Trace Element and Isotope Systematics in Vent Fluids and Sulphides From Maka Volcano, North Eastern Lau Spreading Centre: Insights Into Three-Component Fluid Mixing

1 Introduction

Heat loss during the cooling of the oceanic crust is the driver for hydrothermal fluid circulation (e.g., Von Damm, 1995; German and Seyfried, 2013; Humphris and Klein, 2018). Seawater that percolates through the oceanic crust is chemically modified due to water-rock interaction during low temperature alteration (<250°C) in the recharge zone and temperatures up to 400°C in the reaction and upflow zone, where most of the metals are leached (German and Von Damm, 2003; Bach et al., 2013; Humphris and Klein, 2018). In addition to the leaching process, metals may be contributed by magmatic volatiles, as known from island arc and some back-arc hydrothermal systems (de Ronde et al., 2011; Keith et al., 2018; Seewald et al., 2019; Fox et al., 2020; Martin et al., 2020). As the hot (up to 400°C), low pH (2–4) and metal-bearing fluid rises towards the seafloor strong physicochemical gradients (e.g., temperature, pH, redox, ligand availability) lead to the formation of (subsurface) hydrothermal precipitates due to processes like fluid boiling and fluid-seawater mixing (Diehl et al., 2020; Falkenberg et al., 2021; Keith et al., 2021; Schaarschmidt et al., 2021). Previous studies highlighted that fluid boiling is an efficient process for metal fractionation and precipitation in arc and back-arc hydrothermal systems due to temperature-pressure conditions below the critical point of seawater at 408°C and 301.1 bars (i.e., 3,060 m water depth; Stoffers et al., 2006; Monecke et al., 2014). Hence, the complex interplay of temperature-pressure conditions, host rock composition and magmatic volatile influx can produce hydrothermal fluids and related precipitates with a high compositional variability (Von Damm et al., 1985a; Gamo et al., 1997; Koschinsky et al., 2008; Mottl et al., 2011; Reeves et al., 2011; James et al., 2014; Seewald et al., 2015; Wohlgemuth-Ueberwasser et al., 2015; Keith et al., 2016a; Stucker et al., 2017; Kleint et al., 2019; Diehl and Bach, 2020; Martin et al., 2020).

Fluid boiling leads to the physical separation of a fluid into a low-Cl vapour and a high-Cl liquid phase enriched in Cl relative to seawater (539 mmol/kg Cl; Stoffers et al., 2006; Monecke et al., 2014; Schmidt et al., 2017). Depending on their volatility the dissolved elements either partition into the vapour (e.g., As, Sb, Te) or liquid phase (e.g., Fe, Co, Ni) and/or precipitate (e.g., Au) as a consequence of boiling-induced changes in fluid pH, redox, and ligand availability (e.g., Cl, H₂S; Drummond and Ohmoto, 1985; Grundler et al., 2013; Monecke et al., 2014; Pokrovski et al., 2018). Although, liquid-vapour partitioning and precipitation affect fluid chemistry in predictable ways (e.g., Von Damm et al., 1997; Stoffers et al., 2006; Schmidt et al., 2017) it is still unclear how these processes are recorded in the geological archive by sulphide chemistry (Tardani et al., 2017; Román et al., 2019; Keith et al., 2020; Falkenberg et al., 2021; Nestmeyer et al., 2021; Schaarschmidt et al., 2021).

The chemical composition of vent fluids in active systems is highly beneficial for understanding the ongoing hydrothermal processes, but provides only limited information on the temporal evolution of the hydrothermal system (Von Damm et al., 1997; Schmidt et al., 2017; Humphris and Klein, 2018). To date, time-series data from island arcs and adjacent back-arcs are scarce and underrepresented on a global scale, however, recent investigations of hydrothermal fluids and precipitates from the Southwest Pacific, suggest additional short-lived processes (days to months) affecting the geochemical composition of hydrothermal fluids, which is documented by the mineralogy and chemistry of the precipitates from active and inactive vents and chimney talus (Ditchburn et al., 2012; Berkenbosch et al., 2015, 2019; Seewald et al., 2015; Kleint et al., 2019; Diehl et al., 2020; Falkenberg et al., 2021). The complex physical, chemical and temporal aspects that control the chemical composition of hydrothermal fluids also directly influence the associated hydrothermal precipitates (Ditchburn et al., 2012; Keith et al., 2016b; Berkenbosch et al., 2019; Keith et al., 2021). The sampling of submarine hydrothermal vent systems is challenging and in contrast to the fluid compositions that represent a snapshot in time, associated hydrothermal precipitates accumulate and can document changing processes affecting the fluids during the evolution of the hydrothermal system (de Ronde et al., 2011; Fouquet et al., 2018; Berkenbosch et al., 2019; Diehl et al., 2020; Falkenberg et al., 2021).

Here, we report on vent fluid chemistry in combination with in situ trace element and isotope data of associated sulphide precipitates from two adjacent vent sites at Maka Volcano, North Eastern Lau Spreading Centre, SW Pacific Ocean. This combined approach provides important new insights into the evolution of submarine hydrothermal systems and the preservation of mixing and boiling in the geological record, as preserved by sulphide chemistry.

2 Geological Overview and Sampling Site

The North Eastern Lau Basin in the SW Pacific Ocean (Figure 1A) is known for its intense hydrothermal activity and complex tectonic setting (Zellmer and Taylor, 2001; Lupton et al., 2015; Embley and Rubin, 2018; Baker et al., 2019; Baxter et al., 2020; Anderson et al., 2021). The North Eastern Lau Spreading Centre (NELSC, Figure 1B) is situated in the western part of the Lau back-arc basin and follows a NE-orientation with an estimated spreading rate of ∼42 mm yr⁻¹ (Baxter et al., 2020). It can be subdivided into four segments with the axial Maka volcanic edifice being situated on the southernmost segment (Figure 1B, Baker et al., 2011; Anderson et al., 2021). The geochemical composition of the lavas at the NELSC varies from basaltic to dacitic composition, with basaltic compositions dominating at Maka volcano (Haase et al., 2018; Zhang H. et al., 2019).

FIGURE 1. (A) Large scale bathymetry of the northern Lau Basin, SW Pacific Ocean. The black star marks the study site situated within the North East Lau Basin. Bathymetric data taken from (GEPCO Compilation Group, 2021). (B) Detailed bathymetric map of the North East Lau Basin [black rectangle in (A)]. The sample location at Maka volcano is marked with a black star. (C) Sample location at the top of Maka volcano [black rectangle in (B)]. Bathymetric data presented in (B, C) taken from (Merle et al., 2018). (D, E) Active hydrothermal venting at Maka HF and Maka South. Photographs taken during ROV dives by MARUM QUEST 4000, University Bremen.

Maka is the southernmost axial volcanic edifice at the NELSC, where first evidence for hydrothermal activity was indicated by a plume survey in 2004 (German et al., 2006). The first visual confirmation of actively venting high temperature fluids (∼315°C) was made by Nautilus Minerals Inc. in 2008 (Baker et al., 2011). Since then, several plume studies focused on southernmost segment of the NELSC, leading to the discovery of eight high-temperature vent sites (German et al., 2006; Kim et al., 2009; Resing et al., 2009; Baker et al., 2011, 2019). At the summit of Maka volcano two active hydrothermal vent sites occur at 1,525–1,543 m water depth (Figure 1C), which are characterised by distinct vent structures and fluid discharge (Figures 1D,E). At the Maka hydrothermal field (Maka HF) vigorously venting black smoker-type fluids with temperatures of up to 329°C (Figure 1D) were observed together with smaller chimneys emitting clear fluids with maximum temperatures of 292°C (Haase et al., 2018). Less than a hundred meters south of Maka HF, a second vent site named Maka South was discovered (Figure 1E), where clear to white fluids are discharged at several sites from flanges and diffuse areas with temperatures between 267 and 301°C (Haase et al., 2018).

3 Sampling and Analytical Methods

Sulphide-sulphate samples (n = 6) and pillow lavas with fresh glass rims (n = 4) were recovered during two dives (Table 1) using the ROV MARUM QUEST 4000 (University of Bremen). Chimney talus (116 ROV-17) and fragments from a 15–20 m high active black smoker (116 ROV-13) emitting high temperature fluids (up to 329°C) were recovered at Maka HF. Samples from flanges discharging white smoke of up to 301°C were sampled at Maka South (122 ROV-17) together with talus material and samples from inactive chimneys (122 ROV-18, 22, 24). Pillow lava samples with glass rims were recovered at Maka HF (116 ROV-09, -07) and Maka South (122 ROV-09, 10). Hydrothermal fluids were sampled using Isobaric Gas-Tight (IGT) samplers, which were deployed on the ROV. Fluid samples (n = 6) were taken after the inlet snorkel was stably positioned and temperature reading on the IGT sampler was constant. After retrieving the sampler, multiple aliquots were sequentially taken from each IGT for on-board analysis of O₂, H₂S, SO₄, H₂ and CH₄ concentrations and fluid parameters, such as pH, Eh and salinity (Table 1). Further subsamples were taken for shore-based elemental and isotope analyses of the vent fluids and sulphides (cf. Sections 3.1, 3.2).

TABLE 1. Overview of collected fluid and sulphide samples with maximum fluid temperatures and on-board measured parameters.

3.1 Hydrothermal Sulphide and Volcanic Glass Analysis

3.1.1 Major and Trace Elements

Bulk analyses of sulphide-sulphate samples (n = 5) were performed at Activation Labs in Ontario (Canada) using a combination of thermal desorption (TD ICP-MS) and peroxide fusion inductively coupled plasma mass spectrometry (FUS ICP-MS) and instrumental neutron activation analysis (INAA). Detection limits, blanks, analyses of certified reference materials and the analytical results are presented with respect to the different analytical techniques in the supplementary material (Supplementary Table S1).

The trace element composition of volcanic glass samples from Maka (n = 4) was determined on hand-picked glass separates by a Thermo Scientific XSERIES 2 Quadrupole ICP-MS at the GeoZentrum Nordbayern (Supplementary Table S2). Details about the analytical procedure are presented in Storch et al. (2020).

Major elements in pyrite, chalcopyrite and sphalerite (n = 113) were detected by electron probe micro-analysis using a JEOL JXA-8200 Superprobe at the GeoZentrum Nordbayern (Supplementary Table S3). For the quantitative analyses, a focused beam with an acceleration voltage of 20 kV and a beam current of 20 nA were used. The following reference materials were used for the standardization: FeS₂ (Fe, S), CuFeS₂ (Cu), ZnS (Zn), InAs (As), Ag2Te (Ag), CdS (Cd), InSb (Sb), PbTe (Te, Pb) and Au (Au). Data with analytical totals of <98 and >101 wt% were discarded. Each measuring spot was marked on a back-scattered electron image for the subsequent analysis of the same grain by laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS). Most elements, except Fe, S, Cu and Zn, have concentrations in the trace element range (<0.1 wt%) in pyrite, chalcopyrite and sphalerite, and hence the LA-ICP-MS data were preferentially used for As, Ag, Cd, Sb, Te, Au and Pb.

The trace element analysis of hydrothermal sulphides (n = 96) was performed at the GeoZentrum Nordbayern by a Teledyne Analyte Excite 193 nm laser coupled with an Agilent 7500c quadrupole ICP-MS (Supplementary Table S3). The mass spectrometer operated with a plasma power of 1300 W; He (0.9 L/min) and Ar (0.98 L/min) were used as carrier gases. In addition, Ar also acted as plasma (14.9 L/min) and auxiliary gas (0.9 L/min). A single spot ablation pattern with a frequency of 20 Hz and a fluence of 3.5 J/cm² was used. Total analysis time for each spot was set to 45 s including 20 s of gas blank analysis prior to ablation. A beam size of 25 mμ and on occasion of 15 mμ was used according to the size of the analyzed grain. The following standards were used for the external calibration: Po724 B2 SRM (for Au; Memorial University Newfoundland) and MASS-1 (for V, Mn, Co, Ni, Cu, Zn, Ga, Ge, As, Se, Mo, Ag, Cd, In, Sn, Sb, Te, W, Au, Tl, Pb, Bi). In addition, UQAC-FeS-1 (https://sulfideslasericpms.wordpress.com), a homogenous nano-powder of natural sulphides doped with a range of trace elements for matrixed matched analyses of Fe-sulphides, was used as a secondary standard. Analytical precision and accuracy were monitored by repeated analysis of Po724 B2 SRM, MASS-1 and UQAC-FeS1 (Supplementary Table S4). The relative standard deviation (RSD) in Po724 B2 SRM and MASS-1 for most elements is <10%, except Au (11.4%), Tl (10.2%) and Pb (10.9%). In the secondary standard (UQAC-FeS1) only Ga (11.9%) and Cd (10.8%) showed RSD values > 10%, which are therefore comparable to the primary standards (Supplementary Table S4). Trace element concentrations and detection limits were processed by Glitter (van Achterbergh et al., 2000). Spikes in the time-resolved LA-ICP-MS depth profiles were excluded for the calculation of the trace element concentrations.

3.1.2 Sulphur and Pb Isotopes

Sulphur isotope analysis was conducted at the Westfälische Wilhelms-Universität Münster using Flash EA IsoLink elemental analyzer interfaced to a ThermoScientific Delta V Advantage isotope ratio mass spectrometer (EA-IRMS). The measurements were performed on pyrite (n = 4), chalcopyrite (n = 5) separates without further chemical treatment (Supplementary Table S5). About 100 μg of sulphide powder and 300–500 μg V2O5 were admixed in tin capsules. Measured values are presented in per-mille difference relative to the Vienna-Canyon Diablo Troilite (V-CDT) reference material. Reproducibility, as determined by replicate analyses of the reference material was better than $\pm$0.2‰ (1$\sigma$).

For the Pb isotope analysis (n = 4, Supplementary Table S5), 1–5 mg of sulphide powder was dissolved in 0.4 ml aqua regia (50/50 = 6 M HCl/15 M HNO₃) at 80°C in Teflon beakers. After evapouration on a hot plate the material was redissolved in 1 M HCl. Lead was separated from the rock matrix using 100 μL Eichrom Sr-Spec resin. The sample solution was loaded onto the resin and rinsed with 1 M HCl before collecting Pb in 6 M HCl. All reagents were extra distilled and typical procedural blanks for Pb were 50 pg. The chemical sample treatment in advance of the Pb isotope analysis of the volcanic glass fragments was performed according to an established method as described in Haase et al. (2019). Lead isotope measurements were carried out at the GeoZentrum Nordbayern by a Thermo-Fisher Neptune Plus multicollector ICP-MS using a²⁰⁷Pb/²⁰⁴Pb double spike to correct for instrumental mass fractionation. The double spike, with a²⁰⁷Pb/²⁰⁴Pb ratio of 0.82, was calibrated against a solution of the NBS981 equal atom Pb standard. Samples were diluted with 2% HNO₃ to a concentration of approximately 20 ppb, and an aliquot of this solution was spiked in order to obtain a²⁰⁸Pb/²⁰⁴Pb ratio of about 1. Spiked and unspiked sample solutions were introduced into the plasma via a Cetac Aridus desolvating nebulizer and measured in static mode. Interference of ²⁰⁴Hg on ²⁰⁴Pb was corrected by monitoring ²⁰²Hg. An exponential fractionation correction was applied offline using the iterative method of Compston and Oversby (1969) with a correction of typically 4.5‰ per amu. The NBS981 Pb isotope standard, measured as an unknown together with the samples yielded ²⁰⁶Pb/²⁰⁴Pb, ²⁰⁷Pb/²⁰⁴Pb, ²⁰⁸Pb/²⁰⁴Pb ratios of 16.9417 ± 0.0001, 15.4991 ± 0.0001 and 36.7228 ± 0.0002, respectively. Data are reported relative to our long-term average of the NBS981 Pb standard (16.9410 ± 0.0020 for ²⁰⁶Pb/²⁰⁴Pb, 15.4993 ± 0.0019 for ²⁰⁷Pb/²⁰⁴Pb, 36.7244 ± 0.0046 ²⁰⁸Pb/²⁰⁴Pb).

3.2 Hydrothermal Fluids

3.2.1 On-Board Measurements

The on-board analyses (Table 1) for pH, redox potential (Eh), salinity and dissolved oxygen (O₂) were performed using a WTW pH/Cond 340i multimeter. The following sensors were used for the analysis: SenTix ORP platinum electrode with 3M KCl reference electrolyte for Eh, SenTix 41 probe for pH, WTW TetraCon 325 conductivity probe for salinity and optical WTW FDO sensor for O₂. The on-board analysis of dissolved H₂ and CH₄ was carried out using a headspace extraction method (with gas to liquid volume ratios of >3) and a 7820A Agilent gas chromatograph that operated with N₂ as carrier gas. The gas chromatograph was calibrated with a reference gas containing 253 mol-ppm H₂, 120.1 mol-ppm CH₄, 100.5 mol-ppm CO and 101.1 mol-ppm CO₂. The samples were directly injected after the headspace extraction (including water vapor) and were separated in a Molsieve 60/80 column (Sigma Aldrich, St. Louis, MO) with H₂ and CH₄ being simultaneously detected in a serial set-up of a thermal conductivity detector and a flame-ionization detector, respectively. For each fluid sample with sufficient sample volume the headspace extraction was performed on three individual aliquots to evaluate the precision of the analyses (samples 116 ROV-12 and 122 ROV-15 were exhausted during draw of other aliquots before possible measurement of multiple aliquots). The precision was better than 10% for all samples where multiple aliquots were measured. Concentrations of dissolved H₂S were analysed photometrically following the methylene blue method described by Cline (1969). Dissolved sulphide was initially stabilized in a gelatinous zinc acetate solution, followed by the addition of a colour reagent (N,N′-Dimethyl-1,4-phenylendiamin) and a catalyst (FeCl₃). Photometric measurements were carried out using a ThermoScientific Genesys 10 UV photometer at 660 nm wavelength. The dissolved sulphate (SO₄) concentration was determined gravimetrically by precipitation as insoluble BaSO₄. Volumes of 15–75 ml of the solution were acidified with hydrogen chloride to a pH ≤ 2, heated to near boiling point temperatures (∼85°C) and mixed with BaCl₂ solution (8.5%) to one fifth of the original volume. The BaSO₄ precipitate was subsequently filtered by a pre-weight 0.45 µm cellulose nitrate filter and dried at 40°C.

3.2.2 Major and Trace Elements

The fluid samples were pressure-filtered through 0.2 µm polycarbonate (PC) filters under a laminar flow bench, acidified to pH < 2, using suprapure 30% HCl and stored cool in acid-cleaned PE bottles until further analysis. Concentrations of major and minor elements (B, Br, Cl, Na, Ca, K, Fe, Mn, K, Sr, Li, Mg, Zn, Cu, Si) were determined by inductively coupled plasma – optical emission spectroscopy (ICP-OES, Spectro Ciros), using a matrix-matched calibration at Jacobs University Bremen. Trace and rare Earth elements (REE) were determined by matrix separation following Schmidt et al. (2010) and measured by inductively coupled plasma – mass spectrometry (ICP-MS, Perkin Elmer NexION) at Jacobs University Bremen (Table 2, Supplementary Table S6). Accuracy for major and minor elements was controlled by repeated measurements of the reference materials IAPSO standard seawater (values from Ocean Scientific International Ltd.) and SLRS-6 (certified values from the National Research Council Canada). As quality control for trace and rare Earth elements, the certified reference materials NASS-7 (National Research Council Canada) and TraceCert (Rare Earth Element Mix for ICP, Sigma Aldrich) were measured alongside the samples. The analytical uncertainty of the measured reference materials was ±5% with respect to the certified value.

TABLE 2. Major, minor and trace element data together with gas concentrations and stable and radiogenic isotopes in the vent fluids from Maka HF and Maka South. Notation: b.d.: below detection limit.

3.2.3 Total Dissolved CO₂

Aliquots for the concentration measurements of CO₂ were transferred in helium-filled and evacuated glass vials for gas chromatographic analyses at the University of Bremen (Table 2). The analysis was carried out following a modified procedure after Reeves et al. (2011) with improved headspace-liquid vapour partitioning calculation. Aliquots were acidified with phosphoric acid (25 wt%) to convert dissolved inorganic carbon (DIC) species to CO₂. Subsequently, a Haysep 80/100 column (Sigma-Aldrich, St. Louis, MO) was used with helium as carrier gas in a Thermo Scientific Trace GC Ultra to quantify the CO₂ concentrations. The gas chromatograph was calibrated with pure CO₂ (99.995 vol%) acting as reference gas. A 25 mmol/kg CO₂ standard was measured alongside the samples to verify the measurements. For each sample, two aliquots were measured three times resulting in a precision of better than 5% (on a relative scale). Measurements of the CO₂ standard yielded an average value of 21.8 mmol/kg for the 25 mmol/kg solution suggesting that the accuracy is better than 15%, but also that the sample concentrations may be systematically under-estimated by up to 12%.

3.2.4 Stable and Radiogenic Isotopes

Aliquots of the samples were filtered on-board through 0.2 µm polycarbonate (PC) filters. Afterwards the Sr was separated from the sample matrix using a one-step column with Sr spec as resin. This chemical separation was modified after Pin and Bassin (1992). After separation, the Sr fraction was loaded together with a tantalum emitter onto rhenium filaments and analysed on a ThermoFisher Scientific TRITON Plus thermal ionization mass spectrometer (TIMS) in dynamic mode at MARUM Bremen (Table 2). To correct for instrumental mass fractionation, Sr isotope ratios were normalized to 0.1194 ⁸⁶Sr/⁸⁸Sr The standard reference material NIST SRM 987 was analysed to determine the accuracy of the Sr isotope measurements, yielding 0.71025 ± 0.00002 (2SD_mean, n = 3) for ⁸⁷Sr/⁸⁶Sr, which is in agreement with published TIMS values of NIST SRM 987 [0.710250 ± 0.000044 (2SD_mean, n = 1,596); http://georem.mpch-mainz.gwdg.de].

Boron was separated from the sample matrix using a micro sublimation technique after Wang et al. (2010) and Wilckens et al. (2018). The B isotope analyses were performed by a ThermoFisher Scientific Neptune Plus multicollector-inductively coupled plasma-mass spectrometer (MC-ICP-MS) equipped with a stable introduction system and a high-efficiency x-cone at MARUM Bremen (Table 2). The measurements were carried out by the standard-sample-bracketing method using a 100 ng/g B solution and the NIST SRM 951 boric acid reference material as bracketing standard. Boron isotope values are reported in the conventional δ¹¹B (‰) notation relative to NIST SRM 951. The δ¹¹B value of an internal seawater standard (bottom water from SuSu Knolls, +39.6 ± 0.2, 2SD, n = 5) is in agreement with the literature value (+39.6 ± 0.2, 2SD; Foster et al., 2010), validating the sublimation and measurement technique. The accuracy and precision of the δ¹¹B method was ±0.2‰ (2SD_mean, n = 7) for δ¹¹B, obtained by the repeated analysis of NIST SRM 951.

Fluid samples were analyzed for δ¹⁸O and δ²H at the GeoZentrum Nordbayern. The δ¹⁸O analysis was performed by an automated equilibration unit (Gasbench 2; Thermo Fisher Scientific) in continuous flow mode coupled to a Delta plus XP isotope ratio mass spectrometer (Thermo Fisher Scientific). All samples were measured in duplicates and the reported value is the mean value. The δ²H measurements were carried out by liquid injection into a modified high temperature pyrolysis unit (Thermo TC/EA with CTC PAL autosampler) coupled in continuous flow mode to a Delta V plus isotope ratio mass spectrometer (Thermo Fisher Scientific). The TC/EA unit was converted into an online chromium reduction system based on the principle outlined by Morrison et al. (2001). The glassy carbon tube was removed from the TC/EA and the ceramic tube was filled with quartz wool, quartz glass chips and chromium powder (#10147, 99%, -100 mesh, Alfa Aesar, Karlsruhe, Germany). Temperatures were set to 840°C (reactor temperature) and 85°C (GC column). Injection volume was 1.2 uL by using a 10 uL syringe (SGE Analytical Science, Australia). δ¹⁸O and δ²H values are reported in the standard δ-notation in permil (‰) on the VSMOW/SLAP (Standard Light Antarctic Precipitation/Standard Light Antarctic Precipitation) scale by assigning a value of 0 and –55.5‰ (δ¹⁸O) and 0 and −427.5‰ (δ²H) to VSMOW2 and SLAP2, respectively (Brand et al., 2014). For normalization two laboratory standards that were calibrated directly against VSMOW2 and SLAP2, were measured in each run. External reproducibility based on repeated analyses of a control sample was better than 0.1 and 1‰ (±1 sigma) for δ¹⁸O and δ²H, respectively. The data sets were corrected for memory and instrument drift during the run. Sequence setup and post-run correction procedures were adopted from van Geldern and Barth (2012).

3.2.5 Calculation of Hydrothermal End-Members

Compositions of vent fluids may be affected by seawater mixing prior to or during sampling, when the snorkel of the IGT is not fully centred or as a result of complex vent structures, such as the flanges at Maka South (Figure 1E). Therefore, fluid end-member compositions (Supplementary Tables S7, S8), reflecting the “pure” hydrothermal fluid without a seawater component, were calculated for all fluid samples based on their Mg concentration (Table 2) by linear least square regression through seawater values and extrapolation to 0 Mg (Supplementary Figure S1; Von Damm et al., 1985b). According to the same method, end-member temperatures were calculated for 0 Mg in the vent fluids. Isenthalpic-isobaric mixing paths were constructed using the temporally averaged fluid temperature during sample intake, the Mg concentration of the hydrothermal fluid, as well as the temperature of seawater and the seawater Mg concentration. The composition- and temperature-dependent isobaric heat capacity was calculated for fluid mixing processes at seafloor pressures using the empirical relationship presented in Driesner (2007) that was applied to the thermodynamic dataset of pure water presented in Haar et al. (1984).

4 Results

4.1 Petrography of Sulphide-Sulphate Samples

Pyrite, chalcopyrite, sphalerite and marcasite occur throughout the chimney walls in variable proportions (Table 3) and exhibit distinct textures between the outer chimney wall and the central fluid conduit (Figure 2). Colloform pyrite (py I) typically occurs in the outer chimney wall (Figure 2A), where it is associated with sphalerite (sph I) that overgrew barite (Figure 2B). By contrast, chalcopyrite is only a minor phase in the outer chimney wall and occurs as part of chalcopyrite-sphalerite (sph I) alternations surrounding pyrite (py I, Figure 2A). Marcasite (mrc I) also occurs in this mineral assemblage typically forming a seam around pyrite (py I, Figures 2C,D). Towards the intermediate chimney wall chalcopyrite abundances increase, as commonly reflected by chalcopyrite-sphalerite (sph II) alternations or the occurrence of sub-to euhedral chalcopyrite that surrounds pyrite (py I, Figures 2E,F). Besides chalcopyrite, euhedral pyrite (py II) is a common constituent in the intermediate chimney wall (Figure 2G), which is surrounded by pyrite (py III) and marcasite (mrc II). The inner part of the chimney wall near the central fluid conduit is dominated by chalcopyrite and secondary Cu-sulphides, like bornite, covellite, and chalcocite-digenite that occur in minor amounts particularly in the inactive samples (Figure 2H). Secondary Cu-sulphides were also rarely observed in samples from active chimneys, where they are associated with sphalerite and barite in the outer chimney wall. However, the main mineral assemblage (pyrite, sphalerite, chalcopyrite) in active and inactive vents and in chimney talus are comparable (Table 3). Similarly, mineral abundances and textures do not vary between Maka HF and South, with the exception of marcasite that only occurs at the former vent site (Table 3).

TABLE 3. Summary of mineral abundances in active and inactive vent samples and in chimney talus from Maka.

FIGURE 2. Photomicrographs in reflected light of representative sulphide-sulphate samples from Maka volcano. (A) Colloform pyrite I (py I) surrounded by sphalerite I (sph I) and chalcopyrite (cpy, 122 ROV-24, Maka South). (B) Sulphide assemblage in the outer chimney wall consisting of pyrite I, sphalerite I and minor chalcopyrite (116 ROV-17R, Maka HF). (C) Early pyrite I surrounded by marcasite I (mrc I) and followed by chalcopyrite and late-stage sphalerite II (sph II, 116 ROV-17, Maka HF). (D) Pyrite I and marcasite I assemblage surrounded by chalcopyrite and pyrite II of the high temperature stage (116 ROV-17, Maka HF). (E) Alternations of chalcopyrite and sphalerite II (122 ROV-24, Maka South). (F) Pyrite I surrounded by subhedral chalcopyrite (122 ROV-22, Maka South). (G) High temperature euhedral pyrite II associated with late-stage pyrite III (py III) and marcasite II (mrc II, 116 ROV-17, Maka HF). (H) Chalcopyrite lining around fluid conduit replaced by supergene bornite (bn), covellite (cv) and chalcocite (cc, 122ROV-18, Maka South). Estimations of mineral abundances in the different samples are presented in Table 3.

4.2 Chemistry of Hydrothermal Sulphides and Volcanic Glass

4.2.1 Bulk Sulphide-sulphate Chemistry

The sulphide-sulphate samples from Maka HF and South overlap in their bulk chemical composition. However, differences were observed with respect to the overall compositional range, as reflected by elevated Mn, Co and Bi contents at Maka HF and a tendency towards higher Cu, Zn, Se, Mo, Ag, Cd, Sb and Ba contents (max. values) at Maka South (Figure 3). Compared to other back-arc hydrothermal systems, Maka is enriched in Co and depleted in Ag, Pb and Bi, whereas the contents of most other elements (e.g., Zn, As, Se, Ba, Au) are comparable. In addition, the sulphide-sulphate samples from Maka are enriched in most elements except Mn and Ni compared to pyrite from the same vent site (e.g., Ni, Cu, Zn, Se, Mo and Bi at Maka South) and with respect to the seafloor lavas (except Mn and Ni, Figure 3).

FIGURE 3. Multi-element diagram comparing the composition of the bulk sulphide-sulphate data from Maka HF and Maka South with the volcanic host rocks, as well as with pyrite from the two vent sites at Maka and from Niua South (Falkenberg et al., 2021). The presented data is normalised to the average bulk composition of back-arc sulphide-sulphate samples (Supplementary Table S9).

4.2.2 Major and Trace Elements in Hydrothermal Sulphides

The S and Fe contents of pyrite (n = 65) range from 51.1 to 53.8 wt% and 45.3 to 47.3 wt%, respectively. Variations in the major element composition of chalcopyrite (n = 29) were observed for S (34.2–35.0 wt%), Fe (29.8–30.4 wt%) and Cu (33.0–34.8 wt%). Differences in the major element composition of pyrite and chalcopyrite between Maka HF and Maka South are insignificant. Sulphur (31.8–33.9 wt%), Fe (0.20–14.1 wt%), Cu (0.01–0.68 wt%) and Zn (50.1–65.4 wt%) in sphalerite are in the major and minor element range. The contents of Fe and Zn in sphalerite show a strong negative correlation (R² = 0.96, n = 12). The sphalerite data can be subdivided into a low-Fe (<2.7 wt%) and a high-Fe (>10.6 wt%) group. Both sphalerite types are present at Maka HF, whereas sphalerite compositions from Maka South are restricted to the low-Fe group (Supplementary Table S3).

Significant variations were observed in the trace element composition of pyrite between Maka HF and South. On average, pyrite from Maka HF is enriched in Mn, Co, Ni, Se, Mo, Te, Tl and Bi, but depleted in Cu, Zn, As, Cd, Sb, Au and Pb compared to pyrite from Maka South (Figure 3). Similar trace element concentrations in pyrite between the two vent sites were observed for Ge and Ag. Compared to pyrite from Niua South (Falkenberg et al., 2021), pyrite from Maka rather exhibits a depleted trace element signature and enrichments compared to the former were only observed for Mn, Ni, Se, Te and Bi at Maka HF and Cu at both vent sites. Hydrothermal pyrite from Maka is enriched in most trace elements, except Mn and Ni, relative to the volcanic glass samples from the seafloor lavas (Figure 3).

Trace element variations were also observed between the different pyrite-types (Figure 4). Pyrite I tends to show higher Tl and Pb compared to pyrite II (and III) from the same vent site (Figure 4A). Pyrite II from Maka HF (and South) is enriched in Co and Se compared to pyrite I and III from the same vent site (Figure 4B). Pyrite III shows a wide compositional range with respect to most trace elements, and therefore overlaps with the data of pyrite I and II (Figure 4). Trace element ratios in pyrite between Maka HF and South show systematic variations, such as higher Tl/Pb, Se/Ge (and Co/Ni), as well as lower Sb/Pb and As/Co in pyrite from Maka HF compared to the same pyrite-type from Maka South (Figure 4). A negative correlation is described between Co/Ni – As/Co and Se/Ge – As/Co (Figure 5), in which respect pyrite from Maka South is characterized by a tendency to higher As/Co (>10), as well as lower Co/Ni (<10) and Se/Ge (<10) ratios (Figure 5). Pyrite II from Maka HF generally exhibits the lowest As/Co (<10) and highest Co/Ni (>10) and Se/Ge (>10) ratios. By contrast, pyrite I from Maka South tends to show the highest As/Co (>100) and lowest Co/Ni (<10) and Se/Ge (<10) ratios. Pyrite III from Maka HF and South typically overlap and show intermediate As/Co (10–100), Co/Ni (∼10) and Se/Ge (∼10) ratios. The observed variations in Co/Ni, As/Co and Se/Ge are most significant between the same pyrite-type from the two vent sites (e.g., Se/Ge – As/Co in pyrite II, Figure 5B).

FIGURE 4. Bivariate trace element diagrams of (A) Tl – Pb, (B) Co – Se, (C) Se – Ge, (D) Co – Ni, (E) Sb – Pb, (F) As – Co in pyrite from Maka HF and Maka South. The dashed lines represent the trace element ratios of the respective elements in the diagram (e.g., Tl/Pb in A). Superscripts: 1 = (Falkenberg et al., 2021), 2 = (Román et al., 2019).

FIGURE 5. Diagram showing trace elements ratios of (A) Co/Ni and (B) Se/Ge versus As/Co. Note that pyrite, which formed from black smoker-type fluids with seawater-like chlorinities is characterised by high Co/Ni (>10) and Se/Ge (>10) and low As/Co (>10–100), whereas pyrite that formed from low-Cl vapour-rich fluids and/or by seawater mixing shows low Co/Ni (<10) and Se/Ge (<10) and high As/Co (>10–100). For a more details, please see Section 5.3. Superscripts: 1 = (Falkenberg et al., 2021), 2 = (Román et al., 2019)

4.2.3 Sulphur and Pb Isotopes in Hydrothermal Sulphides

Sulphur isotope (n = 9) values of hydrothermal sulphides from Maka show a narrow range from −0.6 to 1.3‰ compared to other island arc and back-arc vent fields (Figure 6, Supplementary Table S5). The δ³⁴S composition of hydrothermal sulphides from Maka South vary from −0.6 to 1.3‰ and overlap with the compositional range of Maka HF (0.0–0.4‰). With respect to the different hydrothermal sulphides analysed in this study, chalcopyrite (−0.8–1.3%) tends to show slightly more positive δ³⁴S values than the associated pyrite (−0.6–0.0‰, Supplementary Table S5). Irrespective of the vent site and the analysed sulphide, the δ³⁴S data from Maka shows a compositional range similar to mid-ocean ridge basalt (−1.9–0.8‰; Sakai et al., 1984; Labidi et al., 2012; Figure 6).

FIGURE 6. Comparison of δ³⁴S values in hydrothermal sulphides from Maka with other mid-ocean ridge, back-arc and arc related hydrothermal systems. Reference data as listed in Supplementary Table S8 in the supplementary material. The grey field reflects the compositional range of MORB (−1.9–0.8‰, Supplementary Table S8), the dashed line is the average δ³⁴S value of global arc volcanic rocks (5.8 $\pm$ 4.0‰, Supplementary Table S11) and the blue line is the δ³⁴S composition of modern seawater (∼21.2‰; Tostevin et al., 2014).

Variations in the Pb isotope composition of hydrothermal sulphides (n = 6) from Maka are minor, as indicated by ²⁰⁶Pb/²⁰⁴Pb, ²⁰⁷Pb/²⁰⁴Pb and ²⁰⁸Pb/²⁰⁴Pb values that vary from 18.9474 to 18.9922, 15.6040 to 15.6154 and 38.8399 to 38.8623 (Supplementary Table S5). Systematic Pb isotope variations were not observed between the vent sites of Maka HF and South (Figure 7), as well as between the analysed pyrite and chalcopyrite separates (Supplementary Table S5). The Pb isotope composition of fresh volcanic glass from the seafloor lavas varies from 19.0781 to 19.0849, 15.6169 to 15.6244, 38.9312 to 38.9546 in ²⁰⁶Pb/²⁰⁴Pb, ²⁰⁷Pb/²⁰⁴Pb and ²⁰⁸Pb/²⁰⁴Pb, respectively (Haase et al., submitted manuscript¹), without showing any systematic variation between the two vent sites. The volcanic glass samples are characterized by more radiogenic Pb isotopes composition than the hydrothermal sulphide separates (Figure 7).

FIGURE 7. Comparison of (A) ²⁰⁷Pb/²⁰⁴Pb and (B) ²⁰⁸Pb/²⁰⁴Pb versus ²⁰⁶Pb/²⁰⁴Pb between the Maka seafloor lavas and vent sulphides. Note that the seafloor lavas are more radiogenic than the associated hydrothermal sulphides pointing towards a second less radiogenic reservoir in the reaction and/or upflow zone beneath the Maka vent field, which is suggested to be comparable to the seafloor lavas nearby the Maka vent system and from adjacent segments in the NELSC (e.g., <18.95 in ²⁰⁶Pb/²⁰⁴Pb; NELSC data). The mixing-line (black line) quantifies the contribution from these two reservoirs, as preserved in the vent sulphides (see Section 5.4 for details). The dark grey field reflects the compositional range of Pacific pelagic sediments (Regelous et al., 1997; Ewart et al., 1998; Beier et al., 2017) and the light grey field represent the Pb isotope variation of other North Tonga lavas (Regelous et al., 1997; Wendt et al., 1997; Ewart et al., 1998; Falloon et al., 2007; Falkenberg et al., 2021).

4.3 Chemistry of Vent Fluids

4.3.1 Temperature and pH

Measured temperatures of black smoker fluids at Maka HF were 322 ± 2°C (116 ROV-12) and 269 ± 24°C (116 ROV-15) with peak temperatures of 329 and 292°C, respectively. At Maka South, measured fluid temperatures from flanges that discharge white and clear fluids were 196 ± 67.6°C (122 ROV-15), 258 ± 18.8°C (122 ROV-16) and 183 ± 70.6°C (122 ROV-19) with peak temperatures of 301, 298 and 267°C, respectively (Table 1). The calculated temperatures of the end-member fluids are 301–345°C at Maka HF and 379°C at Maka South (Figure 8). End-member fluid temperatures for Maka South could only be determined for one fluid sample (122 ROV-15), which has the lowest Mg contents (31.0 mmol/kg) at this site (Table 2). The isobaric-isenthalpic end-member temperature calculations for the other fluid samples from Maka South (122 ROV-16, -19) yielded unreasonably high temperatures (>1,000°C) for a fluid that was affected by mixing with cold ambient seawater (∼2°C, Figure 8). The measured pH_{(25°C,1 atm)} of the Maka South fluids (2.79–3.33) was higher than those from Maka HF (4.53–5.42, Table 1).

FIGURE 8. Temperature vs. Mg concentrations of the Maka vent fluids and predicted isenthalpic-isobaric mixing paths between hydrothermal end-members and cold seawater (black star). The dashed arrows indicate the direction of mixing paths with heated and compositionally unmodified seawater to explain the high temperatures in combination with the high Mg concentrations at Maka South.

4.3.2 Major and Trace Elements

At Maka HF, the concentrations of Mg and SO₄ in the hydrothermal fluids vary from 5.52 to 13.4 mmol/kg and 0.50–5.35 mmol/L, respectively (Table 2) and are highly depleted compared to seawater (52.8 mmol/kg Mg and 28.1 mmol/L SO₄). The Cl concentrations range from 502 to 549 mmol/kg and are similar to seawater (539 mmol/kg Cl). At Maka South, concentrations of Mg, Cl, Br, Na, Sr, B and U are close to or below seawater concentrations, whereas K, Li, Ca, Mn, Rb, Cs, Ba, Co and Pb are slightly enriched relative to seawater (Table 2). The measured H₂S concentrations are highly enriched at Maka South (up to 37,147 μmol L⁻¹) compared to the fluids from Maka HF (up to 13,059 μmol L⁻¹, Table 1). By contrast, H₂ concentrations at Maka HF (up to 59.6$\pm$3.3 μmol L⁻¹) significantly exceed those from Maka South (up to 17.1 μmol L⁻¹). Measured concentrations of CH₄ and CO₂ in the fluids from both vent sites overlap despite significantly higher Mg concentrations at Maka South (Tables 1, 2). The fluids from Maka HF are characterised by higher REE and Y contents compared to those from Maka South (Supplementary Table S6). Chondrite-normalized REE patterns of the fluids from both vent sites show an enrichment of light over heavy REEs, as well as a positive Eu anomaly (Supplementary Figure S2).

4.3.3 Stable Isotopes

The δ¹¹B values of the vent fluids from Maka HF (26.0–27.6‰) and South (32.6–37.4‰.) are distinct and lower than in seawater (39.6‰, Table 2). Similarly, the Maka HF fluids are less radiogenic with respect to ⁸⁷Sr/⁸⁶Sr (0.70440–0.70448) compared to seawater (0.70915, Table 2). The δ²H_H2O values in the fluids from Maka HF (−0.7–0.2‰) and South (−0.92–−0.14) are comparable with respect to the 1 sigma error of the analysis (±1‰, cf. Section 3.2.4) and overlap with the seawater composition (−0.14 ± 1.0‰, Table 2). By contrast, δ¹⁸O_H2O values are distinct between the two vent sites and compared to seawater (−0.1 ± 0.1‰), as reflected by higher δ¹⁸O_H2O values in the fluids from Maka HF (0.78–1.05‰) compared to those from Maka South (−0.17–0.70‰, Table 2).

4.3.4 Water/Rock Ratios

In order to estimate the intensity of water-rock (W/R) interaction, the W/R ratios were calculated from end-member fluid compositions (Supplementary Tables S7, S8) and metal concentrations measured in fresh volcanic glass from Maka (Supplementary Table S2). The W/R ratios were calculated assuming a 100% extraction rate according to the method described by Von Damm et al. (1985b) (for more detail see Supplementary Table S12). Fluid-mobile elements, such as Li, Rb, Cs and to a lesser extent K are best suited for determining the W/R ratio, as they are affected only to a small degree by secondary reactions (Mottl and Holland, 1978; Von Damm et al., 1985a). In addition, W/R ratios were also estimated from end-member ⁸⁷Sr/⁸⁶Sr isotopic values following the dissolution first – precipitation second model after Berndt et al. (1988). At Maka HF, the calculated W/R ratios yielded similar results for K, Li, Rb, Cs and ⁸⁷Sr/⁸⁶Sr isotope ratios with values in a narrow range between 2 and 3 (Supplementary Table S10).

5 Discussion

5.1 Paragenetic Sequence and Precipitation Conditions

The formation of the vent structures at Maka volcano can be subdivided into an early low temperature (230–260°C), intermediate high temperature (300–390°C) and late low temperature stage (<240°C, Figure 9). As a consequence, the mineralogy and precipitation textures vary in the seafloor vents between the different stages reflecting variable fluid conditions (Figures 2, 9). The earliest stage of chimney formation is typically related to abundant fluid-seawater mixing leading to the formation of a sulphate-rich (e.g., barite) outer chimney wall with a high permeability for seawater ingression (Figures 2A,B; Tivey et al., 1999; Berkenbosch et al., 2012; Webber et al., 2015; Falkenberg et al., 2021; Keith et al., 2021). The formation of early stage colloform and anhedral pyrite I (Figures 2A,B), which post-dates the barite deposition is indicative for disequilibrium precipitation conditions due to strong physicochemical gradients (e.g., temperature) induced by abundant mixing between the hot fluid (290–330°C, Table 1) and cold ambient seawater (∼2°C; Berkenbosch et al., 2012; Keith et al., 2016a). The precipitation of pyrite I is accompanied by early-stage sphalerite I and marcasite I (Figures 2B–D) that likely formed at similar fluid conditions (Murowchick and Barnes, 1986; Metz and Trefry, 2000; Reed and Palandri, 2006). The Fe/Zn ratio of Cu-poor (<2 wt%) sphalerite I can be used to estimate the precipitation temperatures during the early stage (Keith et al., 2014), which are suggested to be in the range of 230–260°C for the Maka hydrothermal system (Figure 9). These temperature estimations also agree with the precipitation conditions of marcasite that typically forms at temperatures <240°C (Murowchick and Barnes, 1986; Nestmeyer et al., 2021). This early sulphide precipitation stage insulates the hot ascending fluids in the central conduit from the surrounding seawater leading to higher precipitation temperatures at the onset of the intermediate-stage (Figure 9). This is reflected by high temperature phases like chalcopyrite and euhedral pyrite II that form in the intermediate chimney wall (Figures 2D–G) and near the central fluid conduit (Figure 2H) indicating more stable precipitation conditions with minor seawater mixing (Butler and Nesbitt, 1999; Berkenbosch et al., 2012; Evans et al., 2020; Falkenberg et al., 2021). The onset of sphalerite II formation overlaps with the pyrite II and chalcopyrite precipitation (Figures 2E,G, 9) and temperature estimations based on the Fe/Zn ratio in sphalerite II yielded temperatures between 300 and 390°C (Keith et al., 2014), which are in the range of typical chalcopyrite precipitation temperatures (300–400°C; Butler and Nesbitt, 1999; Metz and Trefry, 2000; Reed and Palandri, 2006). This also agrees with the measured fluid temperatures at Maka HF (up to 329°C), whereas those from Maka South rather tend to be lower (up to 300°C) than the suggested precipitation temperature of sphalerite II (300–390°C). However, minimum concentrations of 31.0 mmol/kg Mg in the Maka South fluids (Table 2) are indicative for mixing between fluid and seawater during or prior to fluid sampling, suggesting that the endmember temperatures are >300°C and possibly in the range of the proposed sphalerite II precipitation temperatures (300–390°C), which is confirmed by the suggested endmember fluid temperatures for Maka South that reach 380°C (cf. Section 4.3.1). The formation of pyrite III overlaps with sphalerite II (intermediate-stage) and late-stage marcasite II (Figures 2G, 9); however, the precipitation of marcasite II continues beyond the pyrite III deposition implying a final decrease in fluid temperature (<240°C, Murowchick and Barnes, 1986; Nestmeyer et al., 2021). This agrees with the late stage occurrence of bornite, covellite and chalcocite-digenite (Table 3 and Figure 9) either indicating a decreasing fluid discharge resulting in seawater ingress caused by a gradient in fluid and seawater pressure in favour of the latter (Tivey, 1995) or a supergene occurrence by chalcopyrite replacement (Berkenbosch et al., 2012; Keith et al., 2016a; 2021). However, marcasite is absent in the bornite, chalcocite-digenite and covellite bearing samples from Maka South (Table 2), for which reason we propose that marcasite was not stable under the given fluid conditions probably due to elevated fluid temperatures (>240°C, Murowchick and Barnes, 1986; Nestmeyer et al., 2021). Instead, decreasing a (H₂) and a (H₂S) conditions in high temperature fluids (>300°C) could result in a similar sequence of chalcopyrite followed by bornite and eventually chalcocite/covellite as the activities of H₂ and H₂S decrease towards the outer chimney wall (Figure 10), as observed at Maka South. The formation of chalcopyrite and covellite from high temperature fluids can also be related to high-sulphidation conditions (Figure 10), where SO₂ is derived from magmatic fluids (Einaudi et al., 2003), for which we do not see evidence at Maka South (cf. Section 5.4). However, similar conditions may be achieved by the condensation of boiling-induced vapour into hydrothermal fluids or seawater (Naden et al., 2005; Nestmeyer et al., 2021; Schaarschmidt et al., 2021; Voudouris et al., 2021), as suggested for Maka South (cf. Sections 5.2, 5.3). Hence, the paragenetic relations between the different sulphides reflect variations in fluid temperature, a (H₂) and a (H₂S), as well as fluid-seawater mixing in variable proportions and the condensation of boiling induced vapour.

FIGURE 9. Paragenetic sequence of chimney evolution at Maka. Abundances estimated by optical and electron microscopy, as outlined in Table 3. The temperatures combine the in-situ measurements at the active chimneys by temperature probe (intermediate stage) and the estimated precipitation temperatures based on Fe/Zn in sphalerite I (early stage) and II (late stage, cf. Section 5.1).

FIGURE 10. Mineral stabilities in the Cu-Fe-S-O-H system for variable a (H₂) and a (H₂S) at 330°C and 250 bar. The composition of fluids from Maka South (orange circle) and Maka HF (blue circle) were depicted assuming unit activity coefficients for H₂ and H₂S. The orange arrow points in the direction of increasing sulfur fugacity (i.e., high sulphidation). The black arrows represent the suggested sequence of chalcopyrite followed by bornite and chalcocite as H₂ and H₂S decrease towards the outer chimney wall. The orange arrows represent the suggested replacement reactions of chalcopyrite – bornite followed by bornite – chalcocite at Maka South.

5.2 Fluid Processes and Three-Component Mixing

The Cl concentration of the end-member fluids at Maka HF (524–528 mmol/kg) is comparable to seawater (539 mmol/kg), indicating that fluid boiling did not affect the hydrothermal system (Supplementary Table S7; Stoffers et al., 2006; Monecke et al., 2014; Schmidt et al., 2017). This interpretation is supported by isenthalpic-isobaric end-member temperature calculations, which yield temperatures between 301°C and 345°C, placing the fluids in the single-phase field below the boiling curve at seafloor pressure (Supplementary Table S7, Figure 8). The low Mg and SO₄ concentrations in the sampled fluids from Maka HF indicate that entrainment of seawater prior to or during sampling is minor and that the fluid composition is close to the respective end-member (Table 2). The low W/R ratios of 2–3 (cf. Section 4.3.4) and overall enrichment of alkali- (e.g. Li, K, Rb), alkali-earth (e.g., Ca, Sr) and rare Earth elements (Supplementary Tables S7, S8) relative to seawater suggest a rock-dominated hydrothermal system at Maka HF. Furthermore, lower Na/Cl and higher Ca/Cl ratios in the fluids compared to seawater are indicative for Ca-release from the wall rocks caused by albitization (Berndt et al., 1989; Berndt and Seyfried, 1993). The concentration of dissolved CH₄ (24.6 and 51.8 μmol L⁻¹) is on the lower end of the observed concentration range in hydrothermal fluids from back-arcs (Kawagucci et al., 2011; Reeves et al., 2011, 2014; Seewald et al., 2015) and could either be the result of abiotic sources or microbial activity and thermogenesis of sediments and/or organic matter (Welhan, 1988; Seewald et al., 1994; Cruse and Seewald, 2006; Proskurowski et al., 2008; Fiebig et al., 2019).

The concentrations of dissolved H₂ in the hydrothermal end-member fluids at Maka HF also agree with the low W/R ratios, as H₂ is controlled by fluid-mineral equilibria of Fe-bearing sulphides, oxides and aluminosilicate minerals (Seyfried and Ding, 1995; Seyfried et al., 2003). Stable isotope systematics in the end-member fluids from Maka HF support the interpretation of rock-dominated high-temperature hydrothermal conditions, since δ²H_H2O (up to 0.4‰) and δ¹⁸O_H2O (up to 1.18‰, Supplementary Table S7 are shifted towards higher values (within analytical uncertainty) compared to seawater (−0.1‰ in δ²H_H2O, −0.2‰ in δ¹⁸O_H2O), which is consistent with seawater-basalt interaction (Shanks et al., 1995; Shanks, 2001). Furthermore, the ⁸⁷Sr/⁸⁶Sr values of the end-member fluids (0.70414) are significantly less radiogenic than seawater (0.70915, Supplementary Table S7) and calculated W/R-ratios (2) agree well with those (2–3) based on fluid mobile elements, such as K, Li, Rb and Cs (Supplementary Table S10). Likewise, the δ¹¹B values (25.12–26.25) of the end-member fluids are significantly lower than in seawater (39.6, Supplementary Table S7), indicating that the boron concentration and δ¹¹B isotope ratios in the vent fluids are controlled by the composition of the host rocks, as suggested for other arc and back-arc hydrothermal systems (Yamaoka et al., 2015; Wilckens et al., 2018). Additionally, fluids venting at Maka HF show chondrite-normalized REE patterns, which overlap with other high-temperature black smoker systems (Supplementary Figure S2; e.g., Douville et al., 2002). The observed patterns agree with REE solubilities controlled by chloro-complexation, as is expected for high-temperature systems under rock-dominated redox conditions (e.g., Craddock et al., 2010). Hence, the chemistry of the seawater-derived black smoker-type fluids from Maka HF can mainly be attributed to W/R interaction in a rock-dominated non-boiling system with only minor proportions of seawater mixing.

At Maka South, Mg concentrations range from 31.0 to 46.2 mmol/kg (Table 2) and together with the elevated pH_{(25°C,1 atm)} in the vent fluids (4.53–5.42) seawater mixing prior to or during sampling is suggested. Although the vent fluids from Maka South were sampled at p-T conditions below the boiling curve (Table 1), highly variable Cl concentrations (309–490 mmol/kg) that decrease together with the Mg content suggest a low-Cl vapour-rich component in the endmember fluids, which is likely related to subseafloor boiling and subsequent cooling by seawater mixing (Butterfield and Massoth, 1994; Von Damm et al., 1997; James et al., 2014). This is confirmed by isenthalpic-isobaric temperature calculations for end-member fluid compositions that yield temperatures of 379°C (Supplementary Table S7), which are above the boiling curve at seafloor pressure, indicating that the fluids intersected the boiling curve in the subseafloor during their ascent (Figure 8). Concentrations of Cl, Br, Na and SO₄ versus Mg in the vent fluids show a linear correlation, resulting in negative concentrations of Cl (−20.5 mmol/kg), Na (−117 mmol/kg) and SO₄ (−24.0 mmol/L) in the end-member fluids at 0 Mg (Supplementary Figure S1, Supplementary Table S7). These negative concentrations are likely an artefact of the extrapolation to zero-Mg possibly due to the non-conservative behaviour of Mg and other elements during mixing. The low SO₄ concentrations together with the low Ca and Sr contents of the Maka South fluids are indicative for anhydrite precipitation due to subseafloor mixing with entrained seawater. In addition, the negative endmember chloride concentrations also suggest that Mg was probably removed from the mixed fluids, although not as effectively as sulphate. This is consistent with hydrothermal experiments showing that sulphate decreases faster than Mg, when seawater is heated in the presence of basalt to temperatures of 200–300°C (e.g., Seyfried and Mottl, 1982).

It is likely that the Maka South had near 0 mmol/kg Cl prior to mixing and may be interpreted as a “pure” vapour phase that formed by fluid boiling in the subseafloor. Using the SoWat program, the Cl concentration of co-existing vapour and liquid can be calculated in pressure-temperature-composition (p, T, x) space for the NaCl-H₂O system (Driesner, 2007; Driesner and Heinrich, 2007). Assuming a seawater-like chlorinity similar to the black smoker-type fluids from Maka HF (Supplementary Table S7), a temperature of 379°C (cf. isenthalpic-isobaric temperature calculations) and variable pressure conditions for subseafloor (∼175 bar) boiling yield a vapour phase with ∼2 mmol/kg Cl, which agrees with the proposed end-member composition of highly Cl depleted fluids at Maka South (Supplementary Table S7) Assuming an equilibrium partitioning of a 1 kg solution with seawater chlorinity, the results also suggest that ∼91% low-Cl vapour (∼2 mmol/kg) formed by physical separation from a high-Cl liquid (∼9% at ∼6,000 mmol/kg Cl). This may also explain the δ²H_H2O values of the Maka South fluids (−0.92–−0.29), which are more negative than those from Maka HF (−0.67–0.28, Table 2). Experimental studies showed that negative δ²H_H2O values in vapour-dominated fluids can be the result of open-system boiling, if >90% of a vapour fraction is separated by boiling during isobaric heating (Shanks et al., 1995; Berndt et al., 1996). Furthermore, boiling leads to the enrichment of dissolved gases (e.g., H₂S, CO₂, CH₄) in the vapour (Drummond and Ohmoto, 1985; Butterfield et al., 1990). which explains the higher gas concentrations in the sampled fluids from Maka South compared to Maka HF (Table 2 and Supplementary Table S7). However, H₂ contents are lower in the fluids from Maka South than in those from Maka HF, although H₂ is less soluble than H₂S in the aqueous fluid and should hence be even more enriched in the vapor. H₂ may be less controlled by fluid-mineral equilibria of Fe-bearing sulphides, oxides and alumosilicate minerals at the higher W/R ratios at Maka South than at Maka HF. Instead, the compositions are shifted towards higher sulfur fugacities (Kawasumi and Chiba, 2017) (Figure 10), which could point to an influence of the reaction S_o + H₂ = H₂S. Another possibility is that H₂ partitioned so strongly into the vapour phase that it is already depleted in the boiling fluid, which still release the more soluble gaseous species (e.g., H₂S).

The high Mg contents at Maka South (31.0–46.2 mmol/kg) combined with the low chlorinity and elevated gas contents suggest mixing of a vapour-phase with seawater (Figure 11A). However, this simple mixing scenario cannot explain the enrichment of the Cl-complexed elements Si, Li, K, Ca, Mn, Fe, Co, Rb, Sr, Cs and Pb (Figures 11C,D; Berndt and Seyfried, 1990; Pokrovski et al., 2005) in the Maka South fluids relative to seawater. This observation argues against a simple two-component mixing model between a “pure” vapour phase and seawater, but rather suggests mixing of at least three components; of which one is enriched in the above mentioned Cl-complexed elements that are leached from the host rocks (Seyfried et al., 1984; Von Damm et al., 1985a; Berger et al., 1988; Berndt and Seyfried Jr, 1990; Seewald and Seyfried Jr, 1990). We propose that this third fluid component is similar in composition to the black smoker-type fluids that are discharged at the nearby vent site (<100 m) of Maka HF. This three-component mixing model is supported by the SO₄ concentrations in the Maka South vent fluids, which significantly depart from a conservative mixing line between seawater and a Mg- and SO₄-free end-member typically expected for high temperature hydrothermal systems (Supplementary Figure S1E; Mottl and Holland, 1978). Hence, we propose a mixing process between three components for Maka South resulting in a mixing triangle (Figures 11B–E) that includes 1) a black smoker-type fluid that is enriched in alkali-, alkali-earth and trace metals (e.g., Mn, Li, B) 2) a “pure” vapour phase strongly depleted in Cl and enriched in volatile components (e.g., H₂S, CO₂, CH₄) and 3) a seawater contribution (Figure 12). We suggest that the seawater-derived black smoker-type fluid is derived from the rock-dominated Maka HF system and mixed with the “boiled” vapour-rich fluids in the subseafloor (i.e., vapour sensu stricto or condensed vapour) that were derived from a distinct upflow zone. These fluids then mixed with entrained and heated seawater in the subseafloor or during venting (Figure 12) initiating the sulphide precipitation (cf. Section 5.3).

FIGURE 11. Bivariate diagrams of (A) Cl – Mg, (B) Mn – Mg, (C) Li – Cl, (D) Mn – Cl, (E) H₂ – Cl, (F) B – Cl. Linear least squares regression to 0 Mg indicate two different hydrothermal end-members. (C–F) Mixing triangle for three component mixing between (i) black smoker-type fluids similar to Maka HF, (ii) boiling-induced low-Cl vapour and (iii) seawater.

FIGURE 12. Simplified hydrothermal model for the Maka HF and Maka South vent system summarizing the key hydrothermal processes based on the fluid and sulphide data.

5.3 Effects of Fluid Composition on Sulphide Chemistry

Pyrite occurs throughout the paragenetic sequence of chimney formation (Figure 9), as indicated by texturally diverse pyrite-types that reflect the different precipitation conditions during the evolution of the Maka hydrothermal system (cf. Section 5.1). As a consequence, the trace element composition of the different pyrite-types should report on the fluid composition with respect to the different mineralization stages (Figure 9; Wohlgemuth-Ueberwasser et al., 2015; Keith et al., 2016a; Román et al., 2019; Meng et al., 2020; Falkenberg et al., 2021; Nestmeyer et al., 2021). Importantly, trace element concentrations and ratios in pyrite III commonly show a significant variation and overlap with the data of pyrite I and II (Figure 4). This is likely related to the precipitation of pyrite III that commences during the intermediate high temperature stage (300–360°C) and which proceeds into the late low temperature stage (<240°C) reflecting conditions similar to the formation of pyrite I and III, respectively (Figure 9). The effect of fluid temperature on pyrite precipitation is also preserved by the trace element composition of the different pyrite-types (Huston et al., 1995; Maslennikov et al., 2009; Keith et al., 2016b, 2018; Berkenbosch et al., 2019; Meng et al., 2020; Falkenberg et al., 2021). Pyrite I and III are typically enriched in elements (e.g., Sb, Tl, Pb; Figures 4A,E) that are suggested to precipitate at lower fluid temperatures (<250°C; Metz and Trefry, 2000; Reed and Palandri, 2006; Brugger et al., 2016). Similarly, low temperature pyrite I (∼240°C) from Niua South that formed due to abundant fluid-seawater mixing during the early stages of chimney growth shows higher Tl and Pb contents than the associated high temperature pyrite II (300–320°C; Falkenberg et al., 2021). By contrast, Se and Co that typically precipitate at elevated temperatures (>300°C Metz and Trefry, 2000; Maslennikov et al., 2009; Berkenbosch et al., 2019; Meng et al., 2020) are enriched in high temperature pyrite II (300–390°C) relative to pyrite I and III from the same vent site at Maka volcano (Figure 4B). Other trace elements like Ni, Ge and As in pyrite (Figures 4C,D,F) seem to vary irrespective of the precipitation temperature between the different mineralization stages (Figure 9). Pyrite-types from the same stage at Maka HF and South probably formed at comparable fluid conditions (e.g., temperature; cf. section 5.1), implying that variations in trace element concentrations and ratios are not only related to the seafloor precipitation processes, but are rather also the result of mixing and boiling processes that occur during fluid upflow, which are distinct between the two vent sites, as suggested by the fluid data (cf. Section 5.2; Román et al., 2019; Keith et al., 2020; Falkenberg et al., 2021).

Besides temperature, fluid chlorinity also has a significant effect on the solubility of base metals that have an affinity to form Cl-complexes in hydrothermal fluids (e.g., Co, Ni, Zn, Ag, Pb; Metz and Trefry, 2000; Reed and Palandri, 2006; Brugger et al., 2016; Pokrovski et al., 2018). By contrast, elements with a volatile character (e.g., Sb, Te, As, Tl) may also be soluble at low chlorinities, indicating that their precipitation is not primarily controlled by changes in fluid chlorinity (Grundler et al., 2013; Pokrovski et al., 2018; Román et al., 2019; Keith et al., 2020; Nestmeyer et al., 2021). Metalloids like As and Sb and soft metals (like Au) can be transported by low-Cl fluids with high H₂S, because they tend to form bisulfide complexes (e.g., Akinfiev and Zotov, 2001). Hence, variations in vent fluid chlorinity and H₂S contents induced by fluid boiling (cf. Section 5.2) affect the solubility of metals in the fluids and likely the ratio of volatile to non-volatile elements (e.g., As/Co; Pokrovski et al., 2018). The elevated contents in volatile elements like As, Se, Cd and Sb in some of the bulk sulphide-sulphate samples from Maka South may therefore be related to the low-Cl vapour-rich character of the vent fluids at this site (Figure 3). Similarly, previous studies highlighted that boiling signatures and changes in fluid chlorinity may be preserved by the trace element record of pyrite that formed from these fluids (Tardani et al., 2017; Román et al., 2019; Keith et al., 2020; Börner et al., 2021; Falkenberg et al., 2021; Nestmeyer et al., 2021). Nestmeyer et al. (2021) suggested that pyrite from boiling-related low-Cl vapour-rich fluids typically shows higher As/Co ratios (>10–100) due to the preferential partitioning of As into the vapour phase relative to Co that is enriched in the high-Cl liquid. Accordingly, pyrite from Maka South shows higher As/Co ratios (>10–100) than pyrite from Maka HF, which precipitated from black smoker-type fluids with seawater-like chlorinities (524–528 mmol/kg, Supplementary Table S7) that are characterized by lower As/Co ratios (<10–100, Figure 4F). Importantly, pyrite from Maka HF and those that formed under non-to gentle-boiling conditions from the Cerro Pabellón Geothermal System (CPGS) in Chile show comparable As/Co ratio of <10 to 100 (Figure 4F; Román et al., 2019). By contrast, pyrite from the boiling Maka South hydrothermal system overlaps in As/Co with pyrite from the CPGS that formed under vigorous boiling conditions (Figure 4F). Similar variations may be observed with respect to Sb/Pb, where pyrite from Maka South tends to show higher Sb/Pb values (>0.1) than pyrite from Maka HF, however, this is not supported by the pyrite data from the CPGS system (Figure 4E). Hence, the As/Co (and Sb/Pb) ratio may be a useful tool to distinguish pyrite and related mineralizations that formed from fluids with different chlorinities caused by boiling processes.

The fluid data from Maka South points towards a three-component mixing process involving low-Cl vapours induced by boiling, black smoker-type fluids similar to those from Maka HF and seawater (cf. Section 5.2 and Figures 11, 12). This may also be reflected by the bulk sulphide-sulphate data showing a similar compositional range at both sites, but with a volatile element enrichment (e.g., As, Se, Cd, Sb) in some of the Maka South samples, as described above (Figure 3). This is also preserved in the pyrite data from Maka HF and South, as reflected by the compositional overlap in trace element concentrations and ratios between the two vent sites (Figures 4, 5). Interestingly, low temperature pyrite I and III (<260°C; Figure 9) from Maka volcano that are affected by seawater mixing (cf. Section 5.1) tend to show the lowest Se/Ge ratios (<10, Figure 4C) and those from Maka South overlap with pyrite I from Niua South (Falkenberg et al., 2021), which also precipitated due to intense fluid-seawater mixing as also suggest by the fluid data from Maka South (cf. Section 5.2). Hence, the Se/Ge ratio in pyrite may report on mixing processes between hydrothermal fluids and seawater, also indicating that the Se/Ge ratio is not sensitive to boiling-induced fractionation, which is possibly related to the strong tendency of both elements to the vapour phase during fluid boiling likely not affecting their ratio in the fluids and associated pyrite precipitates (Pokrovski et al., 2005, 2018; Brugger et al., 2016).

In addition to the As/Co and Se/Ge in pyrite that are sensitive to fluid boiling and seawater mixing, respectively, the Co/Ni ratio may vary in response to temperature and/or salinity variations in vent fluids (Maslennikov et al., 2009; Brugger et al., 2016; Meng et al., 2020; Falkenberg et al., 2021). Although, there seems to be no clear difference in the Co/Ni value between the different pyrite-types and the two vent sites at Maka (Figure 4D), important information may be provided by the combined use of the Co/Ni and As/Co ratio (Figure 5A). Here, pyrite II from Maka HF that forms at peak temperatures (300–390°C; Figure 9) with only little effect of seawater mixing should best preserve the signature of the black smoker-type fluid, which is characterised by low As/Co (<10) and elevated Co/Ni (>10) possibly reflecting the high temperature and high-Cl (seawater-like) nature of the Maka HF vent fluids (Figure 5A). By contrast, the rather low Co/Ni in the Maka South pyrites is therefore more likely related to lower precipitation temperatures due to mixing with seawater (pyrite I and III) or the low-Cl nature of the vapour-dominated fluids at this site, as also suggested by the high As/Co value in the corresponding pyrites (>10–100; Figure 5A). This is also supported by the negative Se/Ge - As/Co correlation, where the high temperature black smoker-type fluid related pyrite II from Maka HF shows elevated Se/Ge (>10) compared to pyrite from Maka South, where the combination of the low Se/Ge (<10) and high As/Co (>10–100) are indicative for a mixing process between low-Cl vapours and seawater, as also suggested by the fluid data (cf. Section 5.2). Hence, the combined use of trace element ratios in pyrite (e.g., Co/Ni, As/Co, Se/Ge) from seafloor mineralizations (Figure 5) may provide important new insights into the complex nature of mixing and boiling processes during fluid upflow; both of which may occur at the same vent site as indicated by the fluid data from Maka South (Figure 12).

5.4 Leaching Versus Magmatic Volatile Influx: Evidence by S and Pb Isotopes

The S and Pb isotope composition of the hydrothermal sulphides from Maka HF and South indicate that most of the metals are likely leached from the host rocks in the reaction and upflow zone (Supplementary Table S5). Although, the δ³⁴S composition of the local host rock at Maka is unknown the hydrothermal sulphides show δ³⁴S values (−0.6–1.3‰) comparable to MORB (−1.9–0.8‰, Figure 6; Sakai et al., 1984; Labidi et al., 2012), which is supported by He isotope data (3He/⁴He - C/³He) of volcanic rocks and vent fluids from Maka (and NELSC) also indicating host rock compositions that are comparable to MORB (Lupton et al., 2015; Baker et al., 2019). The lack of more negative δ³⁴S values in the hydrothermal sulphides from Maka relative to MORB exclude a contribution of magmatic SO₂, as known from other subduction zone-related hydrothermal systems (Figure 6; Herzig et al., 1998; de Ronde et al., 2011; McDermott et al., 2015; Martin et al., 2020). However, CO₂ concentrations (24.2 and 33.3 mmol/kg) in the end-member fluids are elevated compared to seawater (2.26 mmol/kg, Supplementary Table S7) which is commonly related to fluid boiling, W/R interaction, or magmatic volatile influx in sediment-free hydrothermal systems (Reeves et al., 2011; Seewald et al., 2015, 2019) There is no evidence for fluid boiling at Maka HF (cf. Section 5.2) and the highest CO₂ concentrations reported for seafloor lavas in the northern Lau Basin (144 ppm or 3.28 mmol/kg; Lupton et al., 2015) can only account for ∼3–4 mmol/kg CO₂ in the end-member fluids according to the low W/R ratios of 2–3 at Maka. Therefore, we propose that the Maka hydrothermal system is affected by minor amounts of a CO₂ bearing but almost SO₂-free magmatic volatile phase. Variable degassing activity is common in volcanically active systems, which are either caused by shallow magma intrusions or permeability changes induced by tectonic movement (Tassi et al., 2013; Acocella et al., 2015). Also, independent and decoupled degassing of magmatically derived CO₂ and SO₂ likely relates to different solubilities at changing p-T and redox conditions in the magma (Gaillard and Scaillet, 2014; de Ronde and Stucker, 2015; Diehl, 2019). As a consequence, most of the reduced S that is incorporated in the seafloor sulphides is either derived from the host rocks due to leaching in the reaction and upflow zone (δ³⁴S = −1.9–0.8‰, MORB) or by thermochemical reduction of seawater sulphate (δ³⁴S ∼21.2‰; Tostevin et al., 2014). The S contribution from these reservoirs can be quantified by the two-component mixing equation introduced by Ono et al. (2007). For the calculation only the chalcopyrite data was considered (Supplementary Table S5), because only the inner chalcopyrite lining is suggested to be in equilibrium with the vent fluids (McDermott et al., 2015; Evans et al., 2020). We suggest that the contribution of seawater derived S at Maka is <14% and that >86% of the S is leached from the host rocks in the reaction and upflow zone.

Lead in hydrothermal fluids and the associated sulphide precipitates is typically derived by host rock leaching, and thus Pb isotopes of chimney sulphides provide information on the Pb contribution from the surface lavas and different basement rocks in the reaction and upflow zone (Fouquet and Marcoux, 1995; Verati et al., 1999; Zeng et al., 2017; Zhang X. et al., 2019; Falkenberg et al., 2021). Hence, variations in Pb isotope compositions either reflect a contribution from different reservoirs or are the results of radiogenic decay. However, the very young age of active hydrothermal systems compared to the half-life of the parent isotopes U and Th indicates that effects of radiogenic decay on the Pb isotope composition of hydrothermal sulphides from active seafloor vent systems can be neglected (Zhang X. et al., 2019). Although, the Pb isotope compositions of hydrothermal sulphides from Maka HF and Maka South only vary over a small compositional range, they are distinct with respect to the more radiogenic surface lavas (Figure 7). Hence, a second less radiogenic Pb source in the subsurface, potentially in the reaction zone, is required to explain the Pb isotope variation between the hydrothermal sulphides and the surface lavas (Fouquet and Marcoux, 1995; Falkenberg et al., 2021). Volcanic glass samples from seafloor lavas nearby the Maka vent systems and from adjacent segments along the NELSC revealed Pb isotope compositions that are less radiogenic (e.g., <18.95 in ²⁰⁶Pb/²⁰⁴Pb) than the Maka seafloor sulphides (Figure 7; Haase et al., submitted). Although, the elemental and isotope composition of Pb in crustal rocks beneath the Maka vents is unknown, these samples represent the best possible estimation of this “hidden” Pb reservoir. Thus, a hypothetical two-component mixing-line between the Maka surface lavas and the suggested sub-surface reservoir can be defined (Figure 7). Importantly, the mixing-line intersects the hydrothermal sulphide data indicating that the actual Pb isotope composition of the subsurface rocks that host the hydrothermal system of Maka may be in a comparable range. The combined use of elemental and Pb isotope compositions allows to estimate the Pb contribution from these two reservoirs to the Maka hydrothermal system (Verati et al., 1999). We propose that about 40–50% of the total Pb, as preserved in the seafloor sulphides, was contributed by the “hidden” reservoir in the reaction (or upflow zone). We note that the contribution of Pb derived from seawater mixing can be neglected in this respect due to the highly depleted nature of seawater (0.011 nM Pb; Diehl and Bach, 2020) compared to the volcanic rocks and hydrothermal fluids (Fouquet and Marcoux, 1995; Verati et al., 1999). Hence, the comparison of Pb isotope compositions between fresh seafloor lavas and associated hydrothermal sulphides provides important insights into possible isotope heterogeneities in the oceanic crust (Fouquet and Marcoux, 1995; Verati et al., 1999; Falkenberg et al., 2021). The combined use of S and Pb isotopes indicates that most chalcophile metals at Maka were likely leached from the host rocks and that magmatic volatiles only played a minor role in contributing CO₂, but without a significant effect on the S budget of the hydrothermal system (Figure 12).

6 Summary and Conclusion

Hydrothermal activity at Maka volcano occurs at two vent sites (Maka HF and Maka South), which are in close proximity to each other. The respective hydrothermal fluids and associated sulphide precipitates show distinct compositions that we attribute to W/R interaction in combination with (subseafloor) boiling and mixing processes. The composition of the vent fluids from Maka HF clearly indicates a W/R-dominated hydrothermal system, where dissolved metal concentrations are controlled by the chemical exchange between fluid and host rock. This interpretation is supported by S and Pb isotope data of the sulphide precipitates, also indicating that most of the metals are likely leached from the host rocks. The composition of the actively venting fluids at Maka South suggests a complex three-component mixing process during fluid ascent and discharge at the seafloor, which includes 1) a low Cl vapour phase, 2) a base metal-rich black smoker-type fluid similar to those form Maka HF and 3) a contribution by seawater, which we propose is also preserved by the distinct trace element signatures in pyrite from Maka HF and South. Besides the Cl depleted nature of the Maka South vent fluids, evidence for the influence of a vapour-rich phase is given by elevated contents of volatile elements (e.g.; As, Se Cd, Sb) in bulk sulphide-sulphate samples, as well as higher As/Co ratios (>10–100) and Sb/Pb ratios (>0.1) in pyrite from Maka South compared to Maka HF. The influence of a base metal-rich black smoker-type fluid from Maka HF likely leads to the compositional overlap in the trace element data of pyrite from Maka South and Maka HF. Furthermore, a strong seawater entrainment to Maka South is evident by high Mg contents (31.0–46.2 mmol/kg) and high pH values (4.53–5.42) in the measured fluids. Associated pyrite from Maka South shows low Se/Ge ratios (<10) and high As/Co ratios (>10–100), which we interpret to be indicative for a mixing processes between low-Cl vapours and seawater. Hence, the combined use of vent fluid and sulphide chemistry allows to define new geochemical tracers in hydrothermal sulphides that can provide important insights into understanding W/R interaction, fluid boiling and seawater mixing during the formation of fossil/extinct seafloor mineralizations that formed under temporally variable fluid conditions.

Data Availability Statement

The original contributions presented in the study are included in the article/Supplementary Material, further inquiries can be directed to the corresponding authors.

Author Contributions

LK performed parts of the fluid analyses, interpreted the fluid data, designed parts of the figures and tables, wrote parts of the manuscript; MK supervised the microanalytical study of the sulphide samples, interpreted the sulphide data, designed parts of the figures and tables, wrote parts of the manuscript; DH prepared and investigated the sulphide samples, performed some of the microanalytical analyses; CK, WB, AD, and FW sampled the hydrothermal fluids; WB modelled the mineral stabilities; KH led the research cruise, recovered the sulphide samples; CP, HS performed the S isotope analyses of the sulphides; AD, FW, CP, HS, and RG performed parts of the fluid analyses; RK supervised the LA-ICP-MS laboratory for the sulphide analyses; WB, AK, and KH conceptualized the study; all co-authors reviewed and edited the manuscript.

Funding

This work was funded by project 03G0263 of the German Bundesministerium für Bildung und Forschung (BMBF).

Conflict of Interest

The authors declare that the research was conducted in the absence of any commercial or financial relationships that could be construed as a potential conflict of interest.

Publisher’s Note

All claims expressed in this article are solely those of the authors and do not necessarily represent those of their affiliated organizations, or those of the publisher, the editors, and the reviewers. Any product that may be evaluated in this article, or claim that may be made by its manufacturer, is not guaranteed or endorsed by the publisher.

Acknowledgments

We thank Captain O. Meyer and his crew, V. Ratmeyer and the MARUM ROV team for the excellent support during the scientific work onboard RV Sonne. M. Regelous is thanked for help during the Pb-isotope measurements, C. Abe for his assistance during EPMA, and H. Brätz for her support during the LA-ICP-MS study.

Supplementary Material

The Supplementary Material for this article can be found online at: https://www.frontiersin.org/articles/10.3389/feart.2021.776925/full#supplementary-material

Footnotes

¹Haase, K. M., Schoenhofen, M. V., Storch, B., Beier, C., Regelous, M., Rubin, K. H.,et al. Effects of the Hydrous Domain in the Mantle Wedge on Magma Formation and Mixing at the Northeast Lau Spreading Centre, SW Pacific. Geochem. Geophys. Geosys. submitted manuscript.

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