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).

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.

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).

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).

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).