Mineralogy, geochemistry, and sulfur isotope characteristics of sediment-hosted hydrothermal sulfide minerals from the southern Okinawa Trough

Mineralogy, geochemistry, and sulfur isotope characteristics of sediment-hosted hydrothermal sulfide minerals from the southern Okinawa Trough

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Yamin Yang¹^,³, Zhigang Zeng¹^,²^,³^,⁴^,^*, Xuebo Yin¹, Xiaoyuan Wang¹, Shuai Chen¹, Haiyan Qi¹, Zuxing Chen¹, Bowen Zhu¹^,³

Acta Oceanologica Sinica | 2021, 40(10) : 129 - 143

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Acta Oceanologica Sinica | 2021, 40(10): 129-143

• Marine Geology •

Mineralogy, geochemistry, and sulfur isotope characteristics of sediment-hosted hydrothermal sulfide minerals from the southern Okinawa Trough

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Yamin Yang¹^,³, Zhigang Zeng¹^,²^,³^,⁴^,^*, Xuebo Yin¹, Xiaoyuan Wang¹, Shuai Chen¹, Haiyan Qi¹, Zuxing Chen¹, Bowen Zhu¹^,³

Affiliations

¹ Key Laboratory of Marine Geology and Environment, Institute of Oceanology, Chinese Academy of Sciences, Qingdao 266071, China

² Laboratory for Marine Mineral Resources, Pilot National Laboratory for Marine Science and Technology (Qingdao), Qingdao 266071, China

³ University of Chinese Academy of Sciences, Beijing 100049, China

⁴ Center for Ocean Mega-Science, Chinese Academy of Sciences, Qingdao 266071, China

Published: 2021-10-25 doi: 10.1007/s13131-021-1836-9

Outline

Abstract

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Studies of the mineralogy and sulfur isotope composition of sediment-hosted hydrothermal sulfide minerals in cores are important for understanding the seafloor mineralization environment and material source and reconstructing the hydrothermal history. However, the source of ore-forming materials and the history of hydrothermal activity in the southern Okinawa Trough (SOT) remain unclear. Here, the mineralogy and sulfur isotope characteristics of sulfides from gravity core HOBAB4-S2, collected between the Yonaguni Knoll IV hydrothermal field (HF) and the Tangyin HF, was investigated. Enrichments in Zn (up to 321×10⁻⁶), Cu (up to 73.7×10⁻⁶), and Pb (up to 160×10⁻⁶) and the presence of pyrite, galena, pyrrhotite and minor sphalerite and chalcopyrite provide evidence for the input of hydrothermal materials into the sediments. The pyrite morphologies include disseminated cubic, pentagonal dodecahedron, and framboidal forms. Except for minor framboidal pyrite, euhedral pyrite from core HOBAB4-S2 is mainly of hydrothermal origin with Co/Ni ratios >1 and S/Fe atomic ratios <2 in the Cu-Zn-Pb-rich layers. The occurrences of hexagonal pyrrhotite, high-Co (up to 0.17%) pyrite and high-Fe sphalerite indicate that the hydrothermal precipitates formed at medium-high temperatures and low-sulfur fugacity (fS₂) environments. The δ³⁴S values of sulfides (0.21‰–3.45‰) with low- fS₂ mineral assemblages (e.g., pyrrhotite ± high-Fe sphalerite) in the core are significantly lower than those of magmatic rocks and seawater, indicating possible incorporation of previously formed biogenic sulfur in the sediment. Combined with the age model of the core, it is suggested that hydrothermal activity likely began in the Tangyin HF before AD 1445–1483 and that at least three active episodes may have occurred since then.

Key words

mineralogy / sulfur isotopes / sediment-hosted hydrothermal sulfides / southern Okinawa Trough

Cite this Article

Yamin Yang, Zhigang Zeng, Xuebo Yin, Xiaoyuan Wang, Shuai Chen, Haiyan Qi, Zuxing Chen, Bowen Zhu. Mineralogy, geochemistry, and sulfur isotope characteristics of sediment-hosted hydrothermal sulfide minerals from the southern Okinawa Trough[J]. Acta Oceanologica Sinica, 2021 , 40 (10) : 129 -143 . DOI: 10.1007/s13131-021-1836-9

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1 Introduction

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The Okinawa Trough (OT) is an intracontinental back-arc basin and is located behind the Ryukyu trench-arc system (Lee et al., 1980; Sibuet et al., 1987). The Zn and Pb-enriched hydrothermal mineralization in the OT (Halbach et al., 1993; Ishibashi et al., 2015) resembles that of Kuroko-type deposits. This mineralization has attracted the attention of economic geologists since the discovery of active hydrothermal field (HF) Natsushima 84-1 Knoll in 1986 (Kimura et al., 1988). To date, at least 15 HFs have been reported in the OT (Beaulieu and Szafranski, 2018), including the Tangyin and Yonaguni Knoll IV HFs in the southern OT (SOT). Recently, considerable efforts have been made to investigate the characteristics of sulfide ore deposits (Glasby and Notsu, 2003; Gena et al., 2013; Zhang et al., 2018, 2019), the source of ore-forming material (Cao et al., 2018; Wang et al., 2020; Yang et al., 2020), hydrothermal fluids (Suzuki et al., 2008), the chemical compositions of hydrothermal vent organisms (Zeng et al., 2017a), volcanic rocks (Chen et al., 2018; Guo et al., 2018), the bacterial community of hydrothermal deposits in the SOT (Wang et al., 2018). However, the ore-forming material source of sulfides from the back-arc basin is complex and remains to be clarified despite much research effort. In addition to igneous rocks and seawater, magmatic fluids (Cao et al., 2018; Yang et al., 2020) or sediments (Wang et al., 2020) may have contributed to the hydrothermal systems in the SOT, but this role is still unclear. In addition, the history of hydrothermal activity in the SOT is also uncertain.

As records of the distribution and features of hydrothermal activity, hydrothermal sediments are mixtures of hydrothermal metal-bearing materials and background pelagic sediments (Gurvich, 2006). As the co-precipitates of sulfide ores, hydrothermal components in the sediments are an important part of the sulfide metallogenic system (Gurvich, 2006). Sediment-hosted hydrothermal components usually inherit metalliferous signatures via the discharge of hydrothermal particles or sulfide mass wasting (Feely et al., 1987). The primary input geochemical signals of hydrothermal sulfides are preserved in the sediments since there is no early diagenetic element mobilization (Feely et al., 1987; Kuhn et al., 2000; Agarwal et al., 2020). Therefore, the mineralogical and geochemical characteristics of hydrothermal sulfides can be used to reveal the hydrothermal mineralization environments and depositional processes of sediments (Yang et al., 2020). Moreover, the hydrothermal sediments in cores can be used to determine the history of hydrothermal activity (Gurvich, 2006).

The sediment core HOBAB4-S2 was collected from a seafloor depression between the active Tangyin and Yonaguni Knoll IV HFs in the SOT. In this paper, the major and trace element compositions of sediment samples and the mineralogical, geochemical and sulfur isotope compositions of sulfide minerals are reported. The aims of this study are to identify the hydrothermal components in the sediments, determine the mineralization conditions and infer the possible sulfur sources in the SOT. The occurrence time of hydrothermal activity is then constrained in conjunction with the age model of core HOBAB4-S2.

2 Geological setting

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The OT is an initially expanding back-arc basin behind the Ryukyu Arc in the western Pacific (Sibuet et al., 1987, 1995) (Fig. 1) and is formed by the subduction of the Philippine Sea Plate under the Eurasian continent (Halbach et al., 1993). The most recent extension of the OT commenced approximately 2 million years (Ma) ago at the Plio-Pleistocene boundary (Sibuet et al., 1987; Shinjo, 1998; Shinjo and Kato, 2000). Topographically, the OT is 1 200 km long, 60–100 km wide in the south and 230 km wide in the north (Sibuet et al., 1987). Seismic reflection data suggest that the sediment layer becomes thinner from north to south, with thicknesses of 1–2 km in the SOT (Klingelhoefer et al., 2009).

The OT is divided by the Tokara fault belt (TFB) and the Miyako fault belt (MFB) located at ~127°E and 130°E, respectively, into three segments: the SOT, the MOT and the NOT (Fig. 1a) (Shinjo et al., 1999; Shang et al., 2017; Zhang et al., 2018). The volcanic rocks in the SOT are mainly basalts and rhyolites (Shu et al., 2017) and minor intermediate rocks-dacites (Chen et al., 2018). Sufficient supplies of magma and heat provide favorable conditions for intense submarine hydrothermal activity in the SOT (Glasby and Notsu, 2003; Ishibashi et al., 2015; Zeng et al., 2017a). The emitted vent fluid is enriched in CO₂ and CH₄ volatile gases, indicating interactions between the hydrothermal fluid and organic matter in sediments (Ishibashi et al., 1995).

Tangyin HF in the SOT was discovered by the R/V Kexue during the HOBAB3 cruise in 2014 (Zeng et al., 2017a). It is located on the top of a twin Yuhua Hill seamount approximately 220 m high and 1.5 km wide from east to west. The seamount is characterized by a felsic volcanic basement (Zeng et al., 2017a). Sulfide types are Fe-Zn-rich massive, Zn-Pb-rich massive, and silica-rich precipitates (Wang et al., 2020). Yonaguni Knoll IV HF (Fig. 1b) is located in an elongated valley approximately 1 000 m long and 500 m wide. The valley is covered with muddy sediments and volcanic breccia. The hydrothermal mineralization types can be classified into five groups in these fields: anhydrite-rich chimneys, massive Zn-Pb-Cu enriched sulfides, Mn-rich chimneys, Ba-As chimneys, and pavement and silicified sediments (Suzuki et al., 2008). Previous moored investigations found that the bottom water in the SOT mainly flows toward the south and southeast directions with mean speeds of ~5 cm/s and 3 cm/s, respectively (Chung and Hung, 2000) (Fig. 1).

3 Samples and methods

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3.1 Sample location and chronological framework

The studied samples were recovered by a gravity corer from the SOT during the HOBAB4 cruise of the R/V Kexue at Station S2 in 2016 (Fig. 1). Sediment core HOBAB4-S2 (24°52′49.91″N, 122°37′19.70″E, water depth 1 505 m) was collected from a rifted basin between the Tangyin HF and the Yonaguni Knoll IV HF. The sediments in the core mainly consist of dark gray mud, with occasional silt and fine sand layers. The samples have a slightly fishy rotten egg odor and contain trace shell fragments. The age model of core HOBAB4-S2 was established with the OxCal 4.2 software based on accelerator mass spectrometry (AMS) ¹⁴C dating of mixed planktonic foraminifera and ²¹⁰Pb dating of the upper layers^{11Yang Yamin, Zeng Zhigang, Liu Xiting, et al. Fine-grained turbidites in the southern Okinawa Trough and its implication for earthquakes activity in the past 700 years. Geological Journal (Major revision).}. Then, three sections were chosen for the collection of sulfide mineral grains. The sedimentary age of each section was calculated, and the ages are shown in Table 1.

3.2 Analytical methods

Forty-eight sediment powder samples (<200 mesh) were chosen for the analysis of major elements. Glass discs for wavelength dispersive X-ray fluorescence (WD-XRF) spectrometry analysis were prepared by fusion of 0.5 g of the sediment powder sample with 5.0 g of lithium tetraborate:metaborate (2:1) (lithium borates 67-33, Claisse, Canada) using Pt-Au crucibles and molds employing fusion equipment supplied with gas burners (Fluxy, Claisse, Canada). The major element compositions were determined by XRF on fused glass disks using an AXIOS-Minerals spectrometer at the Institute of Geology and Geophysics, Chinese Academy of Sciences (IGGCAS) in Beijing, China. The loss on ignition (LOI) was measured after heating at 1 000°C for 1 h. The accuracy was monitored by the analysis of the reference material GBW07309. The analytical precision of the major elements was generally better than 5%.

The trace elements of sediment samples were measured at intervals of 1 cm by inductively coupled plasma mass spectrometry (ICP-MS) (ELAN DRC II) at the IOCAS. Every 40 mg sample powder was digested with 1.5 mL of HF and 0.5 mL of HNO₃ in a Teflon vessel for 12 h at 150°C before 1 mL of HNO₃ and 3 mL of deionized Milli-Q water were added to dissolve the residues. Finally, the dissolved samples were diluted to 40 g with deionized Milli-Q water for measurement using ICP-MS. A replicate for every 10 samples was measured to exclude data errors caused by instrument testing. The analytical precision and accuracy were monitored by analyzing the national marine sediment level standard reference materials GBW07315 and GBW07316 and USGS basalt standard substances BCR-2 and BHVO-2. The analytical results for the reference materials were consistent with their certified values and within the recommended ranges for standards. The relative accuracies of the ICP-MS analyses were estimated to be better than ±5%–10% for most elements.

Isolated sulfide grains in the sand-sized fractions (>63 µm) were observed and handpicked after the sand fraction was separated from the silt and clay fraction (<63 µm) via wet sieving. More than 50 sulfide minerals with coarse particles and good crystal forms were picked from the sediments. Then, the samples were ultrasonically cleaned, embedded in epoxy resin, and polished following the method of Dias et al. (2008). Thirty-five and 31 sulfide minerals were selected for major element and sulfur isotope analyses, respectively.

The major element compositions of sulfides were analyzed by a JXA-8230 electron probe microanalysis (EPMA) instrument with a 15 kV acceleration voltage, 20 nA beam current, and 1 μm beam diameter at the State Key Laboratory of Continental Dynamics, Northwest University in Xi’an, China. Many natural and synthetic minerals were used as standard samples to calibrate the sample results, including pyrite (FeS₂), galena (PbS), sphalerite (ZnS), antimony telluride (Sb₂Te₃), and the pure metals gold (Au), copper (Cu), cobalt (Co), nickel (Ni), and manganese (Mn). The minimum detection limits of the elements were less than 0.01 wt%. The elements analyzed included S, Fe, Zn, Cu, Pb, Mn, Au, Ag, Cd, As, Ni and Se. The analytical precision was better than 1%.

In situ sulfur isotope analyses of pyrite and pyrrhotite were performed using a Nu Plasma 1700 multicollector (MC)-ICP-MS instrument (Nu Instruments, Wrexham, UK) with a 193 nm ArF excimer laser ablation system (RESOlution M-50, ASI) at the State Key Laboratory of Continental Dynamics, Northwest University. The operating conditions included an energy density of 3.6 J/cm², a repetition rate of 3 Hz, and a spot size diameter of 30–37 μm, and helium was used as the carrier gas for sample laser ablation. A standard-sample bracketing (SSB) approach was used, and the calibrated standards NBS123, PTST (matrix-matched sulfur minerals), and Py-4 were repeatedly measured before and after each sample to correct for instrument drift and mass bias during the in situ sulfur isotope analysis. The external standard reference materials were Cpy-1 (natural chalcopyrite, δ³⁴S_V-CDT=4.2‰ ±0.3‰, 2SD) and Py-4 (natural pyrite, δ³⁴S_V-CDT = 1.7‰ ±0.3‰, 2SD) (Chen et al., 2017). The external reproducibility of the replicate sulfur isotope analyses is reported at the two-standard deviation (2SD) level. The analytical precision and accuracy of the sulfide measurements were monitored by repeated long-term analysis, and the results were generally better than 0.2‰. The S isotope composition was reported relative to the Vienna Canyon Diablo troilite (V-CDT, δ³⁴S_V-CDT). In detail, the sulfide measurement procedures followed the methods of Bao et al. (2017).

4 Results

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4.1 Major and trace element compositions of sediments

The major element compositions of the sediment samples show variations in SiO₂ (55.35 wt% to 69.31 wt%), Al₂O₃ (10.01 wt% to 16.84 wt%), TFe₂O₃ (3.80 wt% to 6.75 wt%) and MnO (0.05 wt% to 0.07 wt%) (Table 2). The Al/(Al+Fe+Mn) ratios of the sediment samples are 0.64 at 166–167 cm, 0.66 at 312–315 cm, and 0.66 at 444–445 cm.

The trace element concentrations of the sediment samples show wide variations in Cu (18.6×10⁻⁶ to 73.7×10⁻⁶), Zn (76.9×10⁻⁶ to 356.0×10⁻⁶) and Pb (22.1×10⁻⁶ to 160.0×10⁻⁶) (Table 2). Three layers have relatively high Zn, Cu and Pb contents (especially Zn) (Fig. 2 and Table 2). At 167 cm, the Cu, Zn, and Pb concentrations in the sediment can reach 73.7×10⁻⁶, 356.0×10⁻⁶ and 16.0×10⁻⁶, respectively. The contents of Zn and Pb in the sediment are up to 246.0×10⁻⁶ and 92.1×10⁻⁶ at 313 cm and 321.0×10⁻⁶ and 121.0×10⁻⁶ at 445 cm, respectively.

4.2 Sulfide and sulfate mineral compositions

The sulfide and sulfate minerals in the three Cu-Zn-Pb-rich layers of core HOBAB4-S2 contain mainly pyrite, some sphalerite, pyrrhotite, galena and chalcopyrite, arsenopyrite and radiated anhydrite aggregate (Figs 3 and 4). Sulfides in the sediment at 166–167 cm mostly consist of pyrite (Figs 4a–c) with some pyrrhotite and minor sphalerite, chalcopyrite (Figs 4a–c) and arsenopyrite (Fig. 4d). Morphologically, the sulfide particles are mainly granular (Fig. 3a), and euhedral pyrite separates exhibit an intergrown twinning texture (Fig. 3b). Pyrite occurs in the crystal forms of pentagonal dodecahedra (Fig. 3c) and cubes (Fig. 4b). Additionally, trace framboidal pyrite aggregates are observed in the core (Fig. 4c). Some pyrrhotite associated with arsenopyrite is found in the sediment at 166–167 cm (Fig. 4d). Sphalerite is intergrown with pyrite and occasionally tiny chalcopyrite particles (Fig. 4a). Sulfides in the sediment at 312–315 cm are mainly pyrite (Fig. 4e) with some pyrrhotite and sphalerite (Fig. 4f). Sphalerite coexists with pyrite and pyrrhotite. Sulfides from 444–445 cm are analogous to those from 312–315 cm and mainly contain pyrite, some sphalerite (Fig. 4g-h) and minor pyrrhotite (Fig. 4i).

4.3 Chemical compositions of sulfides

The contents of major elements in pyrite are shown in Table 3. For pyrite from 166–167 cm, the contents of S and Fe are 51.88–53.54 wt% (average 52.86 wt%) and 45.64–47.98 wt% (average 47.54 wt%), respectively (Table 3). The Co/Ni and S/Fe atomic ratios are 1.45–31.33 and 1.90–1.97, respectively. The S and Fe contents of pyrite from 312–315 cm are 51.68–53.35 wt% (average 52.56 wt%) and 46.56–48.04 wt% (average 47.53 wt%), respectively (Table 3). The Co/Ni and S/Fe atomic ratios are 0.51–23.20 and 1.91–1.97, respectively. The S and Fe contents of pyrite from 444–445 cm are 50.84–53.26 (average 52.49 wt%) and 46.68–48.2 (average 47.49 wt%), respectively (Table 3). The Co/Ni and S/Fe atomic ratios are 1.86–11.89 and 1.88–1.99, respectively.

The S and Fe contents of pyrrhotite from 166–167 cm are 37.69–38.11 wt% (average 37.9 wt%) and 62.93–63.33 wt% (average 63.13 wt%), respectively (Table 3). The Fe atomic percentages (at%) in pyrrhotite are 49.01%–48.50% (average 48.76%). The S and Fe contents of pyrrhotite from 312–315 cm are 38.00–38.57 wt% (average 38.29 wt%) and 61.89–62.52 wt% (average 62.21 wt%), respectively (Table 3). The Fe atomic percentages (at%) in pyrrhotite are 47.81–48.42 at% (average 48.33 at%). The S and Fe contents of pyrrhotite from 444–445 cm are 38.13–38.43 wt% (average 38.31 wt%) and 61.42–62.07 wt% (average 61.70 wt%), respectively (Table 3). The Fe atomic percentages in pyrrhotite are 47.77–47.99 at% (average 47.87 at%).

The sphalerite in thesesamples contains a certain amount of FeS (mol%) = (Fe–Cu)/(Fe–Cu + Zn); Fe, Cu, and Zn are in mol%; Wang et al., 2020) in addition to the major elements S and Zn. The FeS contents of sphalerite from 312–315 cm and 444–445 cm are 28.16–30.30 mol% and 25.84–31.29 mol%, respectively (Table 4).

4.4 The sulfur isotopes of pyrite and pyrrhotite

In total, 30 LA-ICP-MS spot analyses were performed on representative pyrite and pyrrhotite for sulfur isotope compositions, including 18 spots on pyrite and pyrrhotite from S2-3, four spots on pyrite and pyrrhotite from S2-5, and eight spots on pyrite and pyrrhotite from S2-8. The sulfur isotope dataset is provided in Fig. 5 and Table 5. In addition to three framboidal pyrite aggregates that have relatively low isotopic values of −13.22‰ and −44.06‰, most euhedral-granular pyrite minerals show narrow variable positive δ³⁴S values (Fig. 4). Except for framboidal pyrite, the δ³⁴S values of pyrite and pyrrhotite from S2-3 are 0.22‰–3.45‰ and 0.21‰–1.82‰, respectively (Fig. 5a). The δ³⁴S values of pyrite in the S2-5 samples are 1.01‰–1.02‰, while the sulfur isotope composition of pyrrhotite is 2.09‰ (Fig. 5b). For pyrite from S2-8, the δ³⁴S values range from 0.92‰ to 3.44‰, with an average of 2.41‰ (Fig. 5c).

5 Discussion

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5.1 Identification of the hydrothermal signature in the sediments

5.1.1 The major and trace element compositions of sediments

The Al/(Al+Fe+Mn) ratio can be used to differentiate metalliferous sediments from nonmetalliferous sediments (Boström et al., 1969; Adachi et al., 1986; Rona, 1988; Li et al., 2014; Yu et al., 2018) with the assumption that Al is entirely derived from the continental crust. The observed values in the study sediments range from 0.64 to 0.66, which are slightly lower than that of the continental crust (0.69; Gibbs, 1986) and close to the value of hydrothermal sediments in the SOT (0.57; Yang et al., 2018). The large amount of terrigenous material accumulated in the sediments of the SOT (Diekmann et al., 2008; Li et al., 2009; Dou et al., 2016) may mask the signature of the small-volume hydrothermal fraction and increase the Al/(Al+Fe+Mn) index.

As hydrothermal-sensitive parameters, the trace elements Cu, Zn and Pb have been studied (Zhai et al., 2007; Yang et al., 2020). Cu, Zn and Pb show relatively consistent variations with depth, although the ranges differ (Fig. 2). Several Cu-Zn-Pb-rich layers are found in core HOBAB4-S2 (Fig. 2 and Table 2). Among them, three layers exhibit anomalous Zn, Cu and Pb contents in the sediment samples at 166–167 cm, 312–315 cm and 444–445 cm, which are much higher concentrations than those of surface sediments in the OT (Zn=91×10⁻⁶, Cu=23×10⁻⁶ and Pb=18×10⁻⁶; Zhao et al., 1997). Although the element contents of terrigenous clastic sediments may be affected by grain size in marginal sea basins, high element contents are mainly concentrated in relatively fine-grained sediments (Zhao, 1983). However, the Cu, Zn, and Pb contents in this core show poor correlations with the mean grain size (Supplementary information Fig. S1), indicating a weak influence of particle size on the element contents. The sediments are composed of coarse silt at 166–167 cm, while fine sand is the main component of the sediment at 312–315 cm and 444–445 cm. These features do not seem to be consistent with the influence of element granularity for terrigenous clastic sediments (Zhao, 1983). It is inferred that the enrichments in Cu-Pb-Zn in the HOBAB4-S2 sediments may not be related to terrestrial materials but that other metal-rich hydrothermal materials may be added. Some sulfide aggregates of pyrite, pyrrhotite, sphalerite and minor chalcopyrite are observed in these three layers (Figs 3 and 4). The occurrence of sphalerite may be the main factor for the high Zn contents. Therefore, it is suggested that additional metal-rich hydrothermal materials contribute to the core HOBAB4-S2 sediments.

5.1.2 Pyrite and its genesis

At least three types of pyrite in terms of shape, including disseminated cubic, pentagonal dodecahedral, and framboidal pyrite, are observed in core HOBAB4-S2 (Fig. 4). Pyrite in marine sediments may be derived from hydrothermal pyrite and low-temperature (<100°C) sedimentary pyrite (Hu et al., 2020). These two types have distinct differences in trace elements and sulfur isotopes. Thus, the chemical and sulfur isotope compositions of pyrite can be used to determine its genesis (Bajwah et al., 1987; Hu et al., 2019). Three framboidal pyrites have strong negative sulfur isotope values ranging from −13.22‰ to −44.06‰ (Fig. 5 and Table 5), which are similar to the δ³⁴S values of sedimentary pyrite (–40‰ to –10‰) formed by low-temperature bacterial sulfate reduction (BSR) in modern sediments (Brunner and Bernasconi, 2005). This result suggests that framboidal pyrite may be the product of BSR-related sedimentary pyrite. Since pyrite with framboidal texture is rare in these samples, it is not discussed further. The origin of the euhedral pyrite is discussed.

Many pyrites observed in this core are intergrown with sphalerite (Figs 4a, e, g and h), indicating that they are products of hydrothermal activity. The S/Fe atomic ratio in pyrite can be used to determine the formation conditions of this mineral (Mozgova et al., 2008). Generally, S/Fe is lower than its theoretical value (S/Fe^*=2) for hydrothermal pyrite and close to 2 for sedimentary pyrite (Liu et al., 2018). The S/Fe atomic ratio values of pyrite at depths of 166–167 cm, 312–315 cm, and 444–445 cm are 1.91–1.97, 1.91–1.97, and 1.88–1.99, respectively, with corresponding average values of 1.95, 1.94 and 1.94 (Fig. 6a), indicating that pyrite in these samples has a hydrothermal origin.

Cobalt easily enters the pyrite lattice by isomorphism, whereas Ni is not enriched under high-temperature conditions (Tivey et al., 1995). The different physicochemical conditions for pyrite formation can cause changes in Co and Ni contents and different Co/Ni ratios (Raymond, 1996; Clark et al., 2004; Zhang et al., 2017). Therefore, the Co/Ni ratio in pyrite is regarded as an empirical indicator of the depositional environment (Clark et al., 2004; Koglin et al., 2010). In general, the Co/Ni ratio of hydrothermal pyrite is greater than 1 (Koglin et al., 2010; Zhang et al., 2017; Liu et al., 2018), whereas that of sedimentary pyrite is typically much lower (<1, average 0.63) (Price, 1972; Bajwah et al., 1987). The Co/Ni values of pyrite at depths of 166–167 cm, 312–315 cm, and 444–445 cm are 1.45–31.33, 0.51–23.20, and 1.86–11.89, respectively. Most of the pyrite samples exhibit Co/Ni ratios >1, and only one ratio is <1 (Fig. 6b), which indicates that the pyrite grains from the three layers in core HOBAB4-S2 are mainly of hydrothermal origin.

As mentioned above, the element compositions and sulfur isotope compositions of pyrite from 166–167 cm, 312–315 cm, and 444–445 cm in core HOBAB4-S2 show that most of the pyrite is hydrothermal pyrite and that framboidal pyrite is minor.

5.2 Chemical compositions of hydrothermal minerals and implications for mineralization conditions

Experimental studies of pyrite in hydrothermal systems show that pyrite with pentagonal dodecahedron shapes is dominant when the temperature is between 200°C and 300°C, while the cubic form of pyrite is dominant when the temperature is <200°C or >300–350°C (Cai and Zhou, 1994). Euhedral cubic pyrite (Fig. 4b) is the most common in these samples, followed by pentagonal dodecahedral and framboidal pyrites, reflecting moderate-high temperatures of mineralization. Moreover, the Co contents of pyrite formed at high temperatures, moderate temperatures and low temperatures are >0.1%, 0.01%–0.1% and less than 0.01%, respectively (Liu et al., 2018). The Co contents of pyrite from the sediments at 167–168 cm, 312–315 cm and 444–445 cm are 0.06%–0.17% (except for one below the detection limit), 0.05%–0.15% and 0.06%–0.13% (>0.01%), respectively, suggesting that pyrite was formed in moderate- and high-temperature environments.

The Fe atomic percentages in both monoclinic pyrrhotite and hexagonal pyrrhotite are 33.3–47.5 at% and 47.5–50.0 at%, respectively (Arnold, 1969). The Fe atomic percentages in pyrrhotite (47.77–49.01 at%) in this study fall within the range of values for hexagonal pyrrhotite, which indicates that the pyrrhotite in core HOBAB4-S2 is mainly hexagonal (Carpenter and Desborough, 1964). The formation temperature of hexagonal pyrrhotite was estimated to be at least 257°C based on the Fe atomic content in the Fe-S diagram according to Kissin and Scott (Kissin and Scott, 1982). Moreover, the precipitation of pyrrhotite in response to the low-sulfur fugacity (fS₂) hydrothermal fluid indicates reducing conditions (Suzuki et al., 2008).

Fe²⁺ is known to substitute for Zn²⁺ in the sphalerite structure without destroying the lattice structure, and almost all sphalerite minerals contain a certain amount of Fe (Pring et al., 2008). Some hydrothermal conditions and mineralization processes can be inferred from the content of Fe in sphalerite (Scott, 1983; Koski et al., 1984). High-Fe sphalerite can reflect low-fS₂ conditions (Scott and Barnes, 1971; Keith et al., 2014; Wang et al., 2020). Sphalerite in samples from core HOBAB4-S2 has high FeS contents (25.84–31.29 mol%) (Table 4), indicating low-fS₂ reducing conditions.

The occurrence of euhedral pyrrhotite, pyrite with high Co contents, and high-Fe sphalerite in the samples from core HOBAB4-S2 indicates that the hydrothermal precipitates can be characterized by moderate-high temperatures and low-fS₂ reducing environments. Low-fS₂ mineral assemblages (e.g., pyrrhotite ± high-Fe sphalerite ± isocubanite) have been observed in hydrothermal precipitates from the Hakurei (Kawasumi and Chiba, 2017) and Tangyin HFs (Wang et al., 2020) in the OT. Sediment-rich mid-ocean ridge HFs such as the Guaymas Basin and Middle Valley rift are also low-fS₂ reducing environments, resulting from increased H₂ concentrations during the interaction between hydrothermal fluids and organic-rich sediments (Hannington et al., 1995; Rushdi and Simoneit, 2002; Kawasumi and Chiba, 2017).

5.3 The low sulfur isotope composition

The δ³⁴S values of pyrite and pyrrhotite from 166–167 cm, 312–315 cm and 444–445 cm in core HOBAB4-S2 are 0.21‰–3.45‰, 1.01‰–2.09‰ and 0.92‰–3.44‰, except for three framboidal pyrites (δ³⁴S=−13.22‰ to −44.06‰). These results are similar to the range of sulfur isotope values in sulfides from mature back-arc basins, such as the southern Mariana Basin (0.8‰ to 3.5‰) (Ikehata et al., 2015) and the North Fiji Basin (1.7‰ to 4.8‰, with an average value of 3.4‰) (Kim et al., 2006; Zeng et al., 2017b), and fall within the range of mid-ocean ridge (sediment-free) samples (−0.8‰ to +6.3‰, with an average value of 3.2‰) (Herzig et al., 1998). However, the values are lower than those of sulfides from the Iheya North Knoll HF in the MOT (8.8‰ to 10.5‰, with an average value of 9.9‰) (Ueno et al., 2003) and the Jade HF in the NOT (5.2‰ to 7.2‰, with an average value of 6.3‰) (Halbach et al., 1989; Zeng et al., 2002) (Table 6 and Fig. 7). This contrast indicates that the sulfur sources of the samples from the core are different from those in the Iheya North Knoll and Jade HFs.

The δ³⁴S values of the sulfides from the core (0.21‰–3.45‰) are all lower than those of island arc volcanic rocks (andesite and rhyolite, δ³⁴S=5‰ to 7‰) (Ueda and Sakai, 1984; Woodhead et al., 1987) and seawater-derived S (δ³⁴S=21.0‰±0.2‰) (Rees et al., 1978; Kim et al., 2006). Thus, a two-component mixing model of crustal host rocks and seawater cannot explain the anomalously low δ³⁴S values (down to 0.21‰), indicating the addition of a light sulfur source. The possible causes for the low sulfur isotope compositions in hydrothermal systems include phase separation (Mckibben and Eldridge, 1990), the contribution of magmatic volatiles to hydrothermal systems, and biogenic sulfur (Liao et al., 2018; Meng et al., 2020).

Boiling of hydrothermal fluids and phase separation can lead to the loss of H₂S and H₂ to the vapor phase, changing the oxidation state of the residual fluid phase (Drummond and Ohmoto, 1985; Wang et al., 2020). Increasing the oxidation state would lead to lower δ³⁴S values for H₂S in the fluid phase (Mckibben and Eldridge, 1990). The temperature of hydrothermal fluid needs to reach approximately 342°C to boil at this depth (1 505 m) based on the calculation by Bischoﬀ and Rosenbauer (Bischoff and Rosenbauer, 1985). The formation conditions of sulfides from core HOBAB4-S2 involved moderate-high temperature, and phase separation may not occur under such conditions. Moreover, the δ³⁴S values of sulfides in the Jade (δ³⁴S=6.0‰ to 7.3‰) (Lüders et al., 2001) and Iheya North Knoll HFs (δ³⁴S=8.0‰ to 10.5‰) (Ueno et al., 2003), where subseafloor phase separation of hydrothermal fluid occurred, are positive and homogeneous. These δ³⁴S values are all higher than those (δ³⁴S=0.21‰ to 3.45‰) for sulfide samples from core HOBAB4-S2. Therefore, phase separation may not be a factor for the low δ³⁴S values of sulfide.

The disproportionation of volatile SO₂ into H₂S and H₂SO₄ at temperatures below 400°C during magmatic degassing results in the formation of ³²S-rich sulfide and ³⁴S-rich sulfate, respectively (Herzig et al., 1998). The sulfur isotope fractionation between these phases may reach 16‰–21‰ (Kusakabe et al., 2000), with δ³⁴S<5‰ for the generated sulfides and δ³⁴S>5‰ for the sulfates. The sulfides from the Hine Hina HF in the Lau Basin (−7.7‰ to −2.8‰) (Herzig et al., 1998), the Susu Knolls (−8.0‰ to −3.9‰) and the PACMANUS HF (−0.8‰ to 5.6‰) (Kim et al., 2004; McDermott et al., 2015) in the Manus Basin have negative δ³⁴S values. High-fS₂ mineral assemblages (e.g., bornite ± low-Fe sphalerite ± enargite) indicating the incorporation of magmatic S have also been observed in these HFs (Moss and Scott, 2001; Kawasumi and Chiba, 2017). These results suggest that the sulfur in hydrothermal sulfides from these HFs was derived from magmatic SO₂ disproportionation and generated H₂S (Kim et al., 2004; de Ronde et al., 2005). The δ³⁴S values of sulfide (0.21‰–3.44‰) in the core may be similar to those formed by the disproportionation of magmatic SO₂. However, the low-fS₂ reducing conditions indicated by low-fS₂ mineral assemblages (e.g., high-Fe sphalerite ± pyrrhotite) are different from the high-fS₂ conditions indicative of a contribution from magmatic SO₂. Therefore, it isinferred that the contribution of magmatic volatiles to the hydrothermal fluid may not be responsible for the low sulfur isotope compositions of sulfides from core HOBAB4-S2.

In addition, biogenic sulfur is another source of light sulfur isotopes (Wang et al., 2020). Sulfate-reducing bacteria in sediments can reduce SO₄²⁻ in seawater to S²⁻, resulting in sulfides with negative δ³⁴S signatures (Rudnicki et al., 2001; Seal, 2006). Hydrothermal fluids can leach these previously formed sulfides from sedimentary strata during fluid-sediment interactions (LaFlamme et al., 2018; Wang et al., 2020). This mechanism has been used to explain the lighter δ³⁴S values of hydrothermal pyrite away from the main vent site in the Iheya North Knoll HF (LaFlamme et al., 2018) and the low sulfur in the Tangyin HF (Wang et al., 2020). The sulfur isotope compositions of hydrothermal sulfides in the core (0.21‰–3.45‰) are similar to those of Fe-Zn-rich samples (δ³⁴S=−1.32‰ to 4.97‰) reported by Wang et al. (2020). In addition, framboidal pyrite (δ³⁴S=−13.22‰ to −44.06‰) formed by BSR (MacLean et al., 2008) is observed in core HOBAB4-S2 (Fig. 4c), which confirms the presence of sulfides formed by BSR in the sedimentary strata of the SOT. These sulfides are likely to be leached during high-temperature hydrothermal fluid circulation to provide a light sulfur component to the hydrothermal system. The results support the mechanism proposed by Wang et al. (2020) that biogenic sulfur may be responsible for the light sulfur isotopes of hydrothermal sulfide samples in the Tangyin HF.

5.4 The timing of possible hydrothermal activity in the Tangyin HF

Hydrothermal components in deep-sea sediments are mainly formed in two ways: (1) diffusion and settlement of hydrothermal materials from a neutrally buoyant plume (German et al., 1993; Mills et al., 1993; Dias et al., 2008) and lateral dispersion of the plume to form widely distributed hydrothermal sediments with high Fe (20.01–39.33 wt%) and Mn (4.00–13.65 wt%) contents (Barrett et al., 1987) and (2) mass wasting and erosion of sulfide debris formed near vents. Previous studies have shown that the particle size in the hydrothermal plume is no more than 70 μm (Gurvich, 2006), and the diameters of the majority of particles are less than 2 µm (Walker and Baker, 1988). However, the grain size of sulfides is greater than 63 µm. The contents of Fe (2.66–4.72 wt%) and Mn (0.04–0.05 wt%) in the sediments of core HOBAB4-S2 are lower than those in hydrothermal sediments formed by the dispersion of the hydrothermal plume. Therefore, sulfide materials may migrate and be transported by mass wasting and the vigorous settlement of sulfides from the hydrothermal plume.

Core HOBAB4-S2 is located south of the Tangyin HF and northwest of the Yonaguni Knoll IV HF (Fig. 1b). The plume disperses laterally in the direction of the bottom current to the south or southeast in the SOT (Chung and Hung, 2000), suggesting that the Tangyin HF is more likely to be the source of sulfides in this core. Samples of hydrothermal precipitates from the Tangyin HF are rich in Fe, Zn, Pb, silica and native sulfur (Zhang et al., 2019; Wang et al., 2020). Metalliferous sediments are enriched in pyrite, sphalerite, galena and minor pyrrhotite (Yang et al., 2020). The sulfides in the Yonaguni Knoll IV HF are mainly composed of sphalerite with minor pyrite and chalcopyrite (Zhang et al., 2019). Pyrite, galena, pyrrhotite and minor sphalerite are observed in the sediments of core HOBAB4-S2 (Figs 3 and 4), which are similar to those in the Tangyin HF. Therefore, the sulfides in the core are mainly derived from the nearby Tangyin HF and can shed light on the history of hydrothermal activity.

Three hydrothermal layers in core HOBAB4-S2 may record three episodes of hydrothermal activity in the Tangyin HF. According to the age model of the core, the deposition times of these layers are AD 1777–1815, AD 1600–1639 and AD 1445–1483 (Table 2). The time of hydrothermal activity slightly predates the deposition of sulfides in the sedimentary layers. Therefore, it can be deduced that Tangyin was active before AD 1445–1483, and at least three episodes of hydrothermal activity may have occurred since then.

6 Conclusions

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(1) Hydrothermal components from the Tangyin HF were identified in core HOBAB4-S2.

(2) The chemical compositions of sulfides suggest that those in the sediments formed at medium-high temperatures and under low-fS₂ reducing conditions.

(3) The sulfide-rich samples in core HOBAB4-S2 exhibit a narrow range of δ³⁴S values (0.21‰–3.45‰) with the lower δ³⁴S values (down to 0.21‰), indicating that there may be another light sulfur source. The addition of biogenic sulfur may be the most likely reason for the lower sulfur isotope compositions.

(4) Hydrothermal activity began in the Tangyin HF before AD 1445−1483 and may have occurred at least three times since then.

Acknowledgements

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We thank all the crew and scientists involved in the R/V Kexue during Cruise HOBAB4. We also thank two anonymous reviewers for their constructive comments and suggestions, Wenqiang Yang and Yongsheng Gai for the EPMA analysis of sulfides, Kaiyun Chen and Zhian Bao for the in situ sulfur isotope measurements of sulfides, and Xia Zhang for their help in revising the manuscript and responding to comments.

Funding

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The National Natural Science Foundation of China under contract No. 91958213; the Strategic Priority Research Program of the Chinese Academy of Sciences under contract No. XDB42020402; the National Programma on Global Change and Air-Sea Interaction under contract No. GASI-GEOGE-02; the International Partnership Program of the Chinese Academy of Sciences under contract No. 133137KYSB20170003; the Special Fund for the Taishan Scholar Program of Shandong Province under contract No. ts201511061; the National Key Basic Research Program of China under contract No. 2013CB429700.

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doi: 10.1007/s13131-021-1836-9

Receive Date：2020-10-25
Online Date：2026-03-06
Published：2021-10-25

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Received：2020-10-25
Accepted：2021-03-26

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The National Natural Science Foundation of China under contract No. 91958213; the Strategic Priority Research Program of the Chinese Academy of Sciences under contract No. XDB42020402; the National Programma on Global Change and Air-Sea Interaction under contract No. GASI-GEOGE-02; the International Partnership Program of the Chinese Academy of Sciences under contract No. 133137KYSB20170003; the Special Fund for the Taishan Scholar Program of Shandong Province under contract No. ts201511061; the National Key Basic Research Program of China under contract No. 2013CB429700.

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¹ Key Laboratory of Marine Geology and Environment, Institute of Oceanology, Chinese Academy of Sciences, Qingdao 266071, China

² Laboratory for Marine Mineral Resources, Pilot National Laboratory for Marine Science and Technology (Qingdao), Qingdao 266071, China

³ University of Chinese Academy of Sciences, Beijing 100049, China

⁴ Center for Ocean Mega-Science, Chinese Academy of Sciences, Qingdao 266071, China

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*E-mail: zgzeng@ms.qdio.ac.cn

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表12种不同金属材料的力学参数

科 Family	属数 Number of genus	种数 Number of species	占总种数比例 Percentage of total species (%)	属 Genus	种数 Number of species	占总种数比例 Percentage of total species (%)
鹅膏菌科Amanitaceae	2	11	5.26	鹅膏菌属 Amanita	10	4.78
小菇科 Mycenaceae	2	12	5.74	丝盖伞属 Inocybe	5	2.39
多孔菌科 Polyporaceae	8	14	6.70	蜡蘑属 Laccaria	5	2.39
红菇科 Russulaceae	3	23	11.00	小皮伞属 Marasmius	6	2.87
				小菇属 Mycena	11	5.26
				光柄菇属 Pluteus	5	2.39
				红菇属 Russula	17	8.13
				栓菌属 Trametes	5	2.39

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Table 1. The depths and corresponding sedimentary ages of the sections ^①

Section	Depth/cm	Modeled sedimentary age/AD
Section	Depth/cm	Age min (2σ)	Age max (2σ)
S2-3	165–167	1 760	1 830
S2-5	312–315	1 581	1 656
S2-8	444–445	1 429	1 504

Table 2. Major (wt%) and trace element (10⁻⁶) concentrations of sediments from the SOT

Depth/cm	SiO₂	TiO₂	Al₂O₃	TFe₂O₃	MnO	MgO	CaO	Na₂O	K₂O	P₂O₅	LOI	Total	Cu	Zn	Pb	Cr	Co	Ni	Cd
8–9	57.41	0.79	16.36	6.39	0.06	2.62	2.26	2.46	3.28	0.17	8.78	100.57	59.0	157.0	111.8	102.0	14.3	42.0	1.28
20–21	59.69	0.78	15.14	5.86	0.07	2.29	2.75	2.08	2.88	0.17	7.70	99.42	32.6	110.0	38.4	94.6	13.7	38.3	0.28
30–31	63.72	0.71	13.78	5.37	0.07	2.11	2.74	2.21	2.68	0.20	6.86	100.46	39.3	142.0	72.2	89.1	13.7	36.0	0.48
42–43	59.79	0.79	15.39	5.87	0.07	2.36	2.93	2.23	2.98	0.16	7.80	100.37	42.7	121.0	69.8	93.3	12.9	37.9	0.47
62–63	58.73	0.77	16.02	6.01	0.07	2.40	2.54	2.36	3.18	0.16	8.23	100.47	54.8	148.0	100.0	96.0	13.0	39.0	0.91
70–71	57.47	0.80	16.54	6.32	0.07	2.43	2.52	2.21	3.28	0.17	8.41	100.24	36.2	112.0	42.6	96.0	13.6	39.8	0.36
82–83	57.50	0.76	15.88	6.12	0.06	2.58	2.46	2.44	3.14	0.16	8.68	99.77	65.0	155.0	133.0	94.0	13.2	39.0	1.59
90–91	59.93	0.79	15.62	5.87	0.07	2.39	2.88	2.14	3.05	0.16	7.27	100.17	34.2	109.0	62.9	86.7	12.8	36.0	0.55
100–101	57.55	0.76	15.66	6.08	0.06	2.63	3.34	2.29	3.13	0.15	9.59	101.24	46.9	146.0	75.6	103.0	14.9	42.0	0.92
110–111	57.60	0.79	16.42	6.19	0.06	2.56	2.73	2.22	3.26	0.15	8.47	100.46	36.5	114.0	58.5	86.3	12.3	35.1	0.74
116–117	57.26	0.78	16.25	6.18	0.06	2.56	2.59	2.28	3.26	0.15	8.66	100.03	58.9	157.0	113.0	113.0	15.2	44.4	1.68
130–131	57.15	0.81	16.84	6.33	0.07	2.51	2.68	2.11	3.37	0.15	8.46	100.49	37.9	123.0	52.0	106.0	14.9	42.1	0.52
140–141	58.92	0.81	16.06	5.99	0.07	2.37	2.79	2.06	3.18	0.15	7.92	100.31	32.7	112.0	41.6	102.0	14.5	40.6	0.40
150–151	61.80	0.77	15.00	5.53	0.06	2.21	2.82	2.01	2.91	0.15	7.08	100.35	27.8	98.8	34.1	90.2	13.1	35.9	0.34
160–161	63.60	0.73	13.95	5.17	0.06	2.08	2.88	1.93	2.74	0.15	6.60	99.88	26.7	95.2	35.0	84.4	12.7	34.9	0.33
166–167	61.32	0.75	14.71	6.09	0.06	2.30	2.56	2.18	2.97	0.17	7.48	100.60	73.7	356.0	160.0	95.0	13.9	37.0	1.88
178–179	55.35	0.75	16.81	6.34	0.06	2.59	2.53	2.33	3.41	0.15	9.12	99.44	54.4	147.0	100.0	107.0	13.9	41.2	1.94
188–189	55.79	0.78	16.67	6.35	0.06	2.58	3.05	2.33	3.36	0.15	9.28	100.40	34.5	112.0	42.4	102.0	14.1	40.5	0.32
198–199	64.11	0.75	14.19	5.25	0.06	2.30	3.61	2.18	2.76	0.17	7.42	102.79	25.8	93.1	32.5	87.7	12.6	34.7	0.26
202–203	63.07	0.71	13.33	5.04	0.06	2.23	3.82	2.08	2.64	0.16	7.36	100.49	24.8	91.8	32.2	84.8	13.3	34.7	0.28
212–213	58.41	0.80	16.18	6.16	0.06	2.38	2.68	2.14	3.21	0.16	8.30	100.47	36.2	121.0	47.3	105.0	15.0	41.1	0.50
218–219	57.43	0.77	16.06	6.21	0.05	2.56	2.58	2.30	3.20	0.15	8.52	99.83	60.7	156.0	123.0	108.0	14.5	41.9	1.52
232–233	58.63	0.80	16.26	6.06	0.07	2.35	2.69	2.09	3.19	0.16	8.18	100.48	31.2	105.0	38.0	94.8	13.6	37.8	0.34
242–243	61.47	0.76	14.86	5.55	0.06	2.16	2.81	1.98	2.90	0.16	7.20	99.91	27.2	95.1	33.0	89.9	12.8	35.0	0.28
248–249	66.68	0.70	12.98	4.49	0.05	1.92	2.88	2.07	2.48	0.15	5.92	100.31	18.6	76.9	22.1	72.6	10.1	28.9	0.26
260–261	57.82	0.78	16.02	6.19	0.07	2.48	2.92	2.26	3.14	0.16	8.66	100.51	36.0	114.0	50.5	99.7	14.4	39.4	0.38
270–271	58.34	0.80	16.44	6.22	0.07	2.41	2.81	2.11	3.22	0.16	8.58	101.16	34.9	113.0	41.4	102.0	14.5	39.7	0.40
280–281	58.48	0.77	16.20	6.19	0.06	2.61	2.70	2.30	3.25	0.15	8.70	101.41	62.6	159.0	124.0	110.0	14.6	42.8	1.68
290–291	57.01	0.78	16.25	6.22	0.07	2.58	3.29	2.23	3.23	0.15	8.78	100.58	36.2	115.0	43.0	106.0	14.7	41.8	0.33
300–301	58.52	0.77	15.23	5.85	0.07	2.48	3.58	2.19	3.01	0.17	8.64	100.49	33.7	105.0	40.5	96.6	13.8	38.0	0.30
312–315	69.31	0.60	10.01	3.80	0.05	1.91	4.78	2.10	2.04	0.14	5.78	100.51	27.0	246.0	92.1	57.0	10.6	26.0	0.92
322–323	58.09	0.80	16.15	6.08	0.07	2.39	2.82	2.10	3.21	0.16	8.47	100.33	34.7	114.0	42.6	101.0	14.3	38.7	0.34
334–335	64.69	0.70	13.51	5.04	0.06	2.02	2.93	2.11	2.61	0.20	6.46	100.33	27.0	109.0	40.1	79.6	12.8	32.2	0.36
336–337	58.69	0.78	16.31	6.30	0.06	2.61	2.65	2.39	3.25	0.16	8.58	101.77	53.6	151.0	107.0	105.0	14.2	42.7	1.69
342–343	58.14	0.78	15.97	6.12	0.06	2.45	2.72	2.10	3.15	0.16	8.32	99.97	42.1	134.0	65.8	105.0	14.1	40.0	0.78
352–353	61.07	0.76	14.93	5.63	0.06	2.30	2.97	2.09	2.89	0.16	7.50	100.35	31.8	111.0	42.8	93.5	14.1	37.3	0.34
362–363	59.42	0.80	15.89	5.97	0.07	2.34	2.76	2.06	3.11	0.16	7.87	100.45	33.0	115.0	36.6	100.0	14.0	38.6	0.35
368–369	57.03	0.76	15.73	6.75	0.07	2.59	2.34	2.44	3.17	0.35	8.72	99.94	56.3	158.0	118.0	102.0	13.7	41.3	1.20
382–383	57.69	0.77	16.22	6.33	0.06	2.59	2.58	2.27	3.23	0.17	8.49	100.40	52.5	142.0	111.0	102.0	13.5	39.5	1.01
392–393	60.25	0.78	15.20	5.69	0.07	2.32	3.10	2.03	2.95	0.16	7.74	100.28	30.4	105.0	31.9	94.8	13.6	37.0	0.34
398–399	62.26	0.74	14.99	5.58	0.06	2.32	2.96	2.20	2.97	0.16	7.52	101.75	37.4	168.0	68.5	99.4	14.2	38.4	0.58
412–413	59.90	0.77	15.09	5.72	0.06	2.31	2.98	2.06	2.93	0.16	7.95	99.94	29.8	108.0	32.6	98.2	13.5	37.2	0.27
422–423	64.36	0.72	13.48	4.98	0.06	2.09	3.21	1.96	2.60	0.16	6.65	100.27	24.0	93.0	27.8	84.2	12.5	33.8	0.26
432–433	64.96	0.67	12.60	4.81	0.05	2.06	3.71	2.03	2.48	0.18	6.90	100.45	23.8	118.0	43.1	76.4	12.8	32.1	0.46
444–445	63.93	0.66	12.03	4.82	0.06	2.17	4.27	2.19	2.41	0.19	7.16	99.88	38.8	321.0	121.0	58.0	11.7	28.0	1.17
450–451	57.75	0.78	15.96	6.18	0.06	2.54	2.69	2.26	3.19	0.15	8.92	100.48	53.4	149.0	108.0	101.0	13.5	39.0	1.63
462–463	61.44	0.76	15.03	5.50	0.06	2.21	2.76	1.94	2.91	0.15	10.1	102.88	28.3	105.0	36.0	90.0	12.7	35.4	0.35
472–473	57.47	0.77	16.08	6.29	0.06	2.59	2.49	2.24	3.24	0.18	8.63	100.04	57.6	168.0	124.0	108.0	13.7	40.0	1.28

Table 3. The major element contents (wt%) of pyrite and pyrrhotite from the OT, measured via EPMA

Depth /cm	Sample	Mineral	S	Fe	Pb	Ag	Cd	Sb	As	Se	Zn	Cu	Ni	Co	Au	Mn	Ti	Total	Fe/at%	S/Fe^*	Co/Ni
Note: S/Fe* represents the atomic ratio of sulfur to iron. Py: pyrite; Po: pyrrhotite; − means below the detection limit.
166–167	S2-3-1	Py	52.95	47.71	−	0.01	0.01	0.02	0.01	−	−	−	0.04	0.06	−	−	−	100.82		1.94	1.50
	S2-3-3-1	Py	53.01	47.93	−	0.01	0.02	−	−	0.03	0.02	−	0.01	0.07	0.10	−	−	101.19		1.94	7.00
	S2-3-3-2	Py	53.29	47.77	−	−	−	−	0.02	−	0.02	−	−	0.08	0.07	−	0.03	101.27		1.95
	S2-3-6-1	Py	53.42	48.73	−	−	−	−	0.02	0.02	0.05	−	−	0.08	−	−	−	102.30		1.92
	S2-3-6-2	Py	53.17	47.54	−	−	−	−	−	−	0.03	0.05	−	0.09	0.06	−	−	100.94		1.96
	S2-3-7-2	Py	51.88	47.54	−	−	−	−	0.01	−	−	−	0.03	0.10	−	0.02	−	99.58		1.91	3.33
	S2-3-15-1	Py	52.31	47.80	−	−	−	0.01	−	0.01	0.06	0.02	−	0.13	−	0.01	−	100.35		1.92
	S2-3-18	Po	37.69	63.33	−	−	−	−	−	0.01	−	0.14	0.01	0.09	−	−	−	101.27	49.01	1.03
	S2-3-34	Py	53.06	47.68	−	−	0.03	0.03	0.05	0.03	−	0.04	−	0.12	0.04	−	−	101.09		1.95
	S2-3-48	Py	53.10	47.80	−	0.01	0.02	−	0.07	−	0.05	0.04	0.01	0.08	−	0.08	0.03	101.28		1.94	8.00
	S2-3-65-5	Py	52.25	46.31	−	0.01	0.02	0.08	1.06	−	0.02	0.14	0.01	0.13	0.01	−	−	100.03		1.97	13.00
	S2-3-71	Py	53.03	47.05	−	−	0.01	0.03	0.13	0.08	0.01	−	0	0.01	−	0.02	−	100.45		1.97
	S2-3-83	Py	52.13	47.96	−	−	−	0.04	−	−	−	0.01	0.02	0.06	0.11	0.02	0.01	100.35		1.90	3.00
	S2-3-83	Py	53.54	47.98	−	−	−	0.03	0.01	−	0.03	0.07	0.03	0.07	−	−	−	101.77		1.95	2.33
	S2-3-84	Po	38.11	62.93	−	0.02	−	−	−	0.01	0.01	0.29	0.03	0.12	0.03	−	−	101.53	48.5	1.06
	S2-3-90	Py	53.21	47.38	−	−	−	0.02	0.02	−	−	0.03	−	0.09	0.02	−	0.01	100.78		1.97
	S2-3-104	Py	53.39	47.86	−	0.06	0.01	0.03	−	0.05	0.07	0.02	0.02	0.17	−	0.02	−	101.70		1.95	8.50
312–315	S2-5-5	Py	53.35	48.04	−	−	0.04	0.03	−	−	0.09	0.06	0.03	0.07	−	0.02	0.01	101.74		1.94	2.30
	S2-5-5	Py	53.28	48.02	−	−	0.03	0.04	−	−	0.07	−	0.01	0.10	0.06	0.02	0.02	101.68		1.94	10.00
	S2-5-6-2	Py	52.91	47.70	−	−	−	0.01	0.02	0.02	0.05	0.02	0.01	−	0.08	−	0.01	100.92		1.94
	S2-5-6-3	Py	52.80	47.70	−	0.01	−	−	0.02	0.02	−	−	0.02	0.10	−	0.03	−	100.69		1.94	5.00
	S2-5-7-2	Py	52.49	47.43	−	−	−	0.01	−	−	−	0.01	0.04	0.09	0.02	−	−	100.09		1.94	2.25
	S2-5-7-4	Py	52.09	47.09	−	−	0.01	0.02	−	−	0.06	0.01	0.02	0.08	−	0.01	0.01	99.42		1.94	4.00
	S2-5-8-2	Py	52.33	47.47	−	0.02	−	−	−	−	−	0.05	0.02	0.06	−	0.04	−	99.99		1.93	3.00
	S2-5-8-4	Py	51.68	47.27	−	−	−	−	−	−	0.08	0.02	0.02	0.05	0.06	0.07	0.02	99.30		1.91	2.50
	S2-5-16-3	Py	52.23	47.54	−	−	−	−	−	−	0.17	−	0.13	0.07	0.03	0.03	−	100.19		1.92	0.54
	S2-5-17	Po	38.00	62.52	−	−	0.04	−	−	−	−	0.22	−	0.15	0.14	−	−	101.06	48.42	1.06
	S2-5-18	Py	52.82	47.87	−	0.08	−	0.02	−	0.02	0.06	−	−	0.06	−	0.01	−	100.92		1.93
	S2-5-18-1	Py	52.30	46.56	0.02	−	−	0.01	−	−	0.08	−	0.01	0.12	−	−	−	99.07		1.97	12.00
	S2-5-22	Py	52.38	47.71	−	−	−	−	0.06	−	−	0.06	0.02	0.09	0.03	0.14	−	100.48		1.92	4.50
	S2-5-21-1	Po	38.57	61.89	−	−	0.02	−	−	−	0.11	0.12	0.05	0.12	−	−	−	100.88	47.81	1.09
444–445	S2-8-1-1	Py	53.26	48.01	−	0.04	0.02	−	0.01	0.04	0.01	−	0.04	0.09	0.01	−	−	101.52		1.94	2.25
	S2-8-1-2	Py	53.21	48.20	−	−	−	−	−	0.01	−	−	0.04	0.08	0.04	−	−	101.58		1.93	2.00
	S2-8-1	Py	50.84	47.32	−	0.01	−	0.01	0.10	−	0.04	−	0.01	0.06	0.02	−	0.01	98.42		1.88	6.00
	S2-8-1	Py	51.46	47.96	−	−	−	−	0.14	0.03	0.06	0.03	0.01	0.07	0.02	−	−	99.77		1.88	7.00
	S2-8-1-3	Py	53.09	46.68	−	0.02	−	−	−	−	0.01	−	0.01	0.11	−	0.02	0.05	99.98		1.99	11.00
	S2-8-10-1	Py	52.68	47.13	−	−	−	0.01	0.14	−	−	0.05	−	0.07	0.09	0.11	−	100.26		1.96
	S2-8-18-1	Po	38.13	61.42	−	−	0.01	−	−	0.04	0.25	0.14	0.05	0.08	0.07	−	0.01	100.19	47.85	1.08
	S2-8-18-2	Po	38.43	61.60	−	0.03	−	−	0.05	−	0.04	0.13	0.04	0.13	−	0.03	0.01	100.48	47.77	1.09
	S2-8-20-3	Py	52.70	47.91	−	0.04	0.02	0.02	0.03	−	0.11	−	0.03	0.10	0.02	0.03	−	101.01		1.92	3.33
	S2-8-22-3	Py	51.99	46.93	0.11	−	0.09	0.02	0.06	0.02	0.02	0.04	−	0.06	−	0.02	0.02	99.37		1.94
	S2-8-28-2	Py	53.06	47.18	−	−	−	0.02	0.04	−	−	0.04	−	0.09	0.09	0.01	−	100.52		1.97
	S2-8-28-3	Py	52.59	47.55	−	−	0.04	0.05	0.04	−	0.21	−	−	0.06	0.04	0.01	−	100.58		1.94
	S2-8-37-2	Po	38.37	62.07	−	0.04	0.04	−	−	0.01	0.12	0.19	−	0.10	0.10	−	−	101.05	47.99	1.08

Table 4. The major element (wt%) and FeS contents (mol%) of sphalerite from the SOT as measured via EPMA

Depth/cm	Sample	Mineral	S	Zn	Fe	Cd	Sb	As	Se	Au	Cu	Co	Mn	Ti	Total	FeS/mol%
Note: FeS (mol%)=(Fe–Cu)/(Fe–Cu+Zn); Fe, Cu, and Zn are in mol% (Wang et al., 2020). Sph: sphalerite; − means below the detection limit.
312–315	S2-5-16-4	Sph	34.20	47.28	18.08	0.08	0.05	0.01	0.02	0.06	0.43	0.01	0.12	−	100.33	30.30
312–315	S2-5-21-2	Sph	33.54	48.77	16.52	0.04	0.04	−	−	0.02	0.06	0.05	0.16	0.01	99.21	28.16
444–445	S2-8-28-4	Sph	33.23	49.09	15.37	0.05	−	0.01	−	0.02	0.41	0.05	0.75	−	99.00	26.20
	S2-8-28-4	Sph	33.28	49.22	14.96	0.04	−	−	−	−	0.21	0.01	0.99	0.01	98.73	25.84
	S2-8-37-1	Sph	33.03	46.25	18.95	0.06	0.01	0.04	−	0.01	0.92	−	0.28	0.03	99.57	31.29

Table 5. δ³⁴S values of pyrite and pyrrhotite measured via LA-ICP-MS

Depth/cm	Sample name	Sample	δ³⁴S_V-CDT/‰	2SE
Note: Py: pyrite; Po: pyrrhotite; F-Py: framboidal pyrite.
166–167	S2-3-3	Py	2.58	0.18
	S2-3-6	Py	2.35	0.17
	S2-3-7	Py	1.33	0.16
	S2-3-15	Py	3.15	0.15
	S2-3-18	Po	0.21	0.34
	S2-3-20	Py	2.36	0.13
	S2-3-32	F-Py	−13.22	0.18
	S2-3-43	F-Py	−44.06	0.13
	S2-3-34	Py	2.68	0.17
	S2-3-41	Py	0.22	0.14
	S2-3-46	Py	2.16	0.23
	S2-3-48	Py	1.36	0.15
	S2-3-65	Py	0.51	0.23
	S2-3-66	Py	2.59	0.11
	S2-3-83	Py	2.54	0.17
	S2-3-84	Po	1.82	0.27
	S2-3-90	Py	3.45	0.11
	S2-3-100	Py	2.20	0.11
312–315	S2-5-11	F-Py	−28.49	0.12
	S2-5-16	Py	1.02	0.17
	S2-5-18	Py	1.01	0.18
	S2-5-21	Po	2.09	0.27
444–445	S2-8-1	Py	1.50	0.16
	S2-8-10	Py	2.30	0.17
	S2-8-18	Po	2.01	0.32
	S2-8-20	Py	0.92	0.18
	S2-8-22	Py	3.44	0.14
	S2-8-27	Py	3.29	0.14
	S2-8-27R	Py	3.18	0.13
	S2-8-28	Py	2.63	0.17

Table 6. Sulfur isotope values of sulfide samples from the SOT and other geological settings and δ³⁴S values of other geological bodies. The values in brackets represent the average values

Location	Sample	δ³⁴S_V-CDT/‰	2SE	Data source
SOT	sulfide	0.21 to 3.45		this study
Tangyin	sulfide	−0.75 to 3.15		Yang et al. (2020)
Iheya North Knoll	sulfide	8.8 to 10.5 (9.9±0.2)		Ueno et al. (2003)
Jade	sulfide	5.2 to 7.2 (6.3)		Halbach et al. (1989), Zeng et al. (2002)
Mariana Trough	sulfide	0.8 to 3.5		Ikehata et al. (2015)
North Fiji Basin	sulfide	1.7 to 4.8 (3.4±0.02)		Kim et al. (2006), Zeng et al. (2017b)
Lau Basin-Hine Hina	sulfide	−7.7 to −2.8 (−4.9±0.02)		Herzig et al. (1998)
Manus Basin-PACMANUS	sulfide	−0.8 to 5.6		Kim et al. (2004), McDermott et al. (2015)
Manus Basin-Susu Knolls	sulfide	−8.0 to −3.9		Kim et al. (2004), McDermott et al. (2015)
Mid-ocean Ridges	MORB	0.3	0.5	Sakai et al. (1984)
Mid-ocean Ridges	sulfide	−0.8 to 6.3 (3.2)		Herzig et al. (1998)
Island arc volcanic rocks	andesite, rhyolite	5 to 7		Ueda and Sakai (1984), Woodhead et al. (1987)
Seawater		21	0.2	Kim et al. (2006), Rees et al. (1978)

Fig. 1. Bathymatric map of the Okinawa Trough (OT). a. The locations of the samples and hydrothermal fields in the OT (bathymetric map from http://www.geomapapp.org/index.htm). b. Topography of the Tangyin and Yonaguni Knoll IV HFs and sampling location of core HOBAB4-S2 in the southern Okinawa Trough (SOT) (hydrothermal vent data from Beaulieu et al. (2013) and Tangyin HF data from Zeng et al. (2017a)). The bottom flow directions are from several months of mooring observations (Chung and Hung, 2000).

Fig. 2. Plot showing the core and variations in the Cu, Zn and Pb concentrations.

Fig. 3. SEM image of representative sulfide and sulfate mineral morphologies separated from sediments in core HOBAB4-S2. a. Intergrowth sulfide aggregate is composed of nodular pyrite and minor sphalerite. b. Euhedral pyrite separates with an intergrowing twinning texture. c. Euhedral pyrite separates with a zonal texture. d. Euhedral pyrite with a crystal form of the pentagonal dodecahedron. e. Galena. f. Radiated anhydrite aggregates. Py: pyrite; Sph: sphalerite; Gn: galena; Anh: anhydrite.

Fig. 4. Photomicrographs of typical sulfide mineral assemblages from certain sediment layers in core HOBAB4-S2. The numbers represent δ³⁴S measurements. (a–d) S2-3 separated from 166–167 cm: Intergrowths of pyrite with sphalerite and chalcopyrite (a); Cubic pyrite (b); Framboidal pyrite aggregates (c) and intergrowth of pyrrhotite with arsenopyrite (d). (e, f) S2-5 separated from 312–315 cm: Intergrowths of pyrite with sphalerite and chalcopyrite (e) and Intergrowth of pyrrhotite with subhedral sphalerite (f). (g–i) S2-8 separated from 444–445 cm: Intergrowth of euhedral pyrite with sphalerite (g); Inner chalcopyrite with pyrite on the surface and coexisting sphalerite (h); and Euhedral pyrrhotite (i). Py: pyrite; Cpy: chalcopyrite; Sph: sphalerite; Po: pyrrhotite; Apy: arsenopyrite; F-Py: framboidal pyrite.

Fig. 5. δ³⁴S values of different pyrite and pyrrhotite samples from sediment core HOBAB4-S2 from the SOT. a. S2-3 selected from 166–167 cm; b. S2-5 selected from 312–315 cm; and c. S2-8 selected from 444–445 cm.

Fig. 6. The Fe vs. S and Co vs. Ni diagrams for pyrite. a. Fe vs. S (at%) for pyrite. b. The plot of Co vs. Ni (wt%) for pyrite in core HOBAB4-S2 from the SOT.

Fig. 7. Sulfur isotope values of sulfide samples in hydrothermal sediments from different HFs in the back-arc basin and other geological settings. The δ³⁴S values of mid-ocean ridge basalt (MORB) data are from Sakai et al. (1984). Island arc rock (andesite and rhyolite) data are from Ueda and Sakai (1984) and Woodhead et al. (1987). Modern seawater data are from Rees et al. (1978) and Kim et al. (2006). The sulfur isotopes of sulfides in the Susu Knolls and PACMANUS from the Manus Basin (Kim et al., 2004; McDermott et al., 2015), Lau Basin HFs (Herzig et al., 1998), North Fiji Basin (Kim et al., 2006; Zeng et al., 2017b), Mariana Trough (Ikehata et al., 2015), mid-ocean ridge (sediment-free) (Herzig et al., 1998), Jade HF, NOT (Halbach et al., 1989; Zeng et al., 2002), Iheya North Knoll HF, MOT (Ueno et al., 2003) and Tangyin HF, SOT (Cao et al., 2018; Yang et al., 2020) are shown.

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