgensen et al., 2019). As predicted, this is also seen at the base of the SMT (i.e., 882 cmbsf), where δ 34 S values of both pyrite and elemental sulphur exceed that of seawater sulphate (21; Figs. 2b, 2d). Assimilated, the sulphur isotope systematics of pyrite clearly reveal its genesis. At site 973-4 three pyrite end-members are recognised: (1) externally-sourced pyrite (δ 34 S = 0, 33 S = 0), (2) OSR-derived pyrite (δ 34 S = −46.0, 33 S = 0.12) and (3) AOM-derived pyrite (δ 34 S = 21, 33 S = 0.05). Mixing between externally-sourced and OSR-derived pyrite explains the bulk δ 34 S values ranging from −16.4 to −6.0 in the uppermost ∼100 cm. In δ 34 S-33 S-space mixing is non-linear, producing curved trajectories that drive mixing products to lower 33 S values (Fig. 3b; Ono et al., 2006). Much of the data in this three-isotopespace is encompassed by the mixing field generated by combining these end-members, thereby providing a satisfactory explanation. The higher 33 S values (up to 0.2) found at the fringes of the paleo-SMT, however, fall outside of this mixing envelope and thus require an alternate explanation (Fig. 3b). Exemplified in Fig. 3b, akin to the elemental sulphur record, a negative 33 S-δ 34 S correlation is also seen within the pyrite record. Unlike the elemental sulphur record where the highest 33 S values are found at ∼730 and ∼960 cmbsf, the highest 33 S values in pyrite are observed at ∼630 and ∼860 cmbsf. Accepting that the upper peak of elemental sulphur (∼730 cmbsf) corresponds to contemporary S-front, the current upper and lower S-fronts are currently ∼100 cm deeper in the sediment-pile than they were previously. Yücel et al. (2010) argued that providing elemental sulphur was present to form polysulphide, then the polysulphide pathway would dominate pyrite formation. Elemental sulphur is abundant at the fringes of the SMT in core 973-4, providing ideal conditions for rapid polysulphide-mediated pyrite formation (e.g., Yücel et al., 2010). Given that polysulphides are formed via reaction between elemental sulphur and free sulphide (Zopfi et al., 2004), polysulphides have a free sulphide dependency. Free sulphide availability, however, decreases rapidly from the centre of SMT to the S-front (see section 4.2), where it is further consumed by the upward flux of Fe2+ and the reductive disso-

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lution of reactive Fe-oxides. Accordingly, the efficient conversion of elemental sulphur to pyrite is hindered at the S-front due to the scarcity of the necessary free sulphide. Sluggish pyrite formation at the S-front coupled with the slow descent of the SMT (Liu et al., 2018) results in a non-steady-state system, promoting the depth-dependant differences in the 33 S-δ 34 S systematics of co-existing elemental sulphur and pyrite. Moreover, the lack of free sulphide below ∼920 cmbsf also inhibits transformation of elemental sulphur and Fe monosulphide to pyrite (Riedinger et al., 2017), promoting the large observed isotopic distinction between the two species below the SMT (Figs. 2b, 2d). Taken together, we propose that prolonged oxidative sulphur cycling occurs at the fringes of the SMT, leading to the formation of the negative 33 S-δ 34 S correlation seen in pyrite records, followed by pyritisation at the former S-front (paleo-S-front). Similar high 33 S values are also seen in elemental sulphur, AVS and pyrite extracted from the S-front of multiple sites within the Bornholm Basin, Baltic Sea (unpublished data). Drawing analogy with core 973-4, the transfer of triple-sulphur isotopic signatures diagnostic of oxidative sulphur cycling to refractory solid-phase species has the potential to be a widespread phenomenon and hence a diagnostic geological tracer providing it enters the geological record. 4.4. Implications for contemporary and ancient biogeochemical cycles Core 973-4 contains abundant reactive Fe-oxides, which along with stable methane fluxes, have supported a relatively stable SMT since the LGM. The distribution and importance of OSR-derived pyrite are limited in organic-poor but methane-rich sediments, minimising the impact of mixing with pyrite derived from other sources. These background conditions have sustained an oxidative sulphur cycle, whose isotopic fingerprint has been transferred from elemental sulphur to pyrite via the polysulphide pathway (Figs. 2–3). Informed primarily by porewater data, previous studies have invoked a deep oxidative sulphur cycle in marine sediments (Riedinger et al., 2010; Holmkvist et al., 2011; Treude et al., 2014; Brunner et al., 2016; Pellerin et al., 2018). Unfortunately, with perhaps the exception of diagenetically-susceptible carbonate-associated sulphate (e.g., Fike et al., 2015), porewater sulphates have little preservation potential. Sedimentary pyrite, by contrast, is ubiquitous in modern and ancient sedimentary successions. From the preceding discussion, we propose that oxidative sulphur cycling leads to 33 S-enriched sedimentary sulphides with negative 33 S-δ 34 S systematics. Therefore, these isotope systematics, when combined with solid-phase sulphur and iron distributions, present a novel way to potentially trace the operation of the cryptic sulphur cycle in settings where more labile phases have been consumed or compromised. The isotopic record of sedimentary sulphur-bearing phases, particularly pyrite, is widely used as a proxy to reconstruct ocean chemistry (e.g., Canfield, 2001; Johnston, 2011; Fike et al., 2015). Nevertheless, extracting information pertaining to the operation of Earth’s sulphur cycle is often hindered because of complex diagenetic overprints (e.g., Fike et al., 2015). From the preceding discussion, we suggest that the sulphur isotope data from methanerich but organic-poor sediments predominantly reflect the interplay between OSR, sulphate-driven AOM and oxidative sulphur cycling rather than surficial processes. Mass-dependently fractionated pyrites with high 33 S values up to 0.2, as seen in the South China Sea, are a common feature of the sedimentary record throughout the last c.2.3 billion years of Earth’s history (e.g., Johnston, 2011). While there may be several alternate explanations for these elevated 33 S values, oxidative sulphur cycling clearly leads to the formation of pyrite with high 33 S values and distinctive 33 S-δ 34 S systematics; whether, these signals can be resolved after burial and compaction, however, remains to be seen. Multiple

Fig. 6. A schematic representation of biogeochemical cycling active at site 9734, South China Sea. Idealised authigenic mineral and porewater distributions (cf. Riedinger et al., 2017) are illustrated in the left and middle portions of the figure, while the dominant reactions in proximity to the SMT are given on the right. Abbreviations, again, follow Fig. 2.

sulphur isotope data from contemporary settings is in its infancy, yet, given time, this growing database will undoubtedly illuminate processes that are not resolved by δ 34 S data alone. The minor sulphur isotope record, therefore is anticipated to provide new insight into sulphur cycling and, by extension, permit a better understanding of the ancient sulphur cycle. Methanogenesis was traditionally considered to be the predominant mineralisation process below the SMT (Jørgensen and Kasten, 2006). Besides cryptic sulphur cycling, anaerobic oxidation of methane coupled to iron reduction has recently been shown to be an important process active beneath the SMT (e.g., Beal et al., 2009; Egger et al., 2015). At site 973-4, Liu et al. (2018) reported pervasive vivianite authigenesis in the Fe-oxide-rich sediments below the S-front (∼920 cmbsf), implying deep iron reduction and potentially iron-mediated anaerobic oxidation of methane (Fe-AOM). The relationship between Fe-AOM and oxidative sulphur cycling, however, is essentially unknown. It is possible that methane oxidation by Fe-oxides proceeds immediately beneath the SMT (∼900–1000 cmbsf), catalysed by sulphate (Beal et al., 2009; Sivan et al., 2014). While further microbial studies are required to understand the interplay between the sub-SMT cycles of carbon, sulphur, and iron, our mineralogical and isotopic observations support recent inferences gleaned from porewater chemistry and microbial assays (Holmkvist et al., 2011; Treude et al., 2014), suggesting that the oxidative sulphur cycle and iron-mediated AOM may play a key biogeochemical role beneath the SMT (Fig. 6).

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5. Conclusion

References

Combining multiple sulphur isotope analyses with solid-phase sulphur and iron abundances, uncovers the operation of ironoxide-fuelled sulphide oxidation at the periphery of the SMT. The observed negative 33 S–δ 34 S correlation recorded by elemental sulphur and pyrite reveals iron-driven sulphide oxidation with concomitant disproportionation of the elemental sulphur product— providing the first evidence of the preservation and isotopic signature of the sub-seafloor cryptic sulphur cycle. Analogous negative correlations, therefore, might serve as a diagnostic tracer for oxidative sulphur cycling, allowing its importance to be evaluated in both contemporary and ancient sedimentary environments. Besides the oxidatively-derived negative 33 S–δ 34 S correlation, our multiple sulphur isotope data disclose several mechanisms of pyrite genesis (e.g., externally-sourced, OSR-derived and AOM-derived), demonstrating the utility of multiple sulphur isotopes to provenance mechanisms of pyrite genesis. Along with pervasive vivianite authigenesis, we suggest that the oxidative sulphur cycle and iron-mediated AOM may play key roles in carbon, sulphur and iron cycling beneath the SMT. Thus, a better understanding of these processes in the deep biosphere is paramount to test longstanding hypotheses linking diverse microbial metabolisms and major biogeochemical cycles under energy-limited conditions (e.g., Jørgensen and Kasten, 2006; Treude et al., 2014). These processes, in turn, may be recorded in the isotopic composition by authigenic minerals, allowing their importance to be evaluated throughout Earth’s history.

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Declaration of competing interest The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper. Acknowledgements We acknowledge Xiaoping Liao, William Olszewski, Chao Li, Caixiang Zhang and Zihu Zhang who provided invaluable technical assistance. Genming Luo, David T. Wang, Tina Treude, Alexandra Turchyn, Zhou Wang, Alyssa Findlay and Qi Lin are thanked for formative discussions. Sulphur isotope analysis was performed in the Laboratory of Stable Isotope Geobiology at MIT, funded by the Alfred P. Sloan Foundation via the Deep Carbon Observatory (S.O.). Sampling and initial wet chemical extractions were funded by the State Key R&D Project of China (Grant 2016YFA0601102), National Natural Science Foundation of China (Grants 41772091, 41802025), China National Gas Hydrate Project (Grant DD20160211), and the Fundamental Research Funds for National Universities, China University of Geosciences, Wuhan (Grant CUGCJ1710). Jia.L. acknowledges financial support from the international exchange program administered by School of Earth Sciences, CUG, Wuhan. G.I. gratefully recognises continued support from Roger Summons, and financial backing from the Simons Collaboration on the Origins of Life. G.A. acknowledges financial support from the Israel Science Foundation [2361/19]. This represents Seolfor Solutions contribution #2. Editorial handling by Itay Halevy and reviews by Morgan Raven and an anonymous reviewer are gratefully acknowledged: their insight, expertise and rigour undoubtedly improved the quality and clarity of the final manuscript. Appendix A. Supplementary material Supplementary material related to this article can be found online at https://doi.org/10.1016/j.epsl.2020.116165.

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