Introduction

Mineralogical and geochemical features of seafloor massive sulfide (SMS) deposits at mid-ocean ridge (MOR) spreading centres vary significantly due to the different types of ridge (i.e., fast versus slow spreading) (Hannington et al. 2005; Fouquet et al. 2010). At fast-spreading ridges, hydrothermal fluids only circulate in the upper part of oceanic crust at depths of 1–2 km, which is a region that comprises MOR basalt (MORB) and sheeted dyke, due to the presence of shallow magma chambers (Hannington et al. 2005). In contrast, at slow- to intermediate-spreading ridges, deep-rooted, large-offset detachment faults play an important role in causing amagmatic extension, which allows hydrothermal circulation to occur at much greater depths (~ up to 7 km) and enables fluid interaction with more ultramafic lithologies as compared with fast-spreading systems (McCaig et al. 2007; Escartín et al. 2008). These differences affect the redox state and metal contents of hydrothermal fluids, thereby producing different sulfide mineralogies and contrasts in geochemistry of SMS deposits in MORB- and ultramafic-hosted systems (Hannington et al. 2005; Fouquet et al. 2010; Patten et al. 2016; Knight et al. 2018; Fuchs et al. 2019). Ultramafic-hosted SMS deposits are typically characterised by reduced sulfide assemblages (pyrrhotite–isocubanite–chalcopyrite–Fe-rich sphalerite) and high Cu, Zn, Co, Au, Sn, and Ni contents relative to those of MORB-hosted SMS deposits (Hannington et al. 2005; Fouquet et al. 2010).

Numerous SMS deposits have been discovered along MOR settings since the first discovery of a seafloor hydrothermal venting site at the Galapagos Rift in 1977 (Corliss et al. 1979; Hannington et al. 2011). Although studies have been conducted on the mineralogy and geochemistry of SMS deposits along the Mid-Atlantic Ridge (Marques et al. 2006; Fouquet et al. 2010; Melekestseva et al. 2014; Ren et al. 2021), relatively little is known about SMS deposits in the Central Indian Ridge (CIR). In particular, few studies of the CIR have investigated the hydrothermal processes and genetic environments associated with ultramafic-hosted SMS deposits, such as the Kairei and Cheoeum vent fields (Wang et al. 2014, 2018; Choi et al. 2021). As such, further studies are required to obtain a better understanding of ultramafic-hosted hydrothermal mineralisation in the CIR.

Since 2009, the Korea Institute of Ocean Science and Technology (KIOST) has conducted hydrothermal exploration along the middle part of the CIR (MCIR; 8–17°S; Fig. 1a), which is a slow- to intermediate-spreading ridge (Pak et al. 2017). Eleven oceanic core complexes (OCC) have been recognised in the surveyed areas and generally exhibit hydrothermal plume signatures (Son et al. 2014; Pak et al. 2017). In OCC 1-1 (8.2°S), which shows methane concentrations of up to 13.01 nmol/L and nephelometric turbidity units of up to 0.16 (Pak et al. 2017; Kim et al. 2020), a new hydrothermal site, the Ari vent field (AVF), was discovered by a deep-towed camera during the hydrothermal expedition by R/V ISABU in 2018 (Fig. 1b).

Fig. 1
figure 1

a Tectonic boundaries and distribution of hydrothermal vent fields along the Central Indian Ridge (CIR). The blue box indicates the survey area. b Detailed bathymetric map of segment 1 of the middle part of the CIR. The location of the Ari vent field (AVF) at 8.15°S in oceanic core complex (OCC) 1-1 is marked by a yellow star. The dotted red line indicates the boundaries of the OCC 1-1. Abbreviations: CR, Carlsberg Ridge; MESO, MEteor-SOnne; RTJ, Rodriguez Triple Junction; SEIR, Southeast Indian Ridge; SWIR, Southwest Indian Ridge

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In this study, we conducted a detailed mineralogical investigation and high-resolution geochemical analysis of AVF hydrothermal sulfide samples to characterise the distribution of trace elements and constrain the hydrothermal processes. We compared the geochemical data for pyrite and sphalerite from the AVF with those of other MOR-related SMS deposits and ancient volcanogenic massive sulfide (VMS) deposits in the Urals, in order to distinguish the differences in seafloor hydrothermal mineralisation between mafic- and ultramafic-hosted vent fields. The occurrences of serpentinised ultramafic rocks and reduced sulfide assemblages (pyrrhotite–isocubanite–chalcopyrite–Fe-rich sphalerite) and the distribution of Sn in pyrite and sphalerite indicate that the AVF is one of the few hydrothermal systems in the CIR to have an ultramafic affinity.

Ari vent field

The AVF (8°10.46´S, 68°08.29´E; water depth ~ 3700 m) is located on the OCC 1-1 at the southern inside corner of MCIR segment 1 (Fig. 1b). Its diameter is 150–200 m, as estimated by the deep-towed camera survey (Kim et al. 2020). Basement rocks collected from the OCC 1-1 consist of basalt, gabbro, microgabbro, and serpentinised harzburgite (Yi et al. 2014; Pak et al. 2017). This rock assemblage represents the exhumed lower oceanic crust and mantle, which likely had an important role in determining the redox state and metal contents of the AVF hydrothermal fluids.

The hydrothermal chimneys and mounds are mainly characterised by inactive venting, with diffuse venting being only observed intermittently (Fig. 2). Most chimneys are up to ~ 1.5 m high and, in many cases, are coalesced into a cluster (Fig. 2a). Sulfide mounds without chimney structures are common (Fig. 2b). A thick sediment layer typically covers the surfaces of the inactive chimneys and mounds, where evidence of life was generally absent during the camera survey (Fig. 2a–c). This is in contrast to the small amounts of sediment and abundant vent fauna that are typical of active vent fields in the CIR (Nakamura et al. 2012; Wang et al. 2014). Hydrothermal alteration zones are widespread in the AVF (Fig. 2d). The alteration zone is evident from a reddish brown and/or yellow colour and was likely caused by oxidation of metalliferous sediments by ambient seawater.

Fig. 2
figure 2

Photographs of the Ari vent field. a Inactive chimneys coalesced into a cluster. b Hydrothermal mound covered by thick sediment layers. c, d Hydrothermal alteration zones with a reddish brown and/or yellow colour

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Samples

Hydrothermal sulfide, sulfide-bearing Fe-oxyhydroxide fragments, and consolidated metalliferous sediment samples were recovered using a TV-guided grab (GTV) from the AVF (Fig. 3). The hydrothermal sulfide samples can be classified into two different types according to the major sulfide minerals: (1) Fe–Cu-rich sulfides dominated by pyrite and isocubanite (samples GTV 180101 and 180,103; Fig. 3a, b) and (2) Cu-rich sulfides dominated by isocubanite (samples GTV 180102 and 180106; Fig. 3c, d). Samples GTV 180101 and 180103 have a massive texture, with some cavities being lined by pyrite (Fig. 3a, b). Some greenish fragments of basement rock are included in the matrix of sample GTV 180103 (Fig. 3b). Sample GTV 180102 shows distinct colour zonation (Fig. 3c). The exterior part in contact with seawater is thinly coated with a secondary chalcocite that is dark purple in colour. Sample GTV 180106 is one of the small fragments of Cu-rich sulfides and has a mineralogy similar to that of sample GTV 180102 (Fig. 3d; Table 1). Sample GTV 180202 consists mainly of Fe-oxyhydroxides with minor sulfides (Fig. 3e). X-ray diffraction (XRD) analysis shows peaks of atacamite, hematite, and goethite (ESM 1 Fig. S1). Sample GTV 180402 is consolidated metalliferous sediment that is common around other seafloor hydrothermal vents (Fig. 3f; Hannington et al. 2005). This sediment has various colours, reflecting variable degrees of hydrothermal alteration (Fig. 3f). The sulfide-bearing Fe-oxyhydroxide fragment (GTV 180202) and consolidated metalliferous sediment (GTV180402) samples are from a hydrothermal alteration zone (Fig. 2d).

Fig. 3
figure 3

Photographs of hydrothermal samples collected from the Ari vent field. Hydrothermal sulfides can be classified into two different types according to the major sulfide minerals: a, b Fe–Cu-rich sulfides and c, d Cu-rich sulfides, respectively. e Fe-oxyhydroxide fragment with secondary Cu minerals (white arrows). f Hydrothermally altered, consolidated sediment exhibiting variable degrees of alteration

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Table 1 Mineral abundances of hydrothermal sulfides from the Ari vent field
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Mineralogy and paragenesis

The sulfide samples are classified as Fe–Cu- and Cu-rich, based on the major sulfide minerals (Fig. 4; Table 1). We identified three stages of mineralisation based on the mineral assemblages and textures (Fig. 5).

Fig. 4
figure 4

Photomicrographs and backscattered-electron (BSE) images of sulfide mineral assemblages from the Ari vent field. a–f Fe–Cu-rich sulfide samples: a magnetite (Mgt-A) replaced by isocubanite (Icb-A); b subhedral–euhedral early pyrite (Py-A1) surrounding isocubanite and chalcopyrite (Ccp-A); c early pyrite precipitated along cracks and/or fractures in chlorite (Chl-A); d late colloform pyrite (Py-A2) associated with sphalerite (Sp-A); e galena (Gn-A) inclusion in an early pyrite grain; f uraninite (Urn-A) in early pyrite. g–k Cu-rich sulfide samples: g marcasite (Mrc-B) replaced by coarse-grained isocubanite–chalcopyrite (Icb-B–Ccp-B) aggregates; h relics of early sphalerite (Sp-B1) replaced by isocubanite–chalcopyrite aggregates; i late sphalerite (Sp-B2) replacing isocubanite–chalcopyrite aggregates; j cobaltite (Cbt-B) and electrum (El-B) in cavities of Cu sulfides; k chalcocite (Cct-B) extensively replacing earlier formed minerals. l Sulfide-bearing Fe-oxyhydroxide fragment with atacamite (Atc-C) surrounding pyrite (Py-C) replaced by goethite (Gth-C) and hematite (Hem-C). Abbreviations: Aip, altered isocubanite phase; Po, pyrrhotite; “A”, “B”, and “C” indicate minerals in the Fe–Cu-rich sulfides, Cu-rich sulfides, and sulfide-bearing Fe-oxyhydroxide fragment, respectively. The numbers indicate the generations of pyrite and sphalerite inferred from the textures and mineral assemblages

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Fig. 5
figure 5

Paragenesis of hydrothermal sulfides in the Ari vent field. Abbreviations are as in Fig. 4

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Fe–Cu-rich sulfide samples

Magnetite (Mgt-A), pyrrhotite (Po-A), isocubanite (Icb-A), and chalcopyrite (Ccp-A) are early-formed minerals of stage I mineralisation (Figs. 4a–c and 5). Magnetite and pyrrhotite are commonly replaced by isocubanite (Fig. 4a, b). Chalcopyrite occurs mainly as exsolution lamellae within isocubanite and also as a few discrete grains (Fig. 4b). With progressive mineralisation, the pyrite increases in content and has two morphologically and mineralogically distinct generations (Fig. 4a–f). Subhedral to euhedral early pyrite (Py-A1) surrounds isocubanite and chalcopyrite, indicating that pyrite precipitated after Cu-sulfides (Fig. 4a, b). The early pyrite is often precipitated along fracture zones in chlorite grains, which are relics of basement rocks (Fig. 4c). Stage II mineralisation is represented by late colloform pyrite (Py-A2), sphalerite (Sp-A), and galena (Gn-A) (Fig. 4d, e). The late pyrite forms a lining structure associated with sphalerite (Fig. 4d). Galena fills some cavities in stage I sulfide minerals (Fig. 4e). Stage III is characterised by Fe–Cu-rich sulfide samples that have experienced seawater alteration. In particular, trace amounts of uraninite (< 1 μm in size; Urn-A) sporadically infill some cavities and/or are precipitated on altered surfaces of early pyrite during this mineralisation stage (Figs. 4f; ESM 1 S2a).

Cu-rich sulfide samples

Stage I mineralisation comprises mainly isocubanite (Icb-B) and sphalerite, along with minor pyrrhotite (Po-B), chalcopyrite (Ccp-B), and marcasite (Mrc-B) (Fig. 5; Table 1). Early-formed marcasite shows altered surfaces (Fig. 4g). Sphalerite is more abundant in the Cu-rich sulfide samples than in the Fe–Cu-rich sulfide samples (Table 1). In the former samples, sphalerite shows two different generations. Relics of early sphalerite (Sp-B1) containing numerous inclusions of isocubanite are rarely identified and are replaced by isocubanite (Fig. 4h). With progressive mineralisation, coarse-grained isocubanite becomes predominant, and chalcopyrite appears as exsolution lamellae within isocubanite (Fig. 4g–k). Late sphalerite (Sp-B2) commonly replaces isocubanite with chalcopyrite exsolution (Fig. 4i). Stage II mineralisation is characterised by trace amounts of cobaltite (Cbt-B) precipitated in cavities of isocubanite grains, which is sub- to euhedral (Fig. 4j; ESM 1 S2b). Electrum occurs as small inclusions (El-B; mostly < 1 μm) in cobaltite or infills the cavities in earlier-formed sulfides (Fig. 4j; ESM 1 S2b). Stage III mineralisation is characterised by chalcocite (Cct-B) and an altered isocubanite phase (Aip-B), which extensively replace earlier-formed isocubanite (Fig. 4g, k).

Sulfide-bearing Fe-oxyhydroxide fragment

The sulfide-bearing Fe-oxyhydroxide fragment consists mainly of atacamite (Atc-C), hematite (Hem-C), and goethite (Gth-C), along with trace pyrite (Py-C), isocubanite (Icb-C), chalcopyrite (Ccp-C), galena (Gn-C), and uraninite (Urn-C; Table 1; ESM 1 Fig. S1). Pyrite is replaced by hematite and/or goethite, whereas the other sulfides are present as submicroscopic inclusions in the Fe-oxyhydroxides (Fig. 4l). Uraninite is present as inclusions (< 1 μm in size) in hematite (Fig. 4l inset). All of these minerals are enclosed by later-formed atacamite (Fig. 4l).

Analytical methods

An optical microscope and XRD analysis were used for mineral identification and textural interpretation of the hydrothermal sulfide samples and sulfide-bearing Fe-oxyhydroxide fragment, at KIOST, Busan, South Korea. The semiquantitative analyses of the mineralogy of 19 polished sections are presented in Table 1. XRD analysis was undertaken using a Panalytical X’Pert-PRO diffractometer with a CuKα X-ray source operated at 40 kV and 30 mA. The XRD patterns were recorded over a 2θ range from 5 to 65°, with a 0.01° step size and scan rate of 1°/min (ESM 1 Fig. S1).

Bulk chemical compositions of the hydrothermal sulfides were determined using Au–Ag Fire Assay, 4-Acid Digestion (Code 8 ICP–OES), and Peroxide Fusion Package (Ultratrace 7) at Actlabs (Ancaster, Ontario, Canada). The detection limits for each element are listed in Table 2.

Table 2 Bulk chemical compositions of hydrothermal sulfides from the Ari vent field. All data are individual analyses
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Electron microprobe analysis (EPMA) of individual minerals was conducted using a JEOL JXA-8530F electron microprobe with an accelerating voltage of 15 kV, a beam current of 20 nA, and an electron beam diameter of 5 μm at Gyeongsang National University, Jinju, South Korea. Natural mineral and synthetic standards and Aztec software using ZAF corrections were used for the data calibration: FeS2 (for Fe and S), ZnS (Zn), CuFeS2 (Cu), PbS (Pb), CdS (Cd), Sb2S3 (Sb), InAs (In and As), and pure metal (Mn, Co, and Ni). Results of individual analyses are given in ESM 2 Table S1.

Laser ablation–inductively coupled plasma–mass spectrometry (LA–ICP–MS) analysis was undertaken with a 193-nm excimer LA system (ESI NWR 193, USA) coupled to an Agilent 7700 quadrupole ICP–MS instrument at KIOST. The laser beam diameter was 30–50 μm, depending on mineral grain size, the laser pulse rate was 10 Hz, and the laser energy was 5.7 J/cm2. The total analysis time for each spot was 90 s, comprising 50 s of background measurement followed by 40 s of data acquisition during sample ablation. The following isotopes were measured: 55Mn, 57Fe, 59Co, 60Ni, 65Cu, 66Zn, 69 Ga, 74Ge, 75As, 77Se, 95Mo, 109Ag, 111Cd, 115In, 118Sn, 121Sb, 125Te, 197Au, 205Tl, 208Pb, 209Bi, and 238U. Dwell times for each element were set to 0.02 s, except for Cu, Fe, and Zn, which were set to 0.01 s. External calibration was undertaken using STDGL3 (Belousov et al. 2014). The MASS-1 sulfide reference material (also known as PS-1; Wilson et al. 2002) was analysed as an unknown sample to assess the data quality (ESM 2 Table S2). The results yielded a relative standard deviation (RSD) of < 6% for most elements. The Fe, Zn, and Cu contents determined by EPMA were used as internal standards for quantification of pyrite, sphalerite, and isocubanite, respectively. Contents of Ga and Hg were calculated using MASS-1 as a primary standard, because the Ga and Hg contents of STDGL3 are poorly constrained. The LA profiles for each element were monitored to identify the presence of micron-sized mineral inclusions. Spectra with spikes were not used to calculate trace-element contents. Data calculations were carried out using an in-house Excel spreadsheet and following the method described by Longerich et al. (1996). The entire dataset is presented in ESM 1 Fig. S3 and ESM 2 Tables S3–5.

LA–ICP–MS elemental mapping was undertaken by ablating sets of parallel lines in a grid across each sample. Lines were ablated with a beam size of 9 μm. The spacing between the lines and scan speed was kept constant to match the laser spot size. A laser frequency of 10 Hz was used at a constant laser energy of 5.7 J/cm2. The acquisition time for most elements was set to 0.02 s, but for major elements (Fe, Cu, and Zn), it was 0.01 s. Images were compiled and processed using Iolite software developed by WaveMetrics (Paton et al. 2011).

In situ sulfur isotope analyses of pyrite were conducted with a Neptune Plus multiple collector–ICP–MS (Thermo Fisher Scientific, Bremen, Germany) coupled to a 193-nm GeoLas HD excimer ArF LA system (Coherent, Göttingen, Germany) at the Wuhan Sample Solution Analytical Technology Company Limited, Hubei, China. Helium gas was used to transport the ablated materials into the plasma with a gas flow of 0.5 L/min. Ablation was performed with a laser beam diameter of 44 μm, laser pulse rate of 2 Hz for single spot analyses, and laser energy of 6 J/cm2. To avoid matrix effects, a pyrite standard PPP-1 (Fu et al. 2016) was used as a reference material for correcting the pyrite data. In addition, the in-house reference materials pyrrhotite SP-Po-01 (δ34SVCDT = 1.4‰ ± 0.4‰) and pyrite SP-Py-01 (δ34SVCDT = 2.0‰ ± 0.5‰) were analysed repeatedly as unknowns to assess the data quality. The standard errors for PPP-1, SP-Po-01, and SP-Py-01 are ± 0.08‰, ± 0.08‰, and ± 0.18‰ (2 SD), respectively.

Results

Bulk chemistry

The hydrothermal sulfide samples have Cu contents (1.6–33 wt%) that are much higher than those of Zn (0.01–5.67 wt%) and Pb (0.0008–0.025 wt%; Table 2). These data plot within the sediment-free MOR field, similar to those of other SMS deposits in the CIR (Fig. 6a). Cobalt, Ga, Se, In, and Sn are more concentrated in the Cu-rich sulfide samples as compared with the Fe–Cu-rich sulfide samples, whereas Ni, U, and Mo are more enriched in the latter (Fig. 6b–d; Table 2). The Co contents exhibit a strong positive correlation with Se contents (R2Co–Se = 0.96; Fig. 6b). The U contents are positively correlated with Mo contents (R2U–Mo = 0.76; Fig. 6c) but negatively correlated with Sn contents (R2U–Sn = 0.75; Fig. 6d).

Fig. 6
figure 6

Bulk chemical compositions of hydrothermal sulfides in the AVF. (a) Cu–Zn–Pb ternary diagram modified after Fouquet et al. (1993). Log–log plots of (b) Se versus Co, (c) Mo versus U, and (d) Sn versus U. (e and f) Detailed comparison of the Ari vent field with other SMS deposits from mid-ocean ridges: (e) Sn versus Cu + Zn and (f) Sn versus Fe. Average compositions of sulfides were taken from Hannington et al. (2005), Fouquet et al. (2010), Wang et al. (2014), Cao et al. (2018), Grant et al. (2018), Meng et al. (2020), and Choi et al. (2021). Abbreviations: EPR, East Pacific Rise; CIR, Central Indian Ridge; MAR, Mid-Atlantic Ridge; MESO, MEteor-SOnne; MORB, mid-ocean ridge basalt

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A comparison with other MOR systems shows that high (Cu + Zn) and Sn contents are distinctive characteristics of ultramafic-hosted sulfides (Fig. 6e). The AVF hydrothermal sulfides are relatively poor in Sn compared to other ultramafic-hosted sulfides from the Mid Atlantic Ridge, but the Sn content is distinct between Cu-rich and Fe–Cu-rich sulfide samples (Fig. 6e, f). Copper-rich sulfide samples have an affinity with ultramafic-hosted systems, whereas Fe–Cu-rich sulfide samples are typical of MORB-hosted systems (Fig. 6e). Although Fe contents exhibit no systematic differences between these two types of hydrothermal systems, ultramafic-hosted sulfides, including the AVF sulfide samples, are characterised by a negative correlation between Fe and Sn contents (R2Fe–Sn = 0.75; Fig. 6f).

Chemical compositions of sulfide minerals

Pyrite

Trace element contents of pyrite were only obtained for the Fe–Cu-rich sulfide samples, because the highly altered marcasite in the Cu-rich sulfide samples produced irregular LA–ICP–MS spectra (Figs. 4g; ESM 1 S4a). Most pyrite in the Fe–Cu-rich sulfide samples has smooth LA–ICP–MS time-resolved elemental profiles, but some exhibit irregular spikes of U (ESM 1 Fig. S4b). Cobalt, Ni, Cu, Se, and Sn are more concentrated in early pyrite (Py-A1) as compared with late pyrite (Py-A2), whereas Mn and Tl are more enriched in the latter (ESM 1 Fig. S3). The Co contents generally increase with increasing Te, Se, and Ni, but decrease with increasing Mn and Tl (Fig. 7a–d; ESM 2 Table S3). Some data for Py-A1, which exhibit substantial depletion in Ni at a given Co content (Fig. 7c), also have relatively high Mn and Tl contents comparable to those of Py-A2 (Fig. 7d).

Fig. 7
figure 7

Trace element contents of pyrite determined by LA–ICP–MS. (a) Te versus Co, (b) Se versus Co, (c) Ni versus Co, and (d) Tl versus Mn. The dotted black lines indicate the below detection limit (bdl) of analysis. Abbreviations are as in Fig. 4

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Sphalerite

The AVF samples are dominated by Fe-rich sphalerite, which shows no systematic variation in FeS contents between the Fe–Cu-rich sulfide samples (average = 29 ± 1.5 mol %) and Cu-rich sulfide samples (30 ± 3 mol %; ESM 2 Table S1). In contrast, the trace element contents of the AVF sphalerite are highly variable in the two different types of hydrothermal sulfide samples (ESM 1 Fig. S3), although LA–ICP–MS analysis of the early sphalerite (Sp-B1) in the Cu-rich sulfide samples could not be undertaken due to the large amounts of mineral inclusions (Fig. 4h). Late sphalerite (Sp-B2) in the Cu-rich sulfide samples contains more Co, Ge, As, Se, Ag, Hg, Pb, and Bi as compared with sphalerite (Sp-A) from the Fe–Cu-rich sulfide samples, whereas Mn and Sn are more enriched in the latter (ESM 1 Fig. S3 and ESM 2 Table S4).

The Se contents exhibit a strong positive correlation with Co (R2Se–Co = 0.98 for Sp-A and 0.63 for Sp-B2; Fig. 8a). The Sn contents differ between the Fe–Cu- and Cu-rich sulfide samples: (1) Sn has positive and negative correlations with Se contents in the Fe–Cu- and Cu-rich sulfide samples, respectively (Fig. 8b); (2) the Cu/Sn ratio is very close to ~ 2 in the Fe–Cu-rich sulfide samples, but the relationship is more variable at relatively low Sn contents in the Cu-rich sulfide samples (Fig. 8c); and (3) most individual analyses of the Cu-rich sulfide samples lie on, or close to, the Ga:Sn = 1:1 line, but all data for the Fe–Cu-rich sulfide samples deviate from this line (Fig. 8d).

Fig. 8
figure 8

Trace element contents of sphalerite determined by LA–ICP–MS. (a) Co versus Se, (b) Sn versus Se, (c) Sn versus Cu, and (d) Sn versus Ga. The solid black lines indicate data correlation trends. Abbreviations are as in Fig. 4

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Isocubanite

The Co, Ga, Se, Ag, and In contents of isocubanite (Icb-B) in the Cu-rich sulfide samples are higher than those of isocubanite (Icb-A) in the Fe–Cu-rich sulfide samples, whereas Mn is more enriched in the latter (ESM 1 Fig. S3). The Se contents commonly increases with Co contents (Fig. 9a). In particular, Icb-B is characterised by systematic variations in Zn, Ga, Se, and Sn contents that differ from those of Icb-A (Fig. 9b–d). Specifically, Sn contents are negatively correlated with Se contents (Fig. 9b) but positively correlated with Ga and Zn contents for Icb-B (Fig. 9c, d).

Fig. 9
figure 9

Trace element contents of isocubanite determined by LA–ICP–MS. (a) Se versus Co, (b) Se versus Sn, (c) Sn versus Ga, and (d) Zn versus Ga. Abbreviations are as in Fig. 4

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LA–ICP–MS elemental mapping

Elemental maps were obtained for a Cu-rich sulfide sample to investigate the distribution of trace elements between adjacent minerals (ESM 1 Fig. S5). The maps show that Co, As, Ag, and Pb are incorporated preferentially into marcasite as compared with late sphalerite (Sp-B2) and isocubanite with chalcopyrite exsolution. In particular, the Ga and Sn contents appear to be zoned. The highest contents are confined to replacement boundaries between the late sphalerite and isocubanite with chalcopyrite exsolution.

Sulfur isotopic composition of pyrite

In situ S isotopic compositions (n = 10) of pyrite were obtained from Fe–Cu-rich sulfide samples in accordance with the different mineralisation stages (Fig. 5; Table 3). Early pyrite (Py-A1) has δ34S = 6.2–8.5‰ (average = 7.03‰), whereas late pyrite (Py-A2) has δ34S = 6.6–6.7‰ (average = 6.65‰; Table 3). The data overlap those of other MOR systems (Fig. 10).

Table 3 In situ S isotopic compositions of pyrite in the Fe–Cu-rich sulfide samples
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Fig. 10
figure 10

Sulfur isotope composition of different generations of pyrite in the Ari vent field. Ranges of δ34S values of other MOR systems are modified from Zeng et al. (2017). Other data are from the following: Ding et al. (2021); Tianzuo (Cao et al. 2021); Yuhuang-1 (Liao et al. 2018); seawater (Rees et al. 1978); MORBs (Sakai et al. 1984); gabbro (Alt et al. 1989, 2007; Alt and Anderson 1991

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Discussion

Mineralisation sequence and fluid evolution

The typical exterior–interior mineralogical zones and innermost vent conduits of seafloor chimneys are not observed in the hydrothermal sulfide samples (Fig. 3). In addition, the matrix of sample GTV 180103 contains some greenish fragments of basement rock (Fig. 3b). These results indicate that the collected samples correspond to massive sulfide mounds taken from slightly different locations. Petrographic investigations reveal that Fe–Cu-rich sulfide samples underwent three different stages of mineralisation with decreasing fluid temperature and ƒS2 and increasing ƒO2 (Figs. 4 and 5): subhedral–euhedral pyrite (Py-A1) and isocubanite (Icb-A) dominates stage I; colloform pyrite (Py-A2) and sphalerite (Sp-A) dominates stage II; and stage III represents seawater alteration. The Cu-rich sulfide samples have mineral assemblages and a paragenesis similar to those of the Fe–Cu-rich sulfide samples, but the much higher amount of isocubanite indicates relatively reducing and high-temperature conditions during deposition of the former (Fig. 5; Table 1; Kawasumi and Chiba 2017). This is consistent with LA–ICP–MS analysis showing that the Co and Se contents of sphalerite and isocubanite are higher in the Cu-rich sulfides than in the Fe–Cu-rich sulfides (Figs. 8a and 9a; ESM 2 Tables S4 and 5), given that enrichments in these elements are typical of relatively high-temperature sulfide minerals because the solubility of Co and Se in vent fluids decreases abruptly at temperatures of < 350 °C (Huston et al. 1995; Butler and Nesbitt 1999; Metz and Trefry 2000; Maslennikov et al. 2009; Keith et al. 2016; Meng et al. 2020).

LA–ICP–MS analyses show that the trace element contents of pyrite vary in the different mineralisation stages (Fig. 7; ESM 1 S3). Cobalt, Se, and Ni are enriched in early pyrite (Py-A1) as compared with late pyrite (Py-A2), whereas Mn and Tl are more enriched in the latter (Fig. 7; ESM 2 Table S3). Previous studies have suggested that high contents of Mn and Tl are good indicators of low-temperature mineralisation (< 200 °C; Maslennikov et al. 2009; Meng et al. 2020). As such, the Co–Se-rich early pyrite from the AVF was precipitated from relatively high-temperature fluids as compared with Mn–Tl-rich late pyrite (Fig. 7). However, the temperature dependency cannot explain the enrichment of Ni in early pyrite relative to late pyrite (ESM 2 Table S3), as Ni is typically incorporated into the crystal lattice of pyrite during relatively low-temperature mineralisation (Maslennikov et al. 2009; Keith et al. 2016). The Ni contents of the early pyrite are highly variable at a given Co content (Fig. 7c), suggesting that the fluid temperature was not a major control on the Ni contents of the AVF pyrite. Alternatively, a high ƒS2 of hydrothermal fluids is known to enhance the incorporation of Ni into pyrite (Maslennikov et al. 2009; Li et al. 2017). We suggest that the main-stage mineralisation corresponding to stage I was associated with high ƒS2, thereby enhancing the substitution of Ni into early pyrite (Figs. 5 and 7c). As mineralisation proceeded, the influx of ambient seawater may have decreased the ƒS2 and temperature of the hydrothermal fluids, which ultimately led to the relatively Ni-poor compositions of some early pyrite (Fig. 7c). This is supported by the Ni-poor early pyrite that has relatively high Mn and Tl contents similar to those of late pyrite (Fig. 7d). These results indicate that the Ni contents of the AVF pyrite were likely controlled by ƒS2 rather than the fluid temperature.

Seafloor hydrothermal deposits with an ultramafic affinity are typically characterised by CH4- and H2-rich and H2S-poor hydrothermal fluids as compared with MORB-hosted SMS deposits (Charlou et al. 2002; Nakamura et al. 2009; Schmidt et al. 2011), indicative of low ƒO2–ƒS2 environments. Such a low redox potential and ƒS2 of the AVF hydrothermal fluids is consistent with other lines of geochemical and mineralogical evidence. The AVF sphalerite has high FeS contents (26.6–36.5 mol %), irrespective of the two types of hydrothermal sulfide samples (ESM 2 Table S1). These values are higher than those of many MORB-hosted systems (mostly < 25 mol % FeS; Graham et al. 1988; Hannington et al. 1991; Kawasumi and Chiba 2017), indicating that low ƒO2–ƒS2 conditions facilitated the incorporation of Fe into the crystal lattice of the AVF sphalerite (Scott and Barnes 1971; Kawasumi and Chiba 2017). The AVF sulfide samples have a mineral assemblage of pyrrhotite–isocubanite–chalcopyrite–Fe-rich sphalerite, which is common for other ultramafic-hosted SMS deposits in MOR settings (Fig. 5; Fouquet et al. 2010; Melekestseva et al. 2014; Wang et al. 2014; Choi et al. 2021). A previous experimental study showed that isocubanite began to form, intergrown with chalcopyrite and pyrrhotite, at 335 °C (Lusk and Bray 2002). This is consistent with isocubanite thermometry, which yielded an average formation temperature of ~ 365 °C for the ultramafic-hosted Cheoeum vent field, CIR (Choi et al. 2021). As such, the mineral assemblage in the AVF is indicative of highly reducing conditions and a formation temperature of > 335 °C. In addition, the magnetite replaced by isocubanite appears to be a high-temperature mineral of primary origin during mineralisation stage I (Fig. 4a). A previous study suggested that very low ƒO2–ƒS2 fluid conditions and low H2S contents allow magnetite to precipitate in Cu–Fe-rich submarine hydrothermal chimneys (Fouquet et al. 2010).

The redox state of the AVF hydrothermal fluids varied significantly between different samples and mineralisation stages. For example, enrichments of Co and Se in sphalerite and isocubanite from the Cu-rich sulfide relative to the Fe–Cu-rich sulfide samples indicate that the former samples were formed under more reducing, high-temperature mineralisation, given that these elements are typical of sulfide minerals precipitated from such conditions (ESM 2 Tables S4 and 5; Huston et al. 1995; Butler and Nesbitt 1999; Maslennikov et al. 2009; Keith et al. 2016; Meng et al. 2020; Choi et al. 2023). In particular, substantial enrichment of Te in early pyrite (0.15 − 15.3 ppm) relative to late pyrite (mostly below detection limits) suggests that relatively reducing fluids produced the early pyrite (Figs. 7a; ESM 1 S3), given that a significant decrease in Te solubility can be caused by low ƒO2 conditions (Grundler et al. 2013).

Our results indicate that the AVF sulfide samples were mainly formed by reducing, high-temperature fluids associated with an ultramafic-hosted hydrothermal system. This is consistent with the fact that serpentinisation of ultramafic rocks produces H2- and CH4-rich fluids, resulting in highly reducing conditions (Charlou et al. 2002; Nakamura et al. 2009; Schmidt et al. 2011). The sulfide samples are characterised by three different temporal variations in sulfide minerals with decreasing fluid temperature and ƒS2 and increasing ƒO2 from the main mineralisation stage I (> 335 °C) to relatively low-temperature mineralisation stage II (< 200 °C) and seawater alteration stage III (Fig. 5). This variable mineralisation is readily achieved by mixing between the reducing, high-temperature fluids, and ambient oxidised seawater as mineralisation progressed. However, compared with other ultramafic-hosted sulfides, the AVF sulfide samples have significant differences in U and Sn contents as described below (Fig. 6c–f; Table 2; Fouquet et al. 2010).

Uranium mineralisation

The Fe–Cu-rich sulfide samples and sulfide-bearing Fe-oxyhydroxide fragment contain discrete uraninite inclusions (Fig. 4f, l; ESM 1 S2a). In particular, the Fe–Cu-rich sulfide samples are enriched in Mo (80–225 ppm) and U (7.1–51.9 ppm) as compared with the Cu-rich sulfide samples (Fig. 6c; Table 2). These characteristics suggest that ambient seawater could be a principal source of elevated U content in the AVF, given that submarine hydrothermal fluids are substantially depleted in Mo (mostly < 10 nM) relative to seawater (average = 104 nM; Douville et al. 2002). We suggest that the weak hydrothermal activity (i.e., the predominance of inactive venting; Fig. 2) in the AVF had an important role in the formation of gossan-like altered zones on the seafloor (Maslennikov et al. 2012; Ayupova et al. 2018), which may have increased the U contribution of seawater. It is also considered that seawater circulating through the oceanic crust extracts U. As such, the pristine fluids expelled at the seafloor are U-poor, thereby forming sulfide minerals that are depleted in U (Hegner and Tatsumoto 1989; Mills et al. 1994; Butler and Nesbitt 1999). This is consistent with our LA–ICP–MS analyses, which showed that the AVF sulfide minerals are mostly depleted in U (< 0.5 ppm; ESM 2 Tables S3–5). Although some analyses of early pyrite (Py-A1) in the Fe–Cu-rich sulfide samples are characterised by anomalously high U contents of up to 12.2 ppm (ESM 2 Table S3), the irregular spikes of U in the LA–ICP–MS depth profiles reflect the presence of U-bearing inclusions within Py-A1 (ESM 1 Fig. S4b). Therefore, we suggest that the U contents of the AVF were mainly controlled by the precipitation of uraninite. The uraninite is mainly deposited on the altered surfaces of pyrite and/or hematite (Fig. 4f, l; ESM 1 S2a). Given the thick sediment layers, widespread hydrothermal alteration zones, and abundance of atacamite, chalcocite, hematite, and goethite around the inactive chimneys and/or mounds, protracted submarine weathering occurred in the AVF (Fig. 2; Table 1). As such, the oxidative alteration of Fe-bearing minerals may have facilitated the reduction of U from the hexavalent to tetravalent state, thereby enabling precipitation of uraninite inclusions (Fig. 4f, l; ESM 1 S2a). This is consistent with previous studies that suggested the fixation of seawater-derived U can be induced by the oxidation of Fe minerals (Mills et al. 1994; Ayupova et al. 2018).

Tin mineralisation

In the AVF sulfide samples, sphalerite (average 519 ± 524 ppm Sn; up to 2386 ppm Sn) and, to some extent, isocubanite (average 54.3 ± 135 ppm Sn; up to 939 ppm Sn) are substantially enriched in Sn as compared with pyrite (up to 16.2 ppm Sn; ESM 1 Fig. S3 and ESM 2 Tables S3–5), indicating that sphalerite and isocubanite are the main carriers of Sn in the AVF. This is consistent with the fact that Sn-rich, ultramafic-hosted SMS deposits are characterised by high Cu and Zn contents as compared with Sn-poor MORB-hosted sulfide deposits (Fig. 6e).

The Sn contents of the AVF sphalerite exhibit positive and negative correlations with Se in the Fe–Cu-rich (Sp-A) and Cu-rich sulfide samples (Sp-B2), respectively (Fig. 8b). As such, the sphalerite Sn contents cannot be explained by fluid temperature, as Se enrichments are typical of high-temperature sulfide minerals (Huston et al. 1995; Butler and Nesbitt 1999; Maslennikov et al. 2009; Meng et al. 2020; Choi et al. 2023). With some exceptions, Sn, Cu, and Ga contents are positively correlated with each other in the AVF sphalerite (Fig. 8c, d). In particular, the values of Cu/Sn = 2 and Ga/Sn = 1 reflect the control of sphalerite Sn contents being due to the coupled substitutions 3Zn2+  ↔ 2Cu+  + Sn4+ and 3Zn2+  ↔ Cu+  + Sn2+  + Ga3+, respectively (Cook et al. 2009; Ye et al. 2011). This suggests that determining the oxidation state of Sn (i.e., divalent versus tetravalent) is important for constraining the possible controls on sphalerite Sn contents due to lattice substitution. However, the Sn contents of the AVF sphalerite differ between the Fe–Cu- and Cu-rich sulfide samples. For the former, most data for Sp-A have Cu/Sn ~ 2, whereas most data for Sp-B2 from the Cu-rich sulfide samples have Cu/Sn < 2, especially at relatively low Sn contents (Fig. 8c). This suggests that the incorporation of Sn4+ into sphalerite may be facilitated by relatively oxidising, low-temperature conditions, given that Sp-A is found together with colloform pyrite in the late mineralisation stage II (Figs. 4d and 5), whereas Sp-B2 is precipitated with coarse-grained isocubanite in the main mineralisation stage I (Figs. 4h and 5). In contrast, the positive correlation between Ga and Sn contents with Ga/Sn = 1 is limited to Sp-B2 (Fig. 8d). This suggests that the preferential substitution of 3Zn2+  ↔ Cu+  + Sn2+  + Ga3+ occurs under relatively reducing, high-temperature conditions. These results suggest that the redox state of hydrothermal fluids is an important control on the Sn contents of sphalerite. Considering the much higher Sn contents of Sp-A relative to Sp-B2 (ESM 2 Table S4), Sn4+ is likely the most important form involved in the generation of Sn-rich sphalerite.

Systematic variations in Sn contents are only observed in isocubanite (Icb-B) in the Cu-rich sulfide samples and not in the isocubanite (Icb-A) in the Fe–Cu-rich sulfide samples (Fig. 9b–d). The Sn contents are negatively correlated with Se contents in Icb-B (Fig. 9b), although these elements are typical of relatively high-temperature Cu sulfide minerals (Hutchison and Scott 1981; Huston et al. 1995; Maslennikov et al. 2009). This indicates that fluid temperature had little effect on Sn contents in the AVF isocubanite. In contrast, Sn, Ga, and Zn contents in Icb-B are positively correlated with each other (Fig. 9c, d). The LA–ICP–MS depth profiles are typically flat, suggesting that these elements are present in Icb-B in solid solution. In particular, relics of early sphalerite (Sp-B1) occur within Icb-B (Fig. 4h). This indicates that the early formed Sp-B1 may have been dissolved and re-precipitated by the continuously ascending hydrothermal fluids. The LA–ICP–MS elemental maps also show that Ga and Sn are concentrated in bands along replacement boundaries between Sp-B2 and Icb-B in the Cu-rich sulfide sample (ESM 1 Figs. S5 and 6). Choi et al. (2021) suggested that these Sn–Ga-rich bands formed because these elements were no longer incorporated into the Cu minerals via coupled dissolution and re-precipitation. Although it cannot be completely excluded that the fluids that precipitated Icb-B were initially enriched in Sn and Ga, earlier formed sphalerite may have been one of the sources of Sn for the subsequent remobilisation process, given that the Sn and Ga contents of Icb-B are positively correlated with Zn (Fig. 9c, d). Our results suggest that the redox state of hydrothermal fluids and/or coupled dissolution and reprecipitation of previously deposited Sn-bearing sulfides could be more important factors controlling the Sn content compared to the fluid temperature.

Although the AVF sphalerite and isocubanite are enriched in Sn (ESM 1 Fig. S3), bulk chemical compositions show that the Sn contents differ for the two different types of AVF sulfide samples (Fig. 6d–f). High Sn contents, comparable to those of other ultramafic-hosted sulfides, are confined to the Cu-rich sulfide samples, whereas the Fe–Cu-rich sulfide samples are characterised by an affinity with MORB-hosted sulfides due to the significant Sn depletion (Fig. 6e). This indicates that further explanation is required to account for the anomalous Sn distribution in the Fe–Cu-rich sulfide samples. We suggest that ultramafic-hosted SMS deposits are likely to be depleted in Sn if they are dominated by Fe-rich mineralisation, given that the Fe–Cu-rich sulfide samples in the AVF consist mainly of Sn-poor pyrite (mostly < 1 ppm) (Table 1; ESM 1 Fig. S3). This is consistent with the bulk chemical compositions of other ultramafic-hosted SMS deposits at MOR settings, which exhibit a negative correlation between Sn and Fe contents (Fig. 6f).

Distribution of Sn in pyrite and sphalerite: a comparison of hydrothermal sulfides at MOR settings

In MOR-related hydrothermal systems, one of the most pronounced differences is the much higher Sn contents of ultramafic-hosted SMS deposits relative to MORB-hosted sulfide deposits (Fouquet et al. 2010; Wang et al. 2014; Evrard et al. 2015; Choi et al. 2021). However, it is still unclear why Sn enrichment is associated primarily with ultramafic-hosted hydrothermal mineralisation and which mineral(s) is the main Sn host.

To better understand the distribution of Sn on the mineral scale, we undertook a comparison of pyrite and sphalerite from different types of hydrothermal vent fields (i.e., MORB- versus ultramafic-hosted) in MOR settings (Fig. 11), given that pyrite and sphalerite are major constituents of SMS deposits and can incorporate various trace elements (Maslennikov et al. 2009; Keith et al. 2016; Meng et al. 2020). Sphalerite is substantially enriched in Sn (average > 1000 ppm), whereas most pyrite has very low average Sn contents (< 1 ppm) (Fig. 11). This indicates that sphalerite is one of the main Sn hosts, whereas pyrite is Sn-poor in ultramafic-hosted SMS deposits. Most of the ultramafic-hosted pyrite and sphalerite are enriched in Sn as compared with those in MORB-hosted deposits (Fig. 11). This difference suggests that hydrothermal fluids circulating through ultramafic lithologies could be a more efficient source of Sn as compared with MORB-related hydrothermal fluids. We suggest that the low redox potential of hydrothermal fluids in ultramafic-hosted systems could be important in enhancing the transport of SnCl2 (Sn2+) during hydrothermal circulation. As such, hydrothermal fluids are likely to precipitate Sn-rich minerals in ultramafic-hosted systems. This is consistent with the study of Schmidt et al. (2011) that reported the Sn concentrations of hydrothermal fluids are two orders of magnitude higher in the ultramafic-hosted Nibelungen vent field as compared with the MORB-hosted Red Lion vent field. Therefore, the Sn contents of the AVF pyrite and sphalerite, which are within the range of other ultramafic-hosted systems (Fig. 11), indicate that an ultramafic-hosted hydrothermal system had an important role in forming the AVF.

Fig. 11
figure 11

Average contents of Sn and Se in (a) pyrite and (b) sphalerite for MORB- and ultramafic-hosted vent fields at MOR settings and mafic- and ultramafic-hosted VMS deposits in the Urals. All data were determined by LA–ICP–MS. Average compositions of sulfide minerals are from Maslennikov et al. (2017, 2020), Wang et al. (2017, 2018), Grant et al. (2018), Melekestseva et al. (2018, 2020a, b), Yuan et al. (2018), Meng et al. (2020), Choi et al. (2021), Liao et al. (2021), Ren et al. (2021) and Ding et al. (2022). Contents in parentheses indicate the different types of chimneys in the same vent field

Full size image

Our comparison also shows that Se contents of pyrite and sphalerite differ in the different types of hydrothermal chimneys (i.e., Cu- versus Zn-rich) in the same vent field, although they are not distinguishable between MORB- and ultramafic-hosted vent fields (Fig. 11). For example, pyrite and sphalerite from the Snake Pit and Rainbow sites have elevated Se contents in Cu-rich rather than Zn-rich chimneys (Fig. 11). Huston et al. (1995) demonstrated that mixing of fluids with seawater substantially lowers the Se contents of pyrite. In addition, high Se contents of sulfide minerals are commonly related to relatively reducing, high-temperature mineralisation (Butler and Nesbitt 1999; Maslennikov et al. 2009; Meng et al. 2020). Therefore, the relatively low Se contents of pyrite and sphalerite formed by Zn-rich mineralisation are likely due to the extent of seawater mixing, given that seawater can decrease the temperature and increase the redox potential of hydrothermal fluids. In contrast to Se, Sn-rich sulfide minerals are limited to Zn-rich chimneys in the same vent field (Fig. 11). This suggests that Sn-rich sulfide minerals form from relatively oxidising, low-temperature fluids. Our results show that sphalerite is one of the major host minerals of Sn (Figs. 11b; ESM 1 S3) and, in sphalerite, a much higher proportion of Sn may be precipitated in its tetravalent rather than divalent state (Fig. 8c, d; Cook et al. 2009; Ye et al. 2011; Choi et al. 2021). Given that Sn exists mainly as an Sn(II) aqueous complex (i.e., SnCl2) in hydrothermal fluids (Uchida et al. 2002; Migdisov and Williams-Jones 2005), the oxidative transition from Sn(II) to Sn(IV) for Sn precipitation in sphalerite may have been facilitated by the relatively oxidising, low-temperature conditions. This is also consistent with the higher Sn contents of sphalerite in the Fe–Cu-rich sulfide samples as compared with the Cu-rich sulfide samples (Fig. 11b; ESM 1 S3).

It is generally considered that Fe2+ substitutes for Zn2+ within the sphalerite lattice (Keith et al. 2014; George et al. 2016). Concentrations of Fe and Sn in the AVF sphalerite exhibit a better positive correlation in the Cu-rich sulfide samples (R2 = 0.45) as compared with the Fe–Cu-rich sulfide samples (ESM 1 Fig. S7). This suggests that the direct substitution of Sn2+ for Zn2+ may have been facilitated in the former by the relatively reducing conditions. These features further suggest that the oxidation state of Sn is likely an important control on Sn contents in sphalerite. Our results and comparison allow us to conclude that the geochemistry of pyrite and sphalerite, particularly for Sn, is a more effective approach than bulk compositional analysis of hydrothermal samples for tracing the nature and origins of ultramafic-hosted mineralisation in MOR settings, as the mineralogical compositions of hydrothermal sulfide samples are highly variable in each vent field.

Comparison with VMS deposits on land: genetic and economic implications

It has been widely accepted that seabed hydrothermal venting and its mineralisation are the modern analogues of VMS deposits on land (Maslennikov et al. 2017; Martin et al. 2021). The VMS deposits represent a significant source of the world’s Cu, Zn, Pb, Au, and Au ores, with Co and Sn as by-products (Barrie and Hannington 1999; Hannington et al. 2010). They are conventionally classified into five groups based on host rock compositions: mafic, mafic-siliciclastic, bimodal-mafic, bimodal-felsic, and bimodal-siliciclastic types (Barrie and Hannington 1999). The extensive seafloor exploration during the last two decades recognized the ultramafic-hosted SMS (UM-SMS) deposits (Fouquet et al. 2010; Choi et al. 2021). This contributed to the reclassification of some VMS deposits on land into a sub-class of VMS deposits: the so-called ultramafic-hosted VMS (UM-VMS) deposits (e.g. Outokumpu deposit; Patten et al. 2022).

The UM-VMS and UM-SMS deposits formed as a result of hydrothermal events in volcanic submarine environments are characterised by relatively high contents of critical element such as Co, Ni, Sn, as well as precious and base metals (Fouquet et al. 2010; Maslennikov et al. 2017; Toffolo et al. 2020; Choi et al. 2021; Patten et al. 2022). Among those critical elements, however, significant contents of Sn are unexplained in terms of modern seafloor hydrothermal mineralisation because the source of Sn in VMS deposits was thought to be related to highly evolved magmatism underlying oceanic crust and/or detrital sediments from continents (Bleeker and Hester 1999; Hannington et al. 1999; Serranti et al. 2002). In addition, sulfides of modern SMS deposits have similar but highly variable Sn abundances regardless of the submarine environment (Peltonen et al. 2008 and their references). However, most UM-SMS deposits at MOR settings show consistent and relatively high average concentrations of Sn (up to ~ 2000 ppm) as compared with mafic-hosted SMS deposits (Fouquet et al. 2010; Evrard et al. 2015; Choi et al. 2021), although the precipitation process of Sn into sulfides in UM-SMS deposits is still enigmatic.

Patten et al. (2022) showed relatively high abundance of critical elements including Sn in both UM-VMS and UM-SMS deposits. To better understand Sn mineralisation in SMS and VMS deposits, we plotted the Sn contents of pyrite and sphalerite in the Dergamysh and Buribay VMS deposits in the Urals. The pyrite and sphalerite have higher Sn contents in the Dergamysh deposit than in the Buribay deposit, where ancient chimneys show a genetical affinity with those from ultramafic- and MORB-hosted SMS deposits, respectively (Fig. 11; Maslennikov et al. 2017). The values are similar to those of MOR-related SMS deposits, suggesting that the contribution of different host rocks to the hydrothermal mineralisation is reflected in the distribution of Sn at the mineral scale. This suggests that trace element variations in sulfides from seafloor hydrothermal mineralisation may enhance our understanding of the source of metals in the UM-VMS deposits.

Sulfur source for hydrothermal mineralisation

The δ34S values for pyrite from the different mineralisation stages exhibit a narrow range of 6.2 to 8.5‰, with a median value of 6.7‰ (n = 10; Table 3). These values can be explained by mixing between S of igneous origin (− 2 to + 2‰; Sakai et al. 1984; Alt et al. 1989, 2007; Alt and Anderson 1991) and S from the thermochemical sulfate reduction of seawater (+ 21‰; Rees et al. 1978) (Fig. 10). The proportions of S derived from these reservoirs can be calculated with the two end-member mixing model of Arnold and Sheppard (1981). We estimate that seawater sulfate is 29–40% of the S in the AVF pyrite (Table 3). This indicates that S was mainly derived from the associated igneous host rocks (60–70%), whereas a relatively smaller proportion of seawater S was incorporated into the AVF pyrite by fluid-seawater mixing. Therefore, we conclude that the AVF is a rock-dominated system and that the fluid–rock interaction (i.e. wall-rock leaching) during hydrothermal circulation is the dominant source of S and most metals. In particular, a comparison with other MOR systems shows that relatively high δ34S values (> 10‰) are limited to ultramafic-hosted systems (e.g. Rainbow, Logatchev, and Tianzuo; Fig. 10). This is likely due to the long history of fluid–rock interactions and greater proportion of seawater sulfate when considering the long-lived, deep hydrothermal circulation compared to MORB-hosted systems (Knight et al. 2018; Tao et al. 2020).

Conclusions

Hydrothermal sulfides from ultramafic-hosted mineralisation were collected from the Ari vent field on the slow-spreading middle part of the Central Indian Ridge. The sulfide samples can be classified as Fe–Cu- and Cu-rich types based on the major sulfide minerals. The Fe–Cu-rich sulfide samples record three different mineralisation stages: (1) stage I (subhedral–euhedral pyrite + isocubanite ± chalcopyrite ± magnetite ± pyrrhotite); (2) stage II (colloform pyrite ± sphalerite ± galena ± electrum); and (3) stage III (chalcocite ± uraninite) dominated by seawater alteration. As the AVF mineralisation progressed from stages I to III, the fluid temperature and ƒS2 decreased and ƒO2 increased. The Cu-rich sulfide samples are characterised by mineral assemblages and a paragenesis similar to those of the Fe–Cu-rich sulfide samples, but the more Cu-rich mineralisation with a higher proportion of isocubanite is indicative of relatively high-temperature and reducing mineralisation conditions.

The U-rich (up to 51.9 ppm) and Sn-poor (up to 2.1 ppm) compositions of the AVF sulfide samples are different from those of other ultramafic-hosted SMS deposits. The predominant occurrence of uraninite (< 1 μm in size) on altered surfaces of pyrite and hematite is the main form of U enrichment in the AVF. Ambient seawater was likely the principal source of U, and subsequent oxidative alteration of Fe-bearing minerals may have had an important role in the fixation of seawater-derived U to precipitate the discrete uraninite.

A comparison of sulfide minerals in different types of hydrothermal vent fields at MOR spreading centres reveals that Sn contents vary systematically between MORB- and ultramafic-hosted sphalerite and pyrite, with much higher Sn concentrations in the ultramafic-hosted environments. Sphalerite is one of the major hosts of Sn, whereas pyrite is Sn-poor. Therefore, the lower Sn contents of the Fe–Cu-rich sulfide samples (average = 6.1 ppm Sn) as compared with the Cu-rich sulfide samples (average = 48.9 ppm Sn) in the AVF are most likely due to Fe-rich mineralisation. These results suggest that the geochemistry of sulfide minerals rather than the bulk chemical composition of hydrothermal samples provides a clearer understanding of the nature of hydrothermal mineralisation in MOR settings. Tin could be one of the most effective elements for investigating the ore-forming processes in ultramafic-hosted hydrothermal deposits at MOR settings. This is also evidenced by a comparison of ancient VMS deposits, showing that the Sn contents of pyrite and sphalerite are higher in UM-VMS deposit than in mafic-hosted VMS deposit.

In situ δ34S values (+ 6.2 to + 8.5‰) of pyrite indicate that the S was mainly derived from the host igneous rocks (δ34S − 2 to + 2‰) with a smaller contribution (29–40%) of reduced seawater S (δ34S + 21‰). This indicates that fluid–rock interactions were significant in supplying the S and metals to the fluids in the AVF. Reducing, high-temperature fluids circulating through ultramafic rocks were important in forming the AVF. Such rock-dominated systems influenced by ultramafic-hosted mineralisation may be common along slow-spreading MOR settings.