Large accumulations of 34S-enriched pyrite in a low-sulfate marine basin: The Sturtian Nanhua Basin, South China

Journal Pre-proofs Large accumulations of 34S-enriched pyrite in a low-sulfate marine basin: the Sturtian Nanhua Basin, South China Ping Wang, Thomas J. Algeo, Yuansheng Du, Wenchao Yu, Qi Zhou, Yongjun Qin, Yuan Xu, Liangjun Yuan, Wen Pan PII: DOI: Reference:

S0301-9268(19)30004-X https://doi.org/10.1016/j.precamres.2019.105504 PRECAM 105504

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Precambrian Research

Received Date: Revised Date: Accepted Date:

4 January 2019 9 August 2019 11 October 2019

Please cite this article as: P. Wang, T.J. Algeo, Y. Du, W. Yu, Q. Zhou, Y. Qin, Y. Xu, L. Yuan, W. Pan, Large accumulations of 34S-enriched pyrite in a low-sulfate marine basin: the Sturtian Nanhua Basin, South China, Precambrian Research (2019), doi: https://doi.org/10.1016/j.precamres.2019.105504

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Large accumulations of 34S-enriched pyrite in a low-sulfate marine basin: the Sturtian Nanhua Basin, South China

Ping Wang a, Thomas J. Algeo a, b, c, Yuansheng Du a, Wenchao Yu a, Qi Zhou d, Yongjun Qin e, Yuan Xu a, Liangjun Yuan f, Wen Pan f

a

State Key Laboratory of Biogeology and Environmental Geology, School of Earth Sciences, China University of Geosciences, Wuhan 430074, China

b

State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences-Wuhan, Wuhan 430074, China

c

Department of Geology, University of Cincinnati, Cincinnati, OH 45221-0013, USA

d

Guizhou Bureau of Geology and Mineral Exploration and Development, Guiyang 550004, China

e

Guizhou Geological Survey, Guiyang 550018, China

f

103 Geological Party, Guizhou Bureau of Geology and Mineral Exploration and Development, Tongren 554300, China

Abstract

The Cryogenian of the Nanhua Basin (eastern Guizhou Province, South China) consists of a continuous succession of glacial and post-glacial deposits from the Sturtian Glaciation, including diamictite of the Tiesi’ao Formation and Mn-carbonate and black shale of the



Corresponding author, Tel.: +86 13971241916, fax: +86 27 87481365. E-mail address: [email protected] (Y. Du).

overlying lower Datangpo Formation. Here, we analyzed the sulfur chemistry of these units with the goal of understanding regional to global changes in the marine sulfur cycle accompanying a Snowball Earth event. The study units are characterized by elevated carbonate-associated sulfate (CAS) δ34S (mean +56.0‰, range +49.6 to +62.6‰) and pyrite δ34S compositions (mean +57.5‰, range +48.8 to +66.8‰). Both CAS and pyrite δ34S show water-depth gradients, with mean values increasing from the shallower Lijiawan area (CAS: +49.3‰; pyrite: +44.5‰) to the deeper Xixibao (CAS: +57.9‰; pyrite: +59.8‰) and Gaodi areas (CAS: +62.4‰; pyrite: +61.6‰), reflecting a density-stratified water column with limited vertical mixing. △34S values (i.e., δ34SCAS – δ34Spy) range from –6.5‰ to +8.0‰ with a mean of +0.7‰. These features, which are similar to those for coeval strata globally, are consistent with low seawater sulfate concentrations, but several additional features of the Nanhua Basin deposits do not conform to existing Cryogenian sulfur-cycle models: (1) high total sulfur content (mean 2.2 ± 1.1%), which is difficult to reconcile with low seawater sulfate, and (2) frequent negative △34S values, which indicate that in situ microbial sulfate reduction (MSR) cannot have been the sole control on pyrite δ34S. These features point to quantitatively important hydrothermal sulfur inputs to the Nanhua Basin watermass. Based on these considerations, we propose a new sulfur-cycle model for the Sturtian Nanhua Basin in which hydrothermal emissions supplied large amounts of 34S-enriched H2S to the water column. The released H2S was partly precipitated as syngenetic framboidal pyrite and partly oxidized to sulfate that was removed to the sediment as CAS, thus accounting for the unusual combination of high total sulfur concentrations, similar strongly 34S-enriched sulfur-isotopic compositions for CAS and pyrite, and frequent negative △34S values. As a result of low

seawater sulfate concentrations, both δ34SCAS and δ34Spy developed water-depth gradients through vertical mixing of strongly 34S-enriched hydrothermal sulfide from deep-graben vents with moderately 34S-enriched sulfate from the global ocean. Our model provides new insights into Sturtian-glacial sulfur cycling processes within a semi-restricted marine basin that are likely to have wider applicability to Neoproterozoic marine systems.

Keywords: sulfur isotope; CAS; hydrothermal; thermochemical sulfate reduction; H2S; Neoproterozoic

1. Introduction

The Neoproterozoic Era (~1000-542 Ma) experienced a rise of atmospheric pO2, metazoan evolution, massive metallogenic events, and two major glaciations―the Sturtian Glaciation at ~720-663 Ma and the Marinoan Glaciation at ~654-635 Ma (Kennedy, 1996; Kennedy et al., 1998; Hoffman and Schrag, 2002; Zhang et al., 2008; Macdonald et al., 2010; Shields-Zhou et al., 2012; Lan et al., 2014; Rooney et al., 2015)―that may have been global in extent, as proposed by the Snowball Earth hypothesis (Kirschvink, 1992; Hoffman et al., 1998). Post-Sturtian (~663-654 Ma) marine deposits in Australia, Namibia, Canada, Greenland, and the UK are unusual in containing strongly 34S-enriched pyrite (to ca. +61‰) (Gorjan et al., 2000, 2003; Hurtgen et al., 2002, 2005; Sperling et al., 2016; Parnell and Boyce, 2017; Scheller et al., 2018). The Mn-carbonate unit of the basal Datangpo Formation of the Nanhua Basin of South China, which was deposited during the Sturtian deglaciation,

has yielded the heaviest δ34Spy compositions yet reported (up to +69‰) (Logan et al., 1995; Li et al., 1996, 1999; Chu et al., 2001, 2003; Chen et al., 2008; Shen et al., 2008; Feng et al., 2010; Li-C et al., 2012; Zhang et al., 2013; Wu et al., 2016). All studies to date have attributed this feature to low contemporaneous seawater sulfate concentrations and consequent small MSR-related fractionations. However, this unit contains remarkably high total sulfur (TS) concentrations (mean 2.2%, range 1.3-2.7%; note: all ranges reported in this study are 16th-84th percentiles in order to avoid the influence of outliers). The few sulfur concentration data that are available for coeval units globally are inconsistent, with relatively high values in the basal Arena Formation, East Greenland (mean 2.5%, range 1.9-2.9%; Scheller et al., 2018) but low values in the MacDonaldryggen Member of Elbobreen Formation, Svalbard (mean 0.6%, range 0-1.0%; Kunzmann et al., 2015). The present study was undertaken to investigate the origin of the unusual combination of high pyrite sulfur concentrations and 34S-enriched compositions in the Sturtian Nanhua Basin. We analyzed samples from three sites in northeastern Guizhou Province, South China (the Lijiawan, Xixibao and Gaodi areas; Fig. 1) that represent different depositional water depths. Our analysis included petrographic study and measurement of the δ34S compositions of both pyrite and carbonate-associated sulfate (CAS) as well as generation of whole-rock elemental data. Our main goals are (1) to explain the high concentrations and strong 34S enrichments of both pyrite and CAS in the study units in the context of a new sulfur-cycle model for the Sturtian Nanhua Basin, and (2) to examine the implications of this model for global-scale sulfur cycling in the aftermath of the Sturtian Glaciation.

2. Geological background

2.1. Paleogeography and basin stratigraphy

The supercontinent Rodinia, which was assembled between 1.3 and 0.9 Ga, broke up at ~820 Ma, generating many tectonic units including the Yangtze and Cathaysia Blocks (Fig. 2A-B; Dalziel, 1991; Hoffman, 1991; Moores, 1991; Li et al., 2008; Wang and Pan, 2009; Li-XH et al., 2012). During this breakup, the Nanhua Basin formed through intracontinental rifting between the Yangtze and Cathaysia blocks of the South China Craton. This basin, located along the present southeastern margin of the Yangtze Block (Fig. 2B), accumulated a Neoproterozoic succession comprising four rift stages (Wang and Li, 2003): the first stage marked rift initiation at ca. 820 Ma; the second (ca. 800 Ma) and third stages (ca. 780-750 Ma) represented the main rifting episodes; and the fourth stage (ca. 750-690 Ma) recorded the transition from a rift basin into a cratonic sag basin. The Nanhua Basin preserves a stratigraphically conformable succession of the interval from the Sturtian Glaciation (~720-663 Ma) to the Marinoan Glaciation (~654-635 Ma) (Zhang et al., 2008; Macdonald et al., 2010; Lan et al., 2014). The succession consists of, in ascending order, the Liangjiehe/Tiesi’ao, Datangpo, and Nantuo formations. The diamictite and sandstone of the Liangjiehe and Tiesi’ao formations are Sturtian-age glacial deposits, whereas the glaciomarine diamictite, siltstone, and sandstone of the Nantuo Formation are Marinoan-age glacial deposits. The Datangpo Formation represents interglacial sediments that accumulated between the Sturtian and Marinoan glaciations, comprising two members. The 1st Member is of a Mn-carbonate unit in which rhodochrosite layers alternate with

Mn-bearing black shale layers, overlain by a moderately organic-rich black shale (mean TOC 2.9%), and the overlying 2nd Member consists largely of gray mudstones and siltstones (Yu et al., 2016). The present study examines the uppermost ~10 m of the Tiesi’ao Formation and the 1st Member of the Datangpo Formation. The age of the Mn-carbonate unit of the Datangpo Formation has been determined by U-Pb dating of volcaniclastic zircons. A deep-water tuffaceous bed yielded a TIMS U-Pb age of 662.9±4.3 Ma (Zhou et al., 2004) and a SIMS U-Pb age of 667.3 ± 9.9 Ma (Yin et al., 2006), and a shallow-water tuff yielded an LA-ICP-MS age of 662.7±6.2 Ma (Yu et al., 2017) and CA-ID-TIMS age of 658.8 ± 0.5 Ma (Zhou et al., 2019). These ages are in good agreement and constrain the timing of termination of the Sturtian Glaciation to ~663 Ma (Rooney et al., 2014). A SHRIMP U-Pb age of 654.5±3.8 Ma from an ash bed in the uppermost Datangpo Fm. (immediately below the Nantuo glacial diamictite) provides an age constraint for the onset of the following Marinoan Glaciation (Zhang et al., 2008). These radiometric dates suggest a depositional interval of ~8.5 ± 4 Myr for the Datangpo Formation, which also represents the duration of the Sturtian-Marinoan interglaciation. The Datangpo Formation experienced significant burial temperatures during its history. Although terrestrial organic matter is not present in Cryogenian sediments, the reflectance of bitumen (Rb) can be used to estimate time-integrated maximum burial temperatures as follows. In the study units, bulk manganese ore contains balls of bitumen surrounded by early diagenetic actinomorphic chalcedony, which isolated the bitumen from fluid interaction during burial (see Section 4.1; Zhou et al., 2013; Wu et al., 2016). The bitumen balls yield Rb ranging from 3.24% to 3.34%, which is equivalent to vitrinite reflectances (Ro) of 2.51-2.57%

(Huang, 2016) based on the conversion equation: Ro = 0.668 × Rb + 0.346 (Hood et al., 1975). Maximum burial temperatures (Tmax) can be calculated from the Easy %Ro model: Tmax = (lnRo + 1.78) / 0.0124 + 5.9 lnHr, where Hr = 1 °C Myr‒1, which represents a mean relationship determined from multiple sedimentary basins (Burnham and Sweeney 1989; Sweeney and Burnham, 1990). This calculation yielded Tmax estimates of 256-258 °C for the present study units. Given a surface temperature of 20-30 °C and an average intracratonic geothermal gradient of ~30 °C km‒1, these Tmax values are equivalent to maximum burial depths of ~7-8 km. The present burial depths of the Datangpo Formation in the study area are ~700 m as a consequence of a long history of uplift and erosion during the Cenozoic (Zhou-Q et al., 2016)

2.2. Basinal restriction and Mn-ore mineralization

In the Neoproterozoic Era, ore-grade sedimentary manganese deposits accumulated on many cratons due to contemporaneous oceanic chemical changes (Roy, 2006; Maynard, 2014; Wu et al., 2016; Yu et al., 2016, 2017; Zhou-Q et al., 2016). The deposits consist mainly of Mn-oxides in Namibia and Brazil and Mn-carbonates (primarily rhodochrosite) in India and South China. In South China, the Mn-ores are deep-water deposits that are correlative with “cap carbonates” found directly overlying Neoproterozoic glacial sediments in shallower settings (Shields et al., 2007; Yu et al., 2017). During the Cryogenian, the Nanhua Basin contained a series of NE-SW-trending grabens and horsts, with the grabens generally being wider than the adjacent horsts (Zhou-Q et al.,

2016). Topographic relief on the horsts was only moderate, and their tops are thought to have been fully submarine throughout the study interval. Thus, the surface layer of the Nanhua Basin watermass was continuous, whereas each graben is likely to have developed an isolated deep watermass. However, the deep grabens were episodically ventilated by hyperpycnal flows during the Sturtian deglaciation (Yu et al., 2016). Despite the presence of numerous horsts and grabens, the Cryogenian Nanhua Basin is thought to have been connected to the global ocean (Wang and Li, 2003; Li-C et al., 2012; Yu et al., 2016). One line of evidence supporting the global-ocean connection is the similar ~5‰ rise in δ13C of post-Sturtian cap carbonate observed in the Nanhua Basin (from ‒4.5 to +3.2‰; Yu et al., 2017) and in coeval sections globally, e.g., Siberian Craton (from ‒2.1 to +3.0‰; Macdonald, 2011; Johnston et al., 2012), Congo Craton (from ca. ‒5 to +3‰; Halverson et al., 2002), and Laurentian Craton (from ‒0.8 to +4‰; Macdonald et al., 2009) (note: slight variations in absolute δ13C values reflect local or regional effects). Although the post-Sturtian sea-level rise led to advection of open-ocean waters into the Nanhua Basin, the deeper grabens remained restricted as shown by geochemical and petrographic data (Wang and Li, 2003; Li-C et al., 2012; Yu et al., 2016), implying semi-restriction for the Nanhua Basin as a whole. These hydrographic conditions, in combination with hydrothermal activity (see Section 5.4), contributed to formation of Mn ore bodies in the lowermost Datangpo Formation (Fig. 2C; Zhou-Q et al., 2016; Du et al., 2015; Yu et al., 2016). Two types of deposits are present: massive manganese ores and banded manganese ores. Within the Songtao-Guzhang Graben, a prominent volcanic tuff is present near the middle of the Mn-carbonate unit, providing an

internal correlation horizon (Zhou et al., 2004; Wu et al., 2016). This horizon shows that the massive ores are present only in the lower half of the Mn-carbonate unit, and that they transition from the graben centers (where they are up to ~26 m thick) into thinner banded ores on the graben margins (Fig. S1). The upper half of the Mn-carbonate unit contains banded manganese ore but no massive ore. The massive ores are considered to be closer to the point sources of manganese, i.e., hydrothermal vents in the deep grabens (whose formation was probably related to contemporaneous block faulting), whereas the banded ores formed through alternating metal and siliciclastic sedimentation at greater distances from the point sources. A gradual waning of hydrothermal activity over time is reflected in the more limited Mn-carbonate accumulation in the upper half of the Mn-carbonate unit (Zhou et al., 2013; Du et al., 2015; Yu et al., 2016).

3. Samples and methods

3.1. Study sections and sampling

The present study investigates the Sturtian succession at three sites in eastern Guizhou Province, South China (Fig. 1). Sites 1a and 1b are two nearby sections (<1 km apart) in underground mines of the Lijiawan Manganese Deposit: PM001 (28°04’16” N, 108°47’32” E) and PM002 (28°04’38” N, 108°47’22” E). Site 2 is drillcore ZK2115 (28°07’29” N, 108°51’33” E) from the Gaodi Manganese Deposit. Site 3 is drillcore ZK4207 (28°02’12” N, 109°05’07” E) from the Xixibao Manganese Deposit. From a structural perspective, Sites 1 and 2 are located within the Lijiawan-Daotuo Graben, and Site 3 is within the Xixibao

Graben. The grabens are characterized by much thicker deposits that mostly accumulated under anoxic conditions, compared to the thin, oxic facies of laterally correlative horst deposits (Fig. S1; Zhou et al., 2013; Yu et al., 2016, 2017). Sites 1a-1b are close to the graben margin and therefore preserve relatively thin intervals of the Mn-carbonate unit (1.2 m at 1a and 4.6 m at 1b). In contrast, Sites 2 and 3 are closer to the centers of their respective grabens and preserve thicker sections of the Mn-carbonate unit (11 m at Site 2 and 13 m at Site 3). Although located in different grabens, Sites 2 and 3 show similar thicknesses for the full 1st Member of the Datangpo Formation (41 and 40 m, respectively; Yu et al., 2016) (Fig. S2). Although absolute water depths cannot be determined, Site 1 was distinctly shallower, accumulating in the surface layer of the Nanhua Basin watermass, compared to Sites 2 and 3, which accumulated below the basinal chemocline. A total of 103 samples were collected and analyzed from the three study sites. At Site 1, twelve samples (4 from 1a, 8 from 1b) were collected from the Mn-carbonate unit; at Site 2, a total of 62 samples were collected, including 24 samples of diamictite from the uppermost ~10 m of the Tiesi’ao Formation, 24 samples from the 11-m thick Mn-carbonate unit and 14 samples from the ~30-m-thick black shale unit of the 1st Member of the Datangpo Formation; and at Site 3, 29 samples were collected from the Mn-carbonate unit. Each sample was 30 to 60 g in size. Visible veins were trimmed off, and samples were washed in deionized water and dried prior to grinding to ~200 mesh for geochemical analyses.

3.2. Analytical methods

In preparation for pyrite sulfur isotopic analysis, which was modified after Canfield et al. (1986), ~1 g of powder was reacted with 40 mL of 1 mol/L CrCl2 and 20 mL of concentrated hydrochloric acid in an N2 atmosphere at 200 °C for 2 h. Dissolved pyrite sulfur was reacted with a 3 % AgNO3-10 % NH3·H2O solution, resulting in Ag2S precipitation. The Ag2S precipitates were then collected, filtered and washed with deionized water, dried in an oven at 50 °C for 48 h, and then mixed with an excess of V2O5. The extraction method for trace sulfate from carbonate was modified after the procedure of Burdett et al. (1989), Hurtgen et al. (2002), and Riccardi et al. (2006). The powdered samples (~50 g) were soaked in 10 % NaCl solution for 24 h and then rinsed in ultrapure water for another 24 h, a procedure that was repeated 5 to 6 times to eliminate soluble sulfate. The cleaned and rinsed powders were reacted with 6 M HCl solution to dissolve carbonate, so that the lattice-bound carbonate-associated sulfate (CAS) was liberated to solution as SO42-. After carbonate was completely dissolved, the solutions were filtered using 1 μm filter paper to remove insoluble residues. Approximately 150 mL of saturated BaCl2 solution (~250 g/L) was added to the filtrate to quantitatively precipitate sulfate as BaSO4, and after 24 h the precipitated BaSO4 was collected and dried. Sulfur isotope compositions were analyzed in the State Key Laboratory of Geobiology and Environmental Geology, the China University of Geosciences-Wuhan using a Thermo Fisher Scientific Delta V Plus isotope ratio mass spectrometer coupled with a Flash elemental analyzer. The results are expressed in standard delta notation as per mille (‰) deviations relative to the international Vienna Canyon Diablo Troilite (VCDT) standard: δ34S =

(Rsample/Rstandard ‒ 1) × 1000, where R = 34S/32S. The analytical error for pyrite δ34S was ~0.1‰ (1σ), calculated from replicate analyses of IAEA international standards (δ34S): IAEA S1 (‒0.3‰), IAEA S2 (+22.65‰), and IAEA S3 (‒32.5‰) (Li et al., 2010; Zhou-CM et al., 2016). The analytical error for CAS δ34S was better than ±0.2‰ (1σ) calculated from replicate analyses of IAEA standards (δ34S): NBS-127 (+21.1‰), IAEA-SO-5 (+0.49‰), and IAEA-SO-6 (‒34.05‰). Total organic carbon (TOC) and total sulfur (TS) measurements and scanning electron microscopy (SEM) were conducted at the State Key Laboratory of Biogeology and Environmental Geology, China University of Geosciences (Wuhan), China. TOC and TS were measured using a Jena multi-EA 4000 carbon sulfur analyzer. Analytical errors are better than ±0.2% (1σ) according to replicate analyses of Alpha Resources standard AR4007.

4. Results

4.1. Petrographic results

The Tiesi’ao Formation consists of glacial till with a wide range of clast sizes, thus exhibiting poor sorting (Fig. 3A). The Mn-carbonate unit of Sites 1 and 3 consists of banded manganese ores, whereas that of Site 2 consists of massive manganese ores. Sites 1 and 2 are located in the Lijiawan-Daotuo Deposit, in which the Mn-carbonate unit recorded massive and banded ores, while the Site 3 is located in the Xixibao Deposit, in which the Mn-carbonate unit only consists of banded ores. The banded manganese ores in the Mn-carbonate unit of the basal Datangpo Fm. consist of Mn-carbonate laminae that are dark

black and 0.1-3 mm in thickness, alternating with clay laminae that are light gray and ~0.3 mm in thickness (Fig. 3B, C). The massive manganese ores are similar but contain no clay laminae (Fig. 3D). Some massive ores close to deposit centers contain bitumen balls that are coated by chalcedony and range in diameter from 0.1 to 2 cm (Fig. 3E). Also present are small (<6 μm), dark brown to black, amorphous particles of organic matter around the margins of Mn-carbonate crystals. The Mn-carbonate unit contains a discrete, bedding-parallel, 4-6 cm-thick pyrite layer (Fig. 3F, G), and Tiesi’ao Formation diamictites contain randomly distributed pyrite grains (Fig. 3H). The pyrite in the thick layer consists of framboidal and euhedral-subhedral grains that are surrounded by clay minerals (Fig. 3I-J; Fig. S3). Scattered pyrite framboids in other layers of the Mn-carbonate and black shale units are all stratiform (i.e., bedding-parallel) in their distribution, consistent with a sedimentary origin (Fig. 3K-L).

4.2. Geochemical results

All geochemical data generated for the study samples are given in Table 1. Carbonate-associated sulfate (CAS) δ34S values in the Mn-carbonate unit of the basal Datangpo Formation are significantly elevated at all study sites, with means of +49.3‰ (Site 1a), +49.3‰ (Site 1b), +62.4‰ (Site 2), and +57.9‰ (Site 3) (Figs. 4-6). The minimum δ34SCAS (+38.8‰) is found at Site 1b and the maximum (+69.8‰) at Site 2. Pyrite δ34S values are also significantly elevated, with means of +47.0‰ (Site 1a), +43.2‰ (Site 1b), +61.6‰ (Site 2), and +59.8‰ (Site 3) (Figs. 4-6). The minimum δ34Spy (+37.8‰) is found at Site 1b

and the maximum (+72.4‰) at Site 2. δ34SCAS and δ34Spy also exhibit a strong positive correlation (r = +0.67, p(α) < 0.001) (Fig. 7A). As a consequence of the similarly heavy δ34SCAS and δ34Spy values, △34S (i.e., δ34SCAS – δ34Spy) is small at all study sites, with means of +2.3‰ (Site 1a), +6.1‰ (Site 1b), +7.1‰ (Site 2), and ‒1.8‰ (Site 3) (Figs. 4-6). The minimum △34S is found at Site 3 (‒11.6‰) and the maximum at Site 2 (+13.5‰). Negative △34S values indicate that pyrite is isotopically heavier than co-existing CAS in a sample (as observed in 20 out of 45 samples), and the negative mean △34S of Site 3 means that pyrite in aggregate is isotopically heavier than co-existing CAS, which is highly unusual. Isotopically heavy pyrite is also present at Site 2 in the underlying Tiesi’ao Formation (mean +48.4‰, maximum +60.6‰) and the overlying black shale unit (mean +52.6‰, maximum +67.2‰; Fig. 5), so this is a general feature of the Cryogenian Nanhua Basin, not a feature unique to the Mn-carbonate unit of the Datangpo Formation. Total sulfur (TS) contents are moderately high: mean TS for the Mn-carbonate unit is 2.2 % at Site 1a, 1.4 % at Site 1b, 2.2% at Site 2, and 2.1 % at Site 3, with a maximum of 4.8% at Site 2 (Figs. 4-6). At Site 2, substantial TS is present in both the underlying Tiesi’ao Formation (mean 2.4%, maximum 4.6%) and the overlying black shale unit (mean 2.8%, maximum 3.7%; Fig. 5). This is an important observation, indicating once again that this is a general feature of the Cryogenian Nanhua Basin, not a feature unique to the Mn-carbonate unit of the Datangpo Formation. Total organic carbon (TOC) contents are modest: mean TOC for the Mn-carbonate unit is 0.9% at Site 1a, 1.6% at Site 1b, 2.2% at Site 2, and 2.1% at Site 3, with a maximum of 3.5%

at Site 3 (Figs. 4-6). At Site 2, TOC is similar in the overlying black shale unit (mean 1.8%, maximum 3.1%) but much lower in the underlying Tiesi’ao Formation (mean 0.2%, maximum 2.7%; Figs. 5, 7B). S/TOC ratios are mostly >1, which is unusual in that typical open-marine sediments have S/TOC ratios averaging ~0.3-0.4 (Berner and Raiswell, 1983), and S/TOC ratios are exceptionally high (>10; owing to low TOC) in the Tiesi’ao Formation. This is a significant observation, showing that sulfur and organic carbon accumulation was not tightly coupled in the Cryogenian Nanhua Basin. Total iron (FeT) content in the Mn-carbonate unit has mean values of 3.4% at Site 1a, 2.6% at Site 1b, 2.6% at Site 2, and 2.7% at Site 3 (Figs. 4-6). At Site 2, the underlying Tiesi’ao Formation has a mean of 3.1% and the overlying black shale unit a mean of 3.7% (Fig. 5). Fe in the study units is present mainly as pyrite, as shown by a strong positive correlation of Fe with TS (r = +0.81, p(α) < 0.001) (Fig. 7C). Some of the study samples (~40 of 103 total) show evidence of hydrothermal influence based on an Al/(Al+Fe+Mn) versus Fe/Ti discriminant plot (Fig. 8).

4.3. Primary character of sulfur isotope signals

CAS represents trace amounts of sulfate that substitute for CO32− in carbonate mineral lattices (Pingitore et al., 1995). The δ34S values of CAS have been widely used as a proxy for seawater sulfate δ34S, especially in Precambrian successions where evaporitic sediments are rare (Kampschulte et al., 2001; Hurtgen et al., 2002, 2004, 2006; Kah et al., 2004; Kampschulte and Strauss, 2004; Gellatly and Lyons, 2005; Bottrell and Newton, 2006). In

general, there are only small differences (<5‰) between the S isotopic compositions of CAS and coeval marine evaporites, showing that CAS takes up seawater sulfate without isotopic fractionation (Burdett et al., 1989; Kah et al., 2004; Kampschulte and Strauss, 2004; Lyons et al., 2004). Several arguments suggest that the δ34SCAS values of the present study samples are well-preserved. In the Sturtian Nanhua Basin, the deepwater Mn-carbonate was early diagenetic in origin, forming at or just below the sediment-water interface (Yu et al., 2016) and, thus, likely to have recorded the S isotopic composition of sulfate in the overlying watermass. δ34SCAS can potentially be contaminated by non-CAS sulfate, through oxidation of pyrite during prolonged exposure in outcrop or sample preparation. However, our samples are from deep drillcores and were thus never exposed to near-surface weathering, and their δ34SCAS values show no relationship to either CAS or pyrite sulfur concentrations (Fig. S4A), indicating no measurable influence by pyrite oxidation (cf. Neoproterozoic Otavi Group, Namibia, Fig. S4B). The NaCl solution used to eliminate soluble sulfate can potentially mobilize non-CAS sulfate (e.g., gypsum S) during the extraction procedure (Theiling and Coleman, 2015). However, there was no gypsum in our samples, and pyrite oxidation during CAS extraction was negligible owing to use of a sealed evacuated reaction vessel during the extraction process.

5. Discussion

In the following, we consider key findings of our study in relation to specific issues,

including sulfate concentrations in Sturtian global seawater (Section 5.1), water-depth gradients of sulfur isotopes in the Nanhua Basin (Section 5.2), high sulfur concentrations in Nanhua Basin sediments (Section 5.3), and evidence for hydrothermal activity in the Sturtian Nanhua Basin (Section 5.4). These topics provide background and context to the new sulfur-cycle model for the Sturtian Nanhua Basin that we present (Section 5.5), and which we compare with existing Sturtian sulfur-cycle models (Section 5.6).

5.1. Low sulfate concentrations in Sturtian seawater The Proterozoic ocean is thought to have been characterized by generally low sulfate concentrations (Bottomley et al., 1992; Hayes et al., 1992; Canfield, 1998; Hurtgen et al., 2002, 2004, 2005; Shen et al., 2002, 2003; Lyons et al., 2004). Mesoproterozoic and Neoproterozoic marine units have yielded estimated concentrations mostly in the range of 0.5-5 mM based on both isotopic data and the low CAS content of marine limestones (<300 ppm) (Gorjan et al., 2000, 2003; Shen et al., 2002; Hurtgen et al., 2002, 2004, 2005) as well as rates of change in δ34SCAS values (Kah et al., 2004; Algeo et al., 2015). Especially low sulfate concentrations were inferred in the immediate aftermath of the Sturtian and Marinoan glaciations, when CAS contents fell to lows of 38 ppm and 25 ppm, respectively (Hurtgen et al., 2002). Low Cryogenian sulfate concentrations have been linked to a diminished global hydrologic cycle during Snowball Earth events, limited oxidative weathering of pyrite in subaerial environments, and reduced riverine fluxes of sulfate to the ocean (Kirschvink, 1992; Hoffman et al., 1998; Hoffman and Schrag, 2002). The buildup of ferrous iron and BIFs

during the Sturtian Glaciation are evidence that Fe supply substantially exceeded sulfide availability (Hurtgen et al., 2002). Strongly elevated δ34SCAS and δ34Spy values are generally regarded as evidence of substantial seawater sulfate drawdown through precipitation of 32S-enriched pyrite, causing the residual sulfate pool to become isotopically heavy (Bottrell and Newton, 2006; Algeo et al., 2015). Sturtian seawater sulfate is thought to have been strongly 34S-enriched, as inferred from high δ34SCAS values in all Sturtian marine units globally, e.g., the Tapley Hill Formation in Australia (mean +21.7‰, +10.1 to +27.0‰; Gorjan et al., 2000; Hurtgen et al., 2005) and the Rasthof Formation in Namibia (mean +32.9‰, +12.4 to +51.5‰; Hurtgen et al., 2002, 2005) (Table 2). However, the present study units yield δ34SCAS values (mean +56.0‰, +38.8 to +69.8‰; Table 1) that are ~25 to 35‰ higher than in coeval sections globally. A similar global pattern is seen for δ34Spy: values are moderately elevated for Sturtian formations in Australia (mean +28.7‰ to +31.1‰; Gorjan et al., 2000), Namibia (mean +37.6‰; Gorjan et al., 2003), Canada (mean +31.9‰; Sperling et al., 2016), Greenland (mean +37.2‰; Scheller et al., 2018), and the UK (mean +44.8‰; Parnell and Boyce, 2017). However, the present study units yield δ34Spy values (mean +54.3‰, range +23.8 to +72.4‰ (see also Liu et al., 2006; Chen et al., 2008; Wu et al., 2016; Wang et al., 2016) that are ~10 to 25‰ higher than in coeval sections globally. These observations suggest that the Sturtian global ocean contained somewhat 34S-enriched seawater sulfate, but that Nanhua Basin seawater was even more 34S-enriched. Low seawater sulfate concentrations may be supported also by small differences in sulfur isotope compositions between co-occurring sulfate and pyrite (△34S, or δ34SCAS – δ34Spy).

Small △34S values can be an indication of limited fractionation between cogenetic sulfate and sulfide deposits, which scales logarithmically with seawater sulfate concentrations (Habicht et al., 2002; Algeo et al., 2015). Globally, Sturtian units show low △34S values, e.g., the Tapley Hill Formation in Australia (mean +0.1‰, range 0 to +1.7‰; Gorjan et al., 2000; Hurtgen et al., 2005) and the Rasthof Formation in Namibia (mean +7.5‰, range 0 to +41.4‰; Gorjan et al., 2000, 2003; Hurtgen et al., 2005). In the present study units, △34S is also low (mean +0.7‰, range –6.5 to +8.0‰; n = 45). Per the ‘MSR trend’ of Algeo et al. (2015), △34S values of <10‰ and <3‰ are nominally equivalent to seawater sulfate concentrations of <~1 mM and <~0.1 mM, respectively. The existence of many negative △34S values in the present study units (20 out of 45 samples) suggests a more complex control on S isotopes, because negative △34S values cannot be explained by simple MSR fractionation (see S1 of Supplemental Information). Such negative values, where pyrite δ34S exceeds co-occurring sulfate δ34S, are known as “superheavy” pyrite; Liu et al., 2006) and have been identified not only in Sturtian-age deposits (Gorjan et al., 2000; Hurtgen et al., 2002; Liu et al., 2006; Cui et al., 2018; this study) but also in units of Marinoan (Shen et al., 2008), Ediacaran (Ries et al., 2009) and Early Paleozoic age (Kah et al., 2016; Drake et al., 2018). The conventional explanation, based on studies of the Sturtian, links negative △34S values to Snowball Earth conditions, in which a stratified oceanic water column had a sulfidic deep layer that was 34S-enriched through Rayleigh distillation, and mixing of 34S-enriched sulfide into the surface layer at the end of the glaciation produced the widely observed negative △34S values (Gorjan et al., 2000; Hurtgen et al., 2002). However, such occurrences are not limited to deglaciations, and a more

general explanation for their occurrence is that the Neoproterozoic to Early Paleozoic witnessed a large-scale shift in the marine sulfur cycle characterized by increasingly diverse sulfur reactions contributing to greater production of redox intermediates and disproportionation of sulfur isotopes (Fike et al., 2015).

5.2. Water-depth gradients of S-isotopic compositions in Nanhua Basin Another finding of the present study that is significant for understanding the sulfur cycle of the Sturtian Nanhua Basin is the existence of water-depth gradients in both CAS and pyrite δ34S values in the Mn-carbonate unit. Based on relative water depths (see Section 3.1), δ34SCAS values increase from the shallower Site 1 (mean +49.3‰) to the deeper Site 2 (mean +62.4‰) and Site 3 (mean +57.9‰) (Fig. 9A). Pyrite δ34S values in the Mn-carbonate unit show the same trend, increasing from Site 1 (mean +44.5‰) to Site 2 (mean +61.6‰) and Site 3 (mean +59.8‰) (Fig. 9B). Based on a Student’s t test, the shallow-water (Site 1) and deep-water sections (Sites 2 and 3) have significantly different mean CAS and pyrite δ34S values, although the differences between the two deep-water sections are not significant (see S2 of Supplemental Information). Data from published sources show that the overlying black shale unit exhibits the same depth gradient (Fig. 9C), increasing from the shallower Yangjiaping area (mean +17.7‰, range +6.0 to +29.5‰, n = 6) to the deeper Minle (mean +51.4‰, range +42.6 to +64‰, n = 29) and Gaodi areas (mean +52.6‰, range +31.1 to +65.8‰, n = 14) (Fig. S5; Feng et al., 2010; Li-C et al., 2012; this study). We interpret all of these trends to record a vertical gradient in seawater sulfate δ34S, with the deep watermass of

the Sturtian Nanhua Basin having been more 34S-enriched than the shallow watermass. The existence of water-depth gradients in sulfate δ34S within the Nanhua Basin is consistent with low aqueous sulfate concentrations, as inferred from δ34SCAS, δ34Spy, and △34S data. Low sulfate concentrations would have been essential for generation of vertical variation in seawater sulfate δ34S, and the existence of such gradients effectively precludes high watermass sulfate concentrations. This further highlights the unusual character of the high sediment TS concentrations in the Sturtian strata of the Nanhua Basin (see Section 5.3), for which some other explanation is needed. Thus, multiple lines of evidence―including water-depth gradients in sulfate δ34S within the Nanhua Basin, highly 34S-enriched compositions of CAS and pyrite, and the near-zero △34S values of the study units―nominally support an inference of very low seawater sulfate concentrations in the Sturtian Nanhua Basin. However, other observations suggest a more complex scenario, as described below.

5.3. High sulfur concentrations in Sturtian strata of South China The Sturtian Nanhua Basin contains remarkably high total sulfur (TS) concentrations: 0.8-4.0% (mean 2.4%) in the Tiesi’ao Formation, 0.8-3.2% (mean 2.1%) in the Mn-carbonate unit of the basal Datangpo Formation, and 2.1-3.4% (mean 2.8%) in the overlying black shale. These concentrations are similar to those found in many modern marine sediments (e.g., Berner and Raiswell, 1984; Chanton et al., 1987) that have accumulated from seawater with sulfate concentrations of ~29 mM (Yao and Millero, 1996), and they are far higher than

in modern freshwater sediments (typically <1 %), in which aqueous sulfate concentrations are mostly <500 µM (Holmer and Storkholm, 2001). Thus, given estimates of Neoproterozoic seawater sulfate concentrations of <1 mM (Harrison and Thode, 1958; Canfield, 2001; Habicht et al., 2002; Canfield et al., 2010; Algeo et al., 2015), the TS concentrations of the study units are remarkably high for deposits accumulating in putatively low-sulfate seawater. All of the study units exhibit S/TOC ratios of >1, which significantly exceed those of typical open-marine facies (~0.3-0.4; Berner and Raiswell, 1983). Although S/TOC ratios of ancient marine sediments sometimes increase owing to loss of organic carbon in the burial environment (Raiswell and Berner, 1987), the S/TOC ratios of the Tiesi’ao Formation diamictites are >10 owing to their near-zero TOC contents (mean 0.2%). Furthermore, there is no relationship of TOC and TS (Fig. 7B). All of these observations suggest that accumulation of organic carbon and sulfur was largely uncoupled in the study units. A higher original content of organic carbon in the Tiesi’ao Formation also seems unlikely because of generally low marine productivity during Snowball Earth events (Kirschvink, 1992; Hoffman et al., 1998; Hoffman and Schrag, 2002). This considerations imply that other (i.e., non-MSR) sources of pyrite sulfur were important within the Cryogenian Nanhua Basin. It is difficult to determine whether the high sulfur contents and S/TOC ratios of the Nanhua Basin are typical of the Sturtian glacial interval globally because relatively little TS data are available for non-Chinese Sturtian units. Similar TS concentrations are present in the ~660-Ma basal Arena Formation, East Greenland (mean 2.5%, range 1.9% to 2.9%) (Scheller, et al., 2018), but lower concentrations were reported from the ~717-Ma MacDonaldryggen Member of the Elbobreen Formation, Svalbard (mean 0.6%; 0 to 1.0%)

(Kunzmann et al., 2015) and the Sturtian-equivalent Port Askaig Formation, UK (mean 0.36%; min-max 0-1.9%) (Parnell and Boyce, 2017). Most non-Sturtian Neoproterozoic marine units exhibit quite low total sulfur concentrations. TS contents are 0-0.2% (mean 0.1%) in the pre-Sturtian Russøya Member of the Elbobreen Formation in Svalbard (Kunzmann et al., 2015), 0-1.2% (mean 0.5%) in the post-Marinoan Dracoisen Formation of Svalbard (Kunzmann et al., 2015), and 0-0.9% (mean 0.3%) in the post-Marinoan Doushantuo Fm. of South China (Cui et al., 2015). However, until more sulfur data are reported for Sturtian-age units, the full significance of the high concentrations in the present study units cannot be determined.

5.4. Evidence for hydrothermal influences in the Sturtian Nanhua Basin Earlier studies of Sturtian strata in the Nanhua Basin inferred hydrothermal activity based on various lines of evidence, including positive Eu anomalies, high La/Sc and i(87Sr/86Sr) ratios, and low εNd(t) values (Yu et al., 2016; Wu et al., 2016). Discriminant plots have been widely used to distinguish hydrothermal versus hydrogenous sources of metals to marine sediments (Bonatti, 1975; Toth, 1980; Marchig et al., 1982; Boström, 1983; Peter and Goodfellow, 1996). A Fe/Ti versus Al/(Al+Fe+Mn) crossplot suggests locally strong hydrothermal inputs to the study units, especially in the more Mn-rich samples (>10% Mn) from the Mn-carbonate unit (Fig. 8), an inference consistent with several other types of discriminant plots used by Yu et al. (2016). Evidence of hydrothermalism in the Nanhua Basin spans an extended time interval, including the Sturtian (Yu et al., 2016) and Marinoan

glacial intervals (Huang et al., 2013), the early Cambrian (Coveney and Chen, 1991; Grauch et al., 1991; Steiner et al., 2001; Liu et al., 2015), and the Late Devonian (Chen et al., 2006). Although some studies have argued against hydrothermal activity in the Nanhua Basin (Tang and Liu, 1999; Liu et al., 2006), the large number of studies finding such evidence is difficult to dismiss. Hydrothermal activity can have a strong influence on watermass chemistry and sediment composition in restricted and semi-restricted basins (Wang and Li, 2003; Li-C et al., 2012; Yu et al., 2016). For example, the modern Red Sea has multiple sub-basins with stratified water columns, similar to the Sturtian Nanhua Basin. In these sub-basins, the water column consists of an oxic normal-salinity surface water, an intermediate layer exhibiting sharp redox and salinity gradients in which Fe-Mn-oxyhydroxides precipitate, and an anoxic, metal-rich, high-salinity, high-temperature deep layer (Cocherie et al., 1994; Schmidt et al., 2003; Butuzova et al., 2009). The briny deep layer of these sub-basins exhibits strong hydrothermal influence, as shown by 87Sr/86Sr and εNd(t) values (Cocherie et al., 1994; Anschutz et al., 1995). A second example is the Guaymas Basin of the Gulf of California, another deep, restricted basin with hydrothermal inputs (Byrne and Emery, 1960; Lawver et al., 1975). Hydrothermal fluids are discharged through chimneys at 270 to 320 °C as well as through porous sediment, setting up vertical chemical gradients in the overlying water column with respect to the concentrations and isotopic compositions of DIC, methane, and sulfate (Lonsdale and Becker, 1985; Jørgensen et al., 1990; Elsgaard et al., 1994; Biddle et al., 2012). Hydrothermal activity is likely to have modified the watermass chemistry of the Sturtian

Nanhua Basin. The Mn-carbonate fraction of the basal unit of the Datangpo Formation shows a substantially different chemical composition from the aluminosilicate fraction, being characterized by much higher initial 87Sr/86Sr ratios and lower initial εNd(t) values (Yu et al., 2016). A weathering origin for the manganese in the Mn-carbonate unit is inconsistent with evidence such as a negative correlation of Mn versus Al2O3, and low Al/Mn ratios (0.09 to 4.2, mean 0.68) relative to upper continental crust (105; Rudnick and Gao, 2003). Furthermore, a hydrothermal source is supported by various lines of evidence (Wu et al., 2016; Yu et al., 2016): (1) a positive correlation of Eu/Eu* with Mn content, as positive Eu anomalies are characteristic of hydrothermal fluids; (2) compositions falling in the hydrothermal field of discriminant plots for marine metalliferous sediments (e.g., Fe/Ti versus Al/(Al+Fe+Mn)), with the manganese ores exhibiting the greatest hydrothermal influence (Yu et al., 2016 and this study); and (3) chondrite-normalized REE distributions for the manganese ores and wall rocks that indicate partial derivation of REEs from hydrothermal fluids. Further evidence of hydrothermal influence on sulfur cycling in the Sturtian Nanhua Basin comes from documentation of pyrite produced through thermochemical sulfate reduction (TSR). TSR is an abiotic process that takes place in the sediment column at temperatures >110 °C (Fig. 10; Goldstein and Aizenshtat, 1994; Machel et al., 1995; Worden et al., 1995; Machel, 2001; Jiang et al., 2015). During TSR, sulfate is reduced concurrently with organic matter oxidation, generating reaction products similar to those of MSR. The TSR process may generate dolomite, typically as milky white medium-to-coarse saddle-form crystals, and secondary porosity, although the amount of porosity produced is typically small

(Machel, 2001). Although not always diagnostic, TSR can produce characteristic petrographic features, e.g., milky white, coarse-crystalline cements and/or replacive masses after anhydrite (Machel, 1989; Machel et al., 1995; Worden et al., 1996). Whereas both MSR and TSR produce framboidal pyrite, TSR can also generate pyrite, galena, and sphalerite with banded or colloform textures (Machel, 1989, 2001; Machel et al., 1995). Sulfur isotope signatures can also help to differentiate TSR and MSR: sulfur isotope fractionation during TSR is strongly temperature-dependent, with fractionations (△34S) of <20‰ that become smaller with rising temperature (Kiyosu, 1980; Kiyosu and Krouse, 1990; Watanabe et al., 2009). In contrast, sulfur isotope fractionation during MSR can reach ~46‰ (Canfield, 2001), although both smaller and larger fractionations are known (Habicht et al., 2002; Algeo et al., 2015; see S1 of Supplemental Information). In the Nanhua Basin, evidence of TSR is found within the Mn-carbonate unit at some basinal locales (Cui et al., 2017a, 2017b, 2018). Petrographic features including large framboids, lacy cements, and rings and “flowers” of pyrite crystals, as well as secondary overgrowths on these features, are reasonably diagnostic of a TSR origin. However, the pyrite in our study units consists exclusively of syngenetic framboids and early diagenetic euhedral crystals (Figs. 3, S3), similar to those in marine units of many ages (e.g., Wilkin et al., 1996; Taylor and Macquaker, 2000; Wignall et al., 2005). We found none of the unusual textures identified by Cui et al. (2017a, 2017b, 2018), despite a careful examination of all study samples in an attempt to locate such features. Hydrothermal alteration is a fundamentally local process, since hot fluids tend to be funneled along relatively narrow vertical conduits (Hannington et al., 2001, 2005) and the high temperatures of hydrothermal fluids (to several

hundred °C) are rapidly attenuated upon mixing with cold seawater (Jannasch and Mottl, 1985). We infer that our study sections accumulated in basinal areas at least somewhat removed from the hydrothermal vent systems encountered by Cui et al. (2017a, 2017b, 2018). The timing of pyrite formation in the Mn-carbonate unit of the Datangpo Fm. is under debate. Cui et al. (2017a, 2017b, 2018) inferred a “late diagenetic” origin on the basis of petrographic relationships showing pyrite growth postdating illite and rhodochrosite accumulation. However, the rhodochrosite in the Datangpo Fm. was a synsedimentary precipitate (Yu et al., 2016) and thus does not provide a significant constraint on the timing of diagenetic pyrite growth. In any case, given the pronounced differences in petrographic character of the pyrite documented by Cui et al. (2017a, 2017b, 2018) and that of the present study, it is possible that multiple generations of pyrite formed in the Datangpo Fm. over an extended interval of time. Hydrothermal vent systems can be long-lasting (i.e., over millions of years; Cathles et al., 1997; Pedersen et al., 2010), allowing for such a possibility. Cui et al. (2018) argued that the high δ34S values of the TSR-related pyrite (ca. +56 to +60‰) also favored a secondary origin through strong Rayleigh distillation of an aqueous sulfate pool in the subsurface environment. Although they may well be correct regarding the process of 34S enrichment, this mechanism could have been in operation both during and following Datangpo Fm. deposition and therefore does not represent a robust age constraint.

5.5. A hydrothermal sulfur-cycle model for the Sturtian Nanhua Basin The fundamental conundrum that must be resolved is how, despite strong evidence for

low seawater sulfate concentrations during the Sturtian (i.e., high δ34SCAS and δ34Spy values, pronounced water-depth gradients in δ34SCAS and δ34Spy, and near-zero △34S; see Sections 5.1-5.2), large amounts of reduced sulfur could have accumulated in the present study units (which have mean TS of 2.2%; see Section 5.3; Figs. 4-6). Existing sulfur-cycle models for the Sturtian Nanhua Basin (Logan et al., 1995; Li et al., 1999; Chu et al., 2001; Chen et al., 2008; Shen et al., 2008; Feng et al., 2010; Li-C et al., 2012; Zhang et al., 2013; Wu et al., 2016; see Section 5.6) are focused mainly on explaining S-isotopic variation in the context of low contemporaneous seawater sulfate concentrations (Section 5.1), but they do not and―in some cases―cannot account for all of the data observations of the present study (Sections 5.2-5.3), nor do many of them take into consideration strong evidence for hydrothermal activity in the Nanhua Basin (Section 5.4). A hydrothermal source resolves this conundrum by providing large quantities of sulfur that were rapidly removed from the water column to the sediment. Here, we present a new model for the sulfur cycle of the Sturtian Nanhua Basin based on a dominant hydrothermal source of sulfur (Fig. 10) that accounts for all data observations of the present study as well as for evidence of TSR in the Datangpo Formation (Cui et al., 2017a, 2017b, 2018). Sulfur emissions from hydrothermal vents are typically in reduced (H2S) rather than oxidized form (sulfate). In situ measurements of high-temperature vent fluids at the Main Endeavour Field, Juan de Fuca Ridge (2200 m water depth) yielded H2S concentrations of 17.3 mmol/kg (Ding et al., 2001), and H2S-bearing fluids have been found in many other hydrothermal vent systems, such as the East Pacific Rise (Von Damm, 2000), Mid-Atlantic Ridge (James and Elderfield, 1996; Zielinski et al., 2011) and North Fiji Basin (Sander and

Koschinsky, 2000). These vents are associated with both sulfide and sulfate deposits, with the former sourced mainly from hydrothermal fluids and the latter from cold deep seawater (Styrt et al., 1981; Ono et al., 2007). Rarely, sulfate-rich hydrothermal fluids have been reported, as from the submarine Desmos Caldera located in the eastern Manus Basin, Papua-New Guinea (Gamo et al., 1997) or the Tjornes Fracture Zone near Iceland (Hannington et al., 2001). In the present study units, we infer that hydrothermally vented sulfur in the Sturtian Nanhua Basin was mainly H2S rather than sulfate, because the Tiesi’ao Formation contains large amounts of pyrite (TS = 2.4 ± 1.3%) despite a lack of organic matter (mean TOC 0.2%) to drive MSR (Fig. 6). Evidence for TSR in graben areas of the Nanhua Basin (Cui et al., 2017a, 2017b, 2018) suggests a link to hydrothermal activity since both processes operate at elevated temperatures (>110 °C). Cui et al. (2018) inferred that TSR activity was driven by an upward flux of sulfate along buried faults in the Nanhua Basin, which is consistent with our inference of hydrothermal venting of H2S provided that the rising sulfate was quantitatively reduced to H2S. Cui et al. (2018) also suggested that the gypsum-bearing Ediacaran Dengying Formation (δ34SCAS: to ca. +40‰; Cui, 2015; Cui et al., 2016) was a possible sulfate source for Nanhua Basin hydrothermal fluids, although this seems improbable in that it would require sulfate to have moved stratigraphically downward in order to affect Sturtian-age strata (note: there is insufficient structural deformation in the study area to make this idea plausible; BGMRGZP 1987; Wang et al., 2013). We cannot identify the specific sulfate source of Nanhua Basin hydrothermal fluids but presume that it came from some older (i.e., pre-Sturtian), more deeply buried evaporitic unit. It should be noted that the highest known evaporitic sulfate

δ34S values are +30 to +35‰ in the Phanerozoic (for uppermost Ediacaran-lower Cambrian evaporites; Claypool and Holser, 1980; Strauss, 1997), with similar values for the less complete Precambrian evaporite record (Cameron, 1982; Strauss, 1993). Given a maximum TSR fractionation of <20‰ (Section 5.4), it would be difficult or impossible to achieve observed δ34SCAS and δ34Spy values up to +60 to +70‰ (Figs. 4-6) through a single-step reduction process. This lends support to the idea that a large subsurface sulfur pool was repeatedly fractionated, e.g., through cyclic reduction-oxidation reactions, leading to a strongly 34S-enriched hydrothermal sulfide flux (cf. Cui et al., 2018). In our model, hydrothermal vents in the deep grabens of the Nanhua Basin released massive quantities of 34S-enriched H2S into the overlying water column (Fig. 10). The hydrothermal H2S flux led to direct formation of highly 34S-enriched framboidal pyrite through reaction with Fe2+ in the anoxic deep watermass. At shallower depths, vented H2S was oxidized to sulfate and mixed with less 34S-enriched sulfate from the global ocean, resulting in lower δ34S values for both CAS and pyrite generated through MSR. Vertical mixing of two distinct sulfur sources is evidenced by water-depth gradients in δ34SCAS and δ34Spy in the Mn-carbonate unit (Fig. 9; see Section 5.2). The hydrothermally sourced H2S had high δ34S values (ca. +60‰) based on δ34Spy values from our deep-graben sections (Fig. 9) and TSR-related pyrite (Cui et al., 2018). In contrast, seawater sulfate in the surface layer of the Nanhua Basin, which came from the open ocean, probably had a δ34S value close to +30‰, as inferred from δ34SCAS values of Sturtian deposits of other cratons globally (e.g., Gorjan et al., 2000; Hurtgen et al., 2002, 2005). A contributing factor to the development of a vertical gradient in aqueous sulfate δ34S within the Nanhua Basin was strong water-column

stratification (Yu et al., 2016; see Section 5.6) combined with low dissolved sulfate concentrations in the global ocean (see Section 5.1). Although hydrothermal venting released large amounts of sulfur to the Nanhua Basin water column (as reflected in mean sediment TS of 2.2%), a large respiratory demand for sulfate owing to high organic carbon sinking fluxes led to most aqueous sulfate being rapidly re-reduced to H2S via MSR and removed to the sediment as pyrite (Fig. 10). MSR would have operated both in the water column and below the sediment-water interface of the Nanhua Basin. Keeping in mind that the near-zero △34S values at all sites (Figs. 4-6) are evidence of no significant fractionation between seawater sulfate and H2S at any water depth in the Nanhua Basin, if pyrite at the shallow sites had formed directly from hydrothermally vented H2S, then it would have recorded δ34Spy values as heavy as the deep-water sites. However, because pyrite at the shallow sites exhibits much more 34S-depleted compositions, it must have formed through microbial reduction of aqueous sulfate (MSR), since only sulfate (and not H2S) would have developed a water-depth gradient through mixing. The observation that concentrations of total sulfur (which consists mainly of pyrite sulfur; Section 4.2) are as high at shallow sites (mean 2.7%) as at deep sites (mean 2.2%) indicates that large quantities of hydrothermally sourced H2S were oxidized to sulfate, mixed with 34S-depleted surface-layer sulfate, and then re-reduced to H2S by MSR before deposition as pyrite. These considerations suggest that, although hydrothermal venting was a major source of sulfur to the Sturtian Nanhua Basin (see Section 5.4), this sulfur was largely reprocessed by MSR in the water column prior to final accumulation. Quantitative reduction and removal of hydrothermally sourced sulfur as pyrite was

responsible for the similar δ34S compositions of pyrite and CAS (i.e., near-zero △34S values) throughout the Datangpo Formation (Figs. 4-6). The fact that △34S values are not uniformly near-zero but vary modestly (from –11.6‰ to +13.5‰) and are frequently negative (20 out of 45 analyses being < 0‰; see Section 5.1) was due to some combination of secular or spatial fluctuations in the δ34S of hydrothermally emitted H2S, the timing of pyrite versus carbonate precipitation, or the mixing ratio of surface (global-ocean-derived) sulfate with hydrothermal H2S. As hydrothermal influence waned, sulfate inputs from the open ocean and continental weathering became proportionately more important, resulting in declining pyrite δ34S values: the black shale unit at Site 2 shows an upward trend culminating in δ34Spy values of +30‰ at its top (Fig. 5), i.e., similar to the composition of global seawater sulfate during the post-Sturtian interval (Gorjan et al., 2000; Hurtgen et al., 2002, 2005).

5.6. Comparison with existing sulfur-cycle models for Sturtian Nanhua Basin Our model for the Sturtian Nanhua Basin sulfur cycle (Section 5.5) accounts for all observational data of the present as well as earlier studies. Here, we compare it to earlier published models for the Nanhua Basin and other Sturtian-age depositional systems, including (1) the Snowball Earth model (Chu et al., 2001), (2) the basin restriction model (Li et al., 1999; Chen et al., 2008; Feng et al., 2010; Li-C et al., 2012; Wu et al., 2016), (3) the basin stratification model (Logan et al., 1995; Li et al., 1999; Shen et al., 2008; Li-C et al., 2012), and (4) the Mn reduction model (Zhang et al., 2013). The “Snowball Earth model” emphasizes the role of rapidly rising temperatures and nutrient fluxes following the

termination of the Sturtian Glaciation, leading to near-quantitative utilization of watermass sulfate and limited fractionation between CAS and pyrite δ34S values (Chu et al., 2001). The “basin restriction model” is based on formation of intercontinental seas or isolated basins following the breakup of the supercontinent Rodinia, within which seawater sulfate evolved to higher δ34S and lower concentrations than the coeval open ocean owing to lack of sulfate inputs and efficient sulfide removal (Li et al., 1999; Chen et al., 2008; Feng et al., 2010; Li-C et al., 2012; Wu et al., 2016). The “basin stratification model” is a variant of the basin restriction model that is intended to account for observations of δ34Spy > δ34SCAS; it is similar in invoking strong MSR in an anoxic deep watermass to generate 34S-enriched residual sulfate and H2S, but it assumes that the CAS was derived from an oxic ocean-surface layer with somewhat lower (i.e., less evolved) sulfate S-isotopic compositions (Logan et al., 1995; Li et al., 1999; Shen et al., 2008; Li-C et al., 2012). The “Mn reduction model” invokes a reaction between FeS (the precursor for pyrite formation) and MnO2 (i.e., FeS + 4.5MnO2 + 4H2O  FeOOH + 4.5Mn2+ + SO42- + 7OH-), generated by the reaction of H2S and reactive Fe2+, to account for high pyrite δ34S values (Zhang et al., 2013). It should be noted that these models are not necessarily all mutually exclusive. A common feature of all models to date is low Neoproterozoic seawater sulfate concentrations (cf. Section 5.1). These models mainly attempt to account for high pyrite δ34S values, which have been widely reported for Sturtian sediments (Gorjan et al., 2000, 2003; Hurtgen et al., 2002, 2005; Sperling et al., 2016; Parnell and Boyce, 2017). They invoke 34S-enrichment

of pyrite through strong MSR under conditions of low seawater sulfate

concentrations, which would have resulted in small sulfur isotope fractionations and led to

evolution of seawater sulfate δ34S toward heavier compositions (note that δ34SCAS values average ca. +30‰ globally for the Sturtian; Gorjan et al., 2000; Hurtgen et al., 2002, 2005). Only the basin stratification model includes a mechanism for generating negative △34S values (see Section 5.1), which are impossible to explain based on simple MSR fractionation, and the existing models do not attempt to account for observations of high TS concentrations in Sturtian strata (Section 5.3). The existing models are, for the most part, not incorrect but, rather, incomplete. Our model (Section 5.5) accounts for all observations of the present as well as earlier studies of Sturtian Nanhua Basin strata, and it also has general implications for understanding features of Cryogenian global paleoceanography.

6. Conclusions

The Neoproterozoic Nanhua Basin contains a continuous Sturtian-age succession of glacial to post-glacial sediments. These units show an unusual combination of features, including (1) high δ34S values for CAS (mean +56.0‰, range +49.6 to +62.6‰) and pyrite (mean +54.8‰, range +45.0 to +65.4‰), (2) small △34S values (i.e., δ34SCAS – δ34Spy) (mean +0.7‰, range–11.6‰ to +13.5‰) with a large proportion of negative values (20 out of 45), (3) water-depth gradients in both CAS and pyrite δ34S values, with higher mean values in shallower-water sections, and (4) high total sulfur content (2.2 ± 1.1%). The small △34S values and existence of water-depth gradients in δ34S provide strong evidence of low aqueous sulfate concentrations in the Sturtian Nanhua Basin, consistent with similar inferences for Neoproterozoic basins globally, but this condition is difficult to reconcile with the unusually

high TS content of the study units. Supported by independent evidence of hydrothermal activity in the Sturtian Nanhua Basin, we hypothesize that the high TS content and other observed features were due to large-scale hydrothermal fluxes of H2S. This model accounts for the similar ranges of δ34S values of pyrite and CAS (the latter being due to oxidation of hydrothermal H2S), the negative △34S values (due to random secular or spatial variation in the δ34S of the hydrothermal H2S flux), and the high TS content of the study units despite low aqueous sulfate concentrations (a condition maintained by rapid removal of hydrothermal H2S via reaction with aqueous Fe2+ in the anoxic deep water column of the Nanhua Basin). Our model provides new insights into sulfur cycling processes in a semi-restricted marine basin that are likely to have wider applicability to Neoproterozoic marine systems.

Acknowledgments

We thank the Guochun Zhao and an anonymous reviewer for their constructive comments. We are grateful to the officers of State Key Laboratory of Geobiology and Environmental Geology in CUG for the lab help. We also thank Hua Guo, a lecturer in CUG, for her suggestions on this manuscript. This research is supported by NSFC (Natural Science Foundation of China)-Guizhou Karst Study Center, and Ministry of Land and Resources of the People’s Republic of China “Deep Prospecting and Metallogenic System in the Southeastern Margin of the Upper Yangtze Block (No. 201411051)” Project, and Geological Mineral Exploration and Development Bureau of Guizhou Province “Mineralogical and Geochemical Constraints on the Ore Genesis of the Gas Seepage-type Sedimentary

Manganese Deposit in the Nanhua Rift Basin, eastern Guizhou Province (No. 2016 [30])” Project.

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Table captions

Table 1. Geochemical data for the three study sites (Lijiawan, Xixibao, and Gaodi areas).

Table 2. Compilation of published δ34S data from Cryogenian strata. Note that ranges of pyrite and CAS sulfur-isotopic compositions are reported as minima and maxima.

Figure captions

Fig. 1. (A) Major tectonic subunits of China. (B) Geographic map of eastern Guizhou Province, showing the locations of the study sites.

Fig. 2. (A) Cryogenian paleogeography of South China blocks (Jiang et al., 2003; Zhou et al., 2004). (B) Map of Nanhua Rift Basin on the southeastern margin of the Yangtze Block (Du et al., 2015). (C) Map of Songtao area, eastern Guizhou Province, with study sections marked by yellow stars: PM001 and PM002 in Lijiawan area, drillcore ZK2115 in Gaodi area, and drillcore ZK4207 in Xixibao area (Zhou-Q et al., 2016).

Fig. 3. (A) Diamictite of Tiesi’ao Formation (1640.98 m, drillcore ZK2115); (B-C) Banded manganese ore of Mn-carbonate unit of basal Datangpo Formation (2.29 m, PM002, Site 1); (D-E) Massive ores of Mn-carbonate unit, some containing bitumen balls (D = 1635.40 m

and E = 1636.25 m, drillcore ZK2115); (F-G) Pyrite layers interbedded within manganese layer in Mn-carbonate unit (from ore heap of Gaodi Deposit); (H) Pyrite framboids in Tiesi’ao Fm. (H = 1642.48 m, drillcore ZK2115); (I) Framboids and euhedral to subhedral pyrite grains in Mn-carbonate unit (2.29 m, PM002, Site 1); (J) Pyrite framboids in Mn-carbonate unit of the Datangpo Formation (1635.40 m, drillcore ZK2115); (K-L) Pyrite framboids in black shale unit of Datangpo Formation (1621.40 m and 1607.40 m, respectively, drillcore ZK2115). H-L are SEM photos. Abbreviations: D–detritus; BB– bitumen ball; ML–Mn layer; PL–pyrite layer; CL–clay layer; PF–pyrite framboids; PG– Pyrite grain.

Fig. 4. Geochemical data of Site 1 for the Mn-carbonate unit in two nearby underground mine sections (PM001 and PM002), Lijiawan manganese deposit.

Fig. 5. Geochemical data of Site 2 for the Tiesi’ao Formation and the Mn-carbonate and black shale units of the Datangpo Formation in drillcore ZK2115, Gaodi manganese deposit. The limited number of CAS analyses (n = 4) was due to insufficient sample powder for the remaining samples.

Fig. 6. Geochemical data of Site 3 for the Mn-carbonate unit in the drillcore ZK4207, Xixibao manganese deposit. TOC data are from Yu et al. (2016).

Fig. 7. Crossplots: (A) δ34SCAS versus δ34Spy, (B) TOC versus TS, and (C) Fe versus TS.

Fig. 8. Al/(Al+Fe+Mn) vs Fe/Ti discriminant plot (Marchig et al., 1982; Boström, 1983).

Fig. 9. Distribution of sulfur isotopic compositions by relative water depth of the study sections: (A) CAS δ34S and (B) pyrite δ34S for the Mn-carbonate unit, and (C) pyrite δ34S values in the black shale unit. The between-section differences are statistically significant at p < 0.01 based on a Student’s t test (see SI file).

Fig. 10. Sulfur-cycle model for the Sturtian Nanhua Basin. Within hydrothermal vents, formation of H2S through thermochemical sulfate reduction (TSR) produced masses of petrographically distinctive, 34S-enriched pyrite (Cui et al., 2017a, 2017b, 2018). Large quantities of hydrothermally sourced H2S were released to the water column, partially oxidized, and mixed through the basin, generating sedimentary pyrite and CAS with similar δ34S values. Mixing of sulfate generated in this manner with open-ocean sulfate advected to the surface layer of the Nanhua Basin resulted in water-depth gradients in CAS and in pyrite following MSR.

Table 1 Locati on Site 1a Lijiaw an (PM00 1)

Site 1b Lijiaw an (PM00 2)

Site 2 Gaodi (ZK21 15)

Formation Datangpo Fm. Mn-carbona te unit

Datangpo Fm. Mn-carbona te unit

Datangpo Fm. Black shale unit

dept h

Al2 O3

Si O2

M n

Fe

(m)

%

%

%

L-B120

0.20

18. 3

L-B119

0.50

3.6

L-B118

0.80

8.2

L-B117

1.10

5.2

61. 8 36. 5 24. 8 20. 9

0. 3 11 .5 12 .4 14 .8

L-B298

0.38

L-B299

0.75

L-B300

1.13

L-B301

1.52

6.3

L-B302

1.91

4.3

L-B303

2.29

8.1

L-B304

3.44

3.3

L-B305

3.82

2.3

47. 3 37. 3 57. 6 35. 0 24. 4 44. 1 15. 4 14. 7

5. 1 9. 2 0. 6 12 .0 15 .0 8. 7 16 .9 18 .9

ZK2115 -H49 ZK2115 -H47 ZK2115 -H45 ZK2115 -H43 ZK2115 -H41 ZK2115 -H39 ZK2115 -H35 ZK2115 -H33

1597 .40 1599 .40 1601 .40 1603 .40 1605 .40 1607 .40 1611 .40 1613 .40

17. 1 16. 6 15. 9 16. 9 16. 0 16. 2 16. 6 17. 8

63. 3 63. 5 65. 8 62. 1 62. 2 59. 1 59. 7 60. 6

0. 2 0. 2 0. 1 0. 1 0. 2 0. 4 0. 5 0. 3

Sample No.

13. 6 10. 6 18. 6

TO C

TS

%

%

%

3.3

1.4

2.3

2.3

0.7

0.9

4.7

0.7

3.6

3.4

0.6

2.2

3.9

1.7

2.7

2.9

1.6

1.1

3.7

2.0

2.8

1.7

1.2

0.5

1.9

1.0

0.6

1.9

1.6

1.5

1.9

2.1

0.6

2.7

2.0

1.4

3.8

1.0

2.2

3.8

1.2

2.0

3.0

1.3

1.7

3.9

1.9

2.7

3.5

2.9

2.4

4.1

3.1

3.5

4.2

2.7

3.2

3.8

1.9

3.1

δ34 Spy

δ34S CAS

△34 S

‰

‰

‰

41.1

2.9

47.1

-1.6

51.8

0.0

57.2

8.0

44.3

3.2

53.3

9.1

38.8

0.6

38. 2 48. 7 51. 8 49. 2 41. 1 44. 2 38. 2 47. 4 37. 8 38. 4 46. 8 52. 0 30. 9 31. 5 29. 1 51. 6 45. 8 51. 6 56. 2 56. 9

59.6 49.6

12. 2 11. 8

41.6

3.2

54.6

7.8

52.5

0.5

Fe/ Ti 10. 4 35. 2 15. 6 17. 4 13. 9 14. 6 10. 5 15. 2 19. 3 11. 5 21. 5 34. 0 9.9 10. 5 8.9 10. 7 9.2 11. 2 11. 5 11. 0

Datangpo Fm. Mn-carbona te unit

ZK2115 -H31 ZK2115 -H29 ZK2115 -H27 ZK2115 -H25 ZK2115 -H23 ZK2115 -H21

1615 .40 1617 .40 1619 .40 1621 .40 1623 .40 1625 .40

19. 6 18. 4 17. 9 17. 5 18. 1 18. 6

59. 3 60. 6 59. 8 60. 7 60. 9 60. 0

0. 2 0. 3 0. 4 0. 2 0. 2 0. 2

ZK2115 -H19

1627 .40 1627 .90 1628 .40 1628 .90 1629 .40 1629 .90 1630 .40 1630 .90 1631 .40 1631 .70 1632 .00 1632 .40 1632 .80 1633 .10 1633 .40 1633 .70

14. 2

44. 8 25. 3 24. 5 12. 9 40. 0 14. 9 59. 3 12. 8 63. 8 52. 1 10. 3 35. 5 46. 1 10. 8 14. 0 23. 8

5. 7 18 .3 18 .1 24 .5 9. 7 22 .8 0. 9 24 .0 0. 9 2. 8 26 .1 14 .2 8. 5 24 .5 24 .9 19 .9

WX-32 ZK2115 -H18 WX-33 ZK2115 -H17 WX-34 ZK2115 -H16 WX-35 ZK2115 -H15 WX-36 ZK2115 -H14 ZK2115 -H13 ZK2115 -H12 ZK2115 -H11 ZK2115 -H10 ZK2115 -H9

5.5 4.9 2.4 12. 1 2.2 16. 7 2.6 14. 0 20. 5 2.1 8.7 13. 0 1.0 2.9 4.8

3.6

1.3

3.1

3.4

1.7

2.6

3.6

1.7

2.9

4.5

1.6

3.7

3.5

1.8

2.9

3.7

1.6

3.1

4.1

1.6

3.7

2.7

2.8

2.4

2.3

2.1

1.1

1.4

2.6

0.6

2.8

1.8

2.5

2.0

2.6

1.4

3.7

2.0

3.8

1.6

2.7

0.8

3.2

2.9

3.0

2.5

1.6

2.7

1.5

2.6

0.7

2.8

2.3

2.6

2.6

2.1

2.3

1.1

1.1

0.7

2.5

2.2

2.2

2.8

2.6

2.2

55. 4 65. 1 67. 2 66. 0 63. 8 65. 4

10. 2 10. 0 10. 8 13. 9

64. 2 53. 7 53. 4 56. 7 56. 1 56. 3 54. 0 54. 6 64. 5 56. 6 72. 4 70. 2 67. 2

13. 6 14. 8 22. 2 19. 7 11. 9 29. 7 11. 4 19. 5 10. 7 10. 0 21. 5 14. 6

70. 1 69. 3

9.8 10. 9

64.6

10. 9

69.8

13. 5

62.6

8.0

52.7

-3.8

9.6 22. 3 27. 7 17. 2

Tiesiao Fm.

ZK2115 -H8 ZK2115 -H7 ZK2115 -H6 ZK2115 -H5 ZK2115 -H4 ZK2115 -H3 ZK2115 -H2 ZK2115 -H1

1634 .10 1634 .60 1635 .10 1635 .40 1635 .70 1636 .25 1636 .55 1637 .15

16. 7

16. 9

58. 8 12. 3 41. 1 54. 1 22. 8 20. 5 23. 4 60. 3

0. 2 26 .9 7. 9 2. 2 22 .4 23 .1 21 .2 0. 4

ZK2115 -H51 ZK2115 -H52 ZK2115 -H53 ZK2115 -H54 ZK2115 -H55 ZK2115 -H56 ZK2115 -H57 ZK2115 -H58 ZK2115 -H59 ZK2115 -H61 ZK2115 -H63 ZK2115 -H64 ZK2115 -H65 ZK2115 -H66

1637 .98 1638 .28 1638 .58 1638 .88 1639 .18 1639 .48 1639 .78 1640 .08 1640 .38 1640 .98 1641 .58 1641 .88 1642 .48 1642 .78

14. 4 13. 4 12. 8 10. 9 16. 8 12. 7 14. 7 13. 1 14. 8 12. 6 13. 5 15. 3 14. 8 15. 9

63. 6 70. 3 68. 2 70. 1 58. 4 54. 7 64. 3 55. 0 66. 1 58. 3 62. 4 63. 2 63. 0 63. 2

0. 3 0. 3 0. 6 0. 8 0. 1 1. 1 0. 5 1. 1 0. 3 0. 6 0. 4 0. 1 0. 1 0. 1

1.2 14. 8 15. 8 3.5 3.2 3.8

4.2

3.1

4.1

1.2

2.5

0.5

4.3

2.2

4.3

4.5

2.2

4.8

1.3

1.7

0.8

1.1

1.9

0.5

1.6

2.3

1.0

3.9

2.4

3.7

4.7

0.2

4.5

1.9

0.1

1.7

2.6

0.7

2.5

2.4

0.7

2.2

4.6

2.7

4.3

2.7

0.1

2.5

2.7

0.1

2.5

2.6

0.1

2.2

2.8

0.1

3.2

2.3

0.1

2.0

3.0

0.1

2.9

4.2

0.1

3.4

5.2

0.1

4.6

4.5

0.1

0.0

63. 1

57. 1 68. 8 59. 0 63. 1 65. 4 59. 9 39. 8 47. 1 51. 4 54. 3 45. 0 36. 4 40. 6 40. 8 45. 5 51. 4 55. 2 58. 3 60. 6 59. 9

10. 9 28. 2 11. 9 10. 7 10. 6 8.7 13. 8 11. 1 14. 4 7.1 9.8 10. 4 10. 9 8.1 7.2 7.8 8.3 6.6 7.8 7.6 9.1 7.5

Site 3 Xixiba o (ZK42 07)

Datangpo Fm. Mn-carbona te unit

ZK2115 -H68 ZK2115 -H69 ZK2115 -H71 ZK2115 -H72 ZK2115 -H74 ZK2115 -H75 ZK2115 -H77 ZK2115 -H78 ZK2115 -H79 ZK2115 -H82

1643 .38 1643 .68 1644 .28 1644 .58 1645 .18 1645 .48 1646 .08 1646 .38 1646 .68 1647 .58

ZK4207 -29 ZK4207 -28 ZK4207 -27 ZK4207 -26 ZK4207 -25 ZK4207 -24 ZK4207 -23 ZK4207 -22 ZK4207 -21 ZK4207 -20 ZK4207 -19 ZK4207 -18

887. 85 888. 32 888. 73 889. 17 889. 58 890. 11 890. 53 891. 00 891. 51 891. 80 892. 31 892. 71

5.0 18. 3 8.4 7.6 3.4 11. 3 12. 0 10. 5 8.8 8.9 14. 1 14. 6 12. 3 11. 3 11. 5 8.7 9.7 6.3 12. 6 7.8 7.7 6.5

22. 8 58. 4 53. 1 50. 5 18. 2 65. 9 64. 5 74. 6 72. 6 75. 2

1. 3 0. 1 0. 8 0. 9 1. 0 0. 2 0. 4 0. 2 0. 2 0. 1

61. 1 56. 9 48. 1 44. 3 41. 5 31. 8 32. 4 36. 6 44. 5 39. 0 34. 1 28. 6

1. 8 3. 3 6. 9 9. 8 10 .6 16 .3 14 .5 14 .7 8. 3 12 .6 14 .5 16 .8

2.6

0.0

0.6

4.8

0.1

4.0

2.5

0.1

0.8

3.0

0.0

1.2

2.6

0.1

0.3

3.8

0.1

3.3

2.3

0.1

1.2

1.6

0.1

1.1

3.6

0.0

3.3

2.5

0.1

2.7

2.2

2.3

1.9

3.0

1.9

2.7

2.9

2.3

2.4

2.9

2.5

1.8

2.7

2.0

1.9

2.4

2.7

1.0

3.3

2.3

2.5

2.0

2.2

1.0

3.6

1.4

3.7

2.9

1.7

2.7

3.1

1.6

2.8

2.5

2.2

2.0

21. 5 53. 3 48. 8 50. 3

6.8 12. 2 17. 2 30. 7 17. 2

45. 0 46. 8 46. 4 40. 6 47. 8 62. 8 69. 2 70. 7 65. 8 68. 3 65. 5 62. 4 56. 4 59. 2 54. 2 56. 7 58. 8

7.8 8.8 23. 8 14. 4 53.8

-9.0

7.6

57.6

-11. 6

9.7

61.9

-8.8

59.1

-6.7

63.4

-4.9

67.5

2.0

66.9

4.6

53.8

-2.6

57.2

-2.0

51.1

-3.1

45.6

-11. 1

61.2

2.4

10. 6 12. 3 11. 9 13. 3 17. 3 15. 0 12. 6 12. 9 13. 9 14. 0

ZK4207 -17 ZK4207 -16 ZK4207 -15 ZK4207 -14 ZK4207 -13 ZK4207 -12 ZK4207 -11 ZK4207 -10 ZK4207 -9 ZK4207 -8 ZK4207 -7 ZK4207 -6 ZK4207 -5 ZK4207 -4 ZK4207 -3 ZK4207 -2 ZK4207 -1

893. 22 893. 93 894. 60 895. 22 895. 77 896. 22 896. 63 897. 02 897. 41 897. 83 898. 44 899. 12 899. 41 899. 83 900. 47 900. 76 901. 28

3.7 5.7 10. 4 9.8 10. 7 10. 0 9.3 12. 0 9.0 9.6 8.8 11. 5 9.9 4.0 12. 4 4.0 12. 1

20. 1 27. 5 40. 0 36. 9 39. 6 35. 8 52. 0 49. 6 36. 9 44. 0 45. 5 49. 2 35. 4 22. 0 47. 3 21. 5 42. 4

21 .5 18 .4 12 .2 13 .9 12 .3 13 .8 6. 8 6. 8 13 .7 10 .6 10 .2 7. 4 13 .9 20 .7 6. 4 20 .1 5. 3

2.4

2.6

1.7

2.4

2.2

1.6

3.1

1.7

2.4

2.6

1.7

1.4

2.3

1.5

1.5

2.6

1.6

1.6

3.4

1.5

3.3

2.4

2.3

2.2

2.2

2.1

1.2

2.1

2.0

1.7

2.1

2.0

1.8

1.8

2.0

1.6

2.4

2.2

1.8

2.4

3.5

1.4

4.2

1.6

4.2

3.4

2.8

2.8

3.1

1.7

3.0

62. 4 56. 1 67. 1 62. 9 55. 2 59. 8 62. 1 52. 0 53. 2 54. 1 55. 3 47. 5 57. 7 63. 4 58. 8 59. 8 55. 9

60.5

-1.9

59.5

3.4

60.8

-6.3

62.4

-0.5

57.3

2.1

49.5

-10. 3

53.1

-9.0

59.4

7.4

62.5

9.3

60.2

6.1

51.8

-3.5

49.7

2.2

57.8

0.1

65.5

2.1

57.2

-1.6

59.7

-0.1

54.1

-1.8

19. 1 15. 9 14. 8 13. 4 11. 7 13. 3 20. 3 9.0 13. 8 11. 5 13. 0 9.2 12. 2 20. 4 12. 3 22. 7 10. 6

The Al2O3, SiO2, Mn, Fe, and Fe/Ti data of drillcore ZK4207 in Xixibao area are from Yu et al. (2016).

Table 2 Location

Formation

δ34Spy (‰)

Mean (‰)

δ34SCAS (‰)

Mean (‰)

Reference

Australia

Tapley Hill Fm.

+8.7 ~ +58.0

+28.7 (n = 48)

+10.1 ~ +27.0

+21.7 (n = 7)

Gorjan et al., 2000; Hurtgen et al., 2005

+4.5 ~ +60.7

+31.1 (n = 31)

Gorjan et al., 2000

+16.1 ~ +61.1

+37.6 (n = 8)

Gorjan et al., 2003

Aralka Fm.

Nambia

Canada East Greenlan d

South China

Court Fm. (Gobabis Member) Rasthof Fm. Basal Twitya Fm. Basal Arena Fm.

+12.4 ~ +51.5

+32.9 (n = 27)

Hurtgen et al., 2002, 2005

+11.0 ~ +44.6

+31.9 (n = 8)

Sperling et al., 2016

+10.8 ~ +63.4

+37.2 (n = 19)

Scheller et al., 2018 Liu et al., 2006; Chen et al., 2008; Wu et al., 2016; Wang et al., 2016; this study

Mn-carbon ate unit of Datangpo Fm.

+23.8 ~ +72.4

+54.3 (n = 63)

Black shale unit of Datangpo Fm.

+6.0 ~ +69.0

+48.8 (n = 47)

Large accumulations of

34S-enriched

+38.8 ~ +69.8

+56.0 (n = 45)

Feng et al., 2010; Li-C et al., 2012; this study

pyrite in a low-sulfate marine basin: the

post-Sturtian Nanhua Basin, South China

Highlights:

1. First investigation of sulfur isotopes in pyrite and CAS from Cryogenian strata of the Nanhua Basin, South China 2. Heavy δ34S compositions in pyrite (+29 to +72 ‰, mean +55 ‰) and CAS (+39 to +70 ‰, mean +56 ‰); △34S is ‒12 ‰ to +14 ‰

3. Both pyrite and CAS δ34S show water-depth gradients within the Nanhua Basin due to low seawater sulfate concentrations 4. High total sulfur content (mean 2.2 ± 1.1 %) despite evidence for low contemporaneous seawater sulfate concentrations 5. These features can be explained by massive hydrothermal emissions of 34S-enriched H2S with partial oxidation to sulfate