Introduction

Copper is an essential metal for our existing and changing economy. It is used in electric vehicle batteries, motors, wiring, and battery charging hubs (Nguyen et al. 2021). Copper is mined from several deposit types, including volcanogenic massive sulfide (VMS; Mosier et al. 2009; Tornos et al. 2015), Cu±Mo±Au porphyry (Kirkham and Sinclair 1995; Sillitoe 2010), skarn (Meinert 1992; Xie et al. 2020), iron oxide copper-gold (IOCG; Skirrow 2022), magmatic Cu-Ni (Barnes and Lightfoot 2005; Eckstrand and Hulbert 2007), sediment-hosted copper (Brown 1997, 2014; Rajabpour et al. 2017a), and polymetallic veins (Han et al. 2006; Kojima et al. 2009). The best studied and most extensively explored deposit types are the copper porphyry (69% of global Cu production), sediment-hosted copper (12% of global Cu production), and IOCG (4.7% of global Cu production) deposits, and they constitute the largest sources of copper (Richards 2016; Singer 2017; Chiaradia and Caricchi 2022). The increasing demand for copper has raised concerns regarding its future availability (Elshkaki et al. 2016) and has created opportunities for exploration and mining of smaller scale deposits that occur as polymetallic veins. The research and exploration of polymetallic Sn-Cu deposits are important since they also host critical metals, which are necessary for the functionality of emerging technologies but have supply constraints (Watari et al. 2020), such as In, Sn, W, Bi, and Li (Dill et al. 2013; Torres et al. 2019; Zhou and Wen 2021).

The Herberton Mineral Field (HMF) in Northeast Queensland, Australia, contains a variety of polymetallic Cu deposits, which occur as veins associated with Sn-W mineralization (Fig. 1). Similar deposit styles have been recognized in the Bolivian Sn belt, the Gejiu Sn district in China, and the Camborne-Redruth-St. Day Sn district in England (Sillitoe and Lehmann 2022). A common feature of these metallogenic provinces is that the mineralization is associated with fractionated granites, and displays a well-developed metal zonation, in which Sn, W, Mo, and Cu enrichment is proximal, and Zn, Pb, and Ag occur more distal to the intrusion (Taylor and Cuff 1979; Audétat et al. 2000). In these systems, the different metals may have different sources, with Sn mineralization genetically associated with reduced granitoids (Ishihara 1981; Eugster 1985; Harlaux et al. 2021a; Wu et al. 2021) and Cu with oxidized magmas (Murakami and Ishihara 2013). Tin is generally deposited at higher temperatures (~ 500 °C; Kelly and Turneaure 1970; Lehmann 2021) than Cu (~ 300–400 °C; Sun et al. 2013; Richards and Mumin 2013; Gómez et al. 2021).

Fig. 1
figure 1

a Outline of orogens in eastern Australia and the HMF. b Simplified geological map of the HMF. The age data shown in the legends are from U-Pb zircon geochronology compiled from Purdy et al. (2016), Cheng et al. (2018), and Kumar et al. (2022)

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Metals in magmatic-hydrothermal systems are generally transported under acidic to near neutral conditions as complex ions of chlorine and hydrosulfide (Liu and McPhail 2005; Etschmann et al. 2010; Richards 2011). Typically, metal deposition occurs as a result of changes in temperature, pH, oxidation state, and sulfur fugacity, which can lead to the destabilization of metal complexes and promote metal precipitation (Pokrovski et al. 2013; Schaarschmidt et al. 2021). The evolving hydrothermal conditions during ore deposition are complex, but important geochemical information can be recovered through detailed studies of mineral parageneses, mineral chemistry, and stable isotopes (e.g., sulfur and oxygen) of ore and alteration minerals (Rajabpour et al. 2017b; Li et al. 2020; Fitzherbert et al. 2021; Tashi et al. 2021).

Sulfur isotopes, especially for mineral pairs, provide a powerful geochemical tool to track the sulfur source and constrain fluid interaction and temperature of sulfide deposition (Ohmoto 1972; Seal 2006; Farquhar et al. 2010; Shanks et al. 2014; Walker et al. 2019). The δ34SV-CDT values associated with magmatic processes in mid-ocean ridge basalts from the Pacific-Antarctic ridge range from − 1.57 to + 0.06‰, and the deviation from the 0.0‰ mantle-like signature is due to mixing of depleted mantle and recycled oceanic crust (Labidi et al. 2014). In contrast, δ34S associated with sedimentary processes have a wide variation in isotopic signature and vary from − 50 to + 20‰ (Tostevin et al. 2014). These variations in sulfur isotope values are reflected in the sulfide minerals associated with ore deposits.

In this study, we investigate the hydrothermal processes that were involved in the formation of the polymetallic Baal Gammon and Isabel deposits using thermodynamic modeling in combination with studies on mineral paragenesis, mineral chemistry, and sulfur isotopes. These deposits are ideal for studying hydrothermal processes that operated during the formation of polymetallic vein deposits that consists of Sn-Cu and Sn-Zn styles of mineralization.

Geological setting

Regional geology

The Tasmanides are a series of orogenic belts that extend from South Tasmania to far North Queensland, as part of the Terra Australis Orogen (Fig. 1a; Cawood 2005). They encapsulate the geological history of eastern Australia from the Cambrian to the Triassic and record evidence for the breakup of Rodinia, the formation of Gondwana, the collision of Gondwana and Laurasia to form Pangea, and the breakup of Pangea (Glen 2005; Klootwijk 2013). The orogenic belts within the Tasmanides include the Delamerian, Thomson, Lachlan, Mossman, and New England Orogens (Fig. 1a; Rosenbaum 2018). The Delamerian Orogen is best exposed in the southern part of the Tasmanides, and it formed from the middle Cambrian to the early Ordovician (550–470 Ma), mostly as a result of contractional deformation (Flöttmann et al. 1994, 1998; Foden et al. 2020). The Lachlan Orogen (450–340 Ma) is positioned east of the Delamerian Orogen and consists of thrust fault systems developed in a turbidite-dominated sequence (VandenBerg 1999; Gray and Foster 2004; Hough et al. 2007). The Thomson Orogen (1180–495 Ma) is poorly exposed, as it lies under a cover of younger sedimentary basins, except for a series of metamorphic terranes in North Queensland that consist of Neoproterozoic to Early Paleozoic volcano-sedimentary sequences (Glen et al. 2013; Spampinato et al. 2015b; Spampinato et al. 2015a; Purdy et al. 2016; Cross et al. 2018; Dirks et al. 2021; Edgar et al. 2022a, 2022b; 2023). The Mossman Orogen temporally overlaps with the New England Orogen and encompasses convergent margin rocks that are mostly composed of marine sediments (Henderson et al. 2013) intruded by Carboniferous and Permian granitoids with associated volcanics (Champion 1991; Cheng et al. 2018; Kumar et al. 2022). The New England Orogen (375–200 Ma) is the youngest orogen and consists of multiple cycles of compression and extension (Jessop et al. 2019). The Tasmanides are extensively mineralized with precious (Au), base (Cu, Pb, and Zn) and critical metals (Kumar et al. 2022).

Geology of the Herberton Mineral Field

The HMF is situated in the Mossman Orogen and contains Au, Cu, Zn, Sn, W, Mo, and critical metal deposits (ca. 2400 lode deposits; Blake 1972). These deposits are hosted by meta-sedimentary rocks of the Hodgkinson Formation and felsic rocks of the Kennedy Igneous Association (Fig. 1b; Cheng et al. 2018; Kumar et al. 2022). The Mossman Orogen consists of a multiply deformed turbidite sequence (the Hodgkinson Formation) in the east, and fossiliferous limestone, chert, meta-basalt, and sandstone of the Chillagoe Formation in the west (Withnall and Henderson 2012; Henderson et al. 2013). The Kennedy Igneous Association comprises granitoids and volcanic rocks that have been grouped into four supersuites, based on petrological similarities (White et al. 2001), which intruded the Hodgkinson Formation. Rocks belonging to three of these supersuites occur near the Baal Gammon and Isabel deposits (Fig. 1b). The O’Briens Creek Supersuite, which is the oldest, hosts most of the Sn and W deposits and consists of highly fractionated granites with U-Pb zircon ages between ca. 365 and 317 Ma (Champion 1991; Cheng et al. 2018; Kumar et al. 2022). The Ootann Supersuite is widespread in the HMF and records U-Pb zircon ages of between 320 and 302 Ma (Champion and Chappell 1992; Cheng et al. 2018). The Almaden Supersuite mostly consists of granodiorite and is the youngest with U-Pb zircon ages at ca. 300 Ma (Champion and Chappell 1992). Two volcanic units, the Featherbed Volcanics and the Koolmoon Volcanic Group, occur near the Baal Gammon and Isabel deposits (Fig. 1b). The Featherbed Volcanics (310–275 Ma; Cheng et al. 2018) extend to the western part of the HMF and lie over the Ootann Supersuite, whereas the Koolmoon Volcanic Group (ca. 290 Ma; Kumar et al. 2022) occurs mainly in the eastern section overlying the Almaden and O’Briens Creek Supersuites (Fig. 1b).

Geology of the Baal Gammon deposit

The Baal Gammon deposit consists mainly of Cu(+Zn-In) mineralization that overprints earlier Sn and is hosted in the meta-sedimentary rocks of the Hodgkinson Formation and in the United North Australian Mine Porphyry (UNA Porphyry; Fig. 2a, b). The UNA Porphyry intruded the meta-sedimentary rocks as an irregularly shaped dike. The meta-sedimentary rocks consist of fine- to coarse-grained sandstone that dip gently to the SW with bedding-parallel fractures and shear zones that dip 20–40° to the SW. Both the meta-sedimentary rocks and the UNA Porphyry are crosscut by quartz-feldspar porphyry dikes mapped as part of the Slaughter Yard Creek Volcanics (Kumar et al. 2022; Fig. 2a, b). Regionally, the Slaughter Yard Creek Volcanics trend NNW-SSE, but locally in the Baal Gammon open pit, they trend approximately E-W. They are also affected by hydrothermal biotite and chlorite alteration.

Fig. 2
figure 2

a Simplified geological map of the Baal Gammon deposit. b Simplified cross-section showing mineralization overprinting the UNA Porphyry. c Hand-sample of the UNA Porphyry with an electron-backscatter image insert showing cassiterite (Cst) surrounded by chlorite (Chl) and quartz (Qz). d Drill core sample of chalcopyrite (Ccp) and pyrrhotite (Po) vein cross-cutting UNA Porphyry. e Simplified geological sketch map of the northern pit wall showing silica-cap (Sc) surrounding the UNA Porphyry and the supergene (Sg) zone with an insert image of copper carbonates

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Three stages of mineralization are recognized in the Baal Gammon deposit. Stage I consists of Sn mineralization hosted within the greisenized UNA Porphyry as hydrothermal cassiterite (Fig. 2c). Stage II consists of Cu, Zn, and In mineralization in the form of massive to semi-massive sulfides, sulfide infill breccia, sulfide veins (Fig. 2d), and disseminated sulfides that are located in the UNA porphyry and along the contact of the UNA Porphyry and the meta-sedimentary rocks. Stage III mineralization consists of supergene alteration of the sulfide mineralization and occurs as Cu-Fe oxides and carbonates (Fig. 2e).

Geology of the Isabel deposit

The Isabel deposit contains early Sn that is overprinted by Zn-Pb-Ag-In mineralization in the form of polymetallic veins hosted by brecciated meta-sedimentary rocks of the Hodgkinson Formation (Fig. 3a, b). Fault-controlled breccia surrounds the deposit and consists of angular meta-sandstone clasts with a jigsaw-fit texture within a fine-grained quartz-feldspar matrix. Dolerite and felsic porphyry dikes, mapped as part of the Slaughter Yard Creek Volcanics, intruded the meta-sedimentary rocks. The hydrothermal breccia surrounding the polymetallic veins consists of angular meta-sandstone clasts with a jigsaw-fit pattern, with fine-grained sulfide mineralization as infill (Fig. 5d).

Fig. 3
figure 3

a Simplified geological map and b cross-section of the Isabel deposit. c, d Drill core sample of sphalerite (Sp) and galena (Gn) veins with an insert of a photomicrograph in reflected light of cassiterite (Cst) surrounded by sphalerite (Sp) and chalcopyrite (Ccp)

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The polymetallic vein contains an early stage of Sn in the form of cassiterite (stage I), overprinted by a later stage (stage II) of Zn-Pb-In mineralization as sulfides (Figs. 3c and 4b). The stage I cassiterite was observed mostly as inclusions in sphalerite veins. Stage II sulfide mineralization consists of SW-NE oriented, large (10 m long and 1–2 m wide), sphalerite-rich veins, which were crosscut by thin chalcopyrite-pyrrhotite veins (Fig. 5e), sphalerite and galena veins with stannite (Fig. 5f), and hydrothermal breccia with sphalerite infill.

Fig. 4
figure 4

Mineral paragenetic sequences for the: a Baal Gammon deposit and b Isabel deposit

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Fig. 5
figure 5

Photomicrographs of ore samples from the Baal Gammon (a–c) and Isabel (d–f) deposits under reflected light. a Chalcopyrite (Ccp) and pyrrhotite (Po) with sphalerite (Sp) inclusions in massive sulfides at the contact of quartz-feldspar porphyry and UNA Porphyry. b Exsolution lamellae of pyrrhotite in chalcopyrite in sulfide veins located in the UNA Porphyry. c Framboidal pyrite (Py) in massive sulfides from the meta-sedimentary rocks. d Sphalerite (Sp) and quartz (Qz) in a hydrothermal breccia. e Sphalerite with thin veins of chalcopyrite and pyrrhotite. f Galena (Gn) vein with sphalerite and stannite (stn)

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Materials and methods

Samples

The Baal Gammon samples (Fig. 2a, b) were taken from sulfide veins in the deeper sections of the UNA Porphyry (BIN035), from massive sulfides located at the contact between a quartz-feldspar porphyry dike and the UNA Porphyry (BIN019), and from massive sulfide mineralization that occurs at the contact between meta-sedimentary rocks and the UNA Porphyry (BIN005, BIN030, and BIN042). The Isabel samples (Fig. 3b) were taken from sphalerite veins (IN013 and IN049) and from a hydrothermal breccia (IN024) with sphalerite infill. Trace elements were analyzed for sulfides from all samples, whereas sulfur isotopes were analyzed for chalcopyrite-pyrrhotite pairs from samples BIN035, BIN005, BIN019, and IN013.

Laser-ablation inductively coupled plasma mass spectrometry (LA-ICP-MS)

Laser-ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) was used for trace element analyses of in-situ sphalerite, chalcopyrite, pyrrhotite, stannite, galena, and arsenopyrite ore assemblage from the Baal Gammon and Isabel deposits. Laser analyses were performed using an Analyte G2 Excimer Laser Ablation system coupled to a Thermo iCap-TQ ICP-MS instrument at the Advanced Analytical Centre, James Cook University. Laser ablation of the sulfides was carried out with an energy density of 2 J/cm2 at a rate of 5 Hz on a spot size of 40 μm. NIST610 was analyzed and used as a reference standard, whereas MASS-1 and GSE-1 standards were used as external standards to monitor matrix effects. A mixture of helium and argon was used as the carrier gas for transporting the ablated analyte into the ICP-MS analysis chamber. A total of 39 isotopes were analyzed during spot analyses, and the major element concentrations (e.g., Cu for chalcopyrite and Zn for sphalerite) measured through electron microprobe. Kumar et al. (2023) was used as an internal standard during data reduction to yield quantitative element concentrations. The isotopes 113In and 115In were measured and the isobaric interference between 115In and 115Sn were monitored, and a correction factor was applied. The LA-ICP-MS trace element data was reduced using Iolite software (Paton et al. 2011; Fisher et al. 2017), and the laser peaks were inspected for spikes, which may indicate mineral inclusions. Summaries of the trace element results are provided in Tables 1 and 2.

Table 1 A summary of trace element concentrations (ppm) for sulfide minerals from the Baal Gammon deposit
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Table 2 A summary of trace element concentrations (ppm) for sulfide minerals from the Isabel deposit
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Secondary ion mass spectrometry (SIMS)

The secondary ion mass spectrometry (SIMS) analysis of the in-situ chalcopyrite-pyrrhotite mineral pair for sulfur isotope ratios (32S, 33S, and 34S) was performed on a CAMECA IMS1280 ion microprobe at the Centre for Microscopy Characterisation and Analysis at the University of Western Australia. Representative chalcopyrite and pyrrhotite ore samples from the Baal Gammon and Isabel deposits were mounted within 5 mm of the central portion of an epoxy mount that was cut at 1 cm from the edge to accommodate S isotope reference materials. The mount was thoroughly cleaned using distilled water and ethanol, then dried and coated with a 10-nm-thick gold layer.

During S isotope analysis, the sample surface was sputtered over a 10 × 10 μm area with a 10 kV, Gaussian Cs+ beam at ~ 2 nA intensity and a total impact energy of 20 keV. Secondary ions were sputtered in the double focusing mass spectrometer within a 110 μm entrance slit and focused in the center of a 3000 μm field aperture (× 100 magnification). Energy filtering was performed using a 40 eV band pass, and a 5 eV gap was set toward the high-energy boundary. All S isotopes were collected simultaneously in Faraday cup detectors fitted with 1010 Ω (L’2, 32S) and 1011 Ω (L1, 33S and H1, 34S) resistors, operating at a mass resolution of ~ 2500 amu. The 32S1H and 33S peaks were not completely resolved under these conditions; thus, the magnetic field was offset slightly to the low-mass boundary to avoid interference from 32S1H on 33S peaks. The magnetic field was regulated using a NMR control. Each analysis included a pre-sputtering measurement over a 15 × 15 μm area for 30 s, and automated centering of the secondary ions in the field aperture, contrast aperture and entrance slit. Each analysis then consisted of 20 four-second cycles. The analytical session was monitored for drift using two bracketing standards after analysis of every five samples. Instrumental mass fractionation (IMF) was corrected using a matrix matched reference material (Nifty for chalcopyrite and Alexo for pyrrhotite, and reference values can be found in LaFlamme et al. (2016)). The used IMF correction procedures are detailed in LaFlamme et al. (2016) and Kita et al. (2009).

Geochemical modeling

Geochemical modeling of metal complexes and sulfide precipitation was conducted with The Geochemist’s Workbench (GWB) Community Edition 13 (Bethke et al. 2022) software to investigate the: (1) stability of metal chlorides over a range of temperature and pH conditions and (2) sulfide ore formation as a function of sulfur fugacity (fS2) and pH. The thermodynamic parameters (ΔGf and Keq) for modeling were taken from the default thermodynamic database contained in the GWB software. The metal chloride species that were modeled included Cu, Zn, Fe, and In complexes, with concentrations (10−6 M) similar to those observed in active hydrothermal systems (e.g., Seward et al. 2014; Simmons et al. 2016). The fS2-pH phase diagrams for the sulfide ores were modeled for Cu, Zn, Fe, Co, and S between 0 and 300 °C. The modeling was focused on metal chloride complexes due to low-sulfidation character of these deposits, and the method was modified from Kumar et al. (2023) including temperature constraints.

Results

Petrographic observations

The mineral paragenesis (Fig. 4) and sulfide ore petrography (Fig. 5) shows that there are two types of chalcopyrite-pyrrhotite mineral pairs at Baal Gammon with only one type at Isabel. At Baal Gammon, the chalcopyrite-pyrrhotite mineral pairs include massive chalcopyrite in equilibrium with pyrrhotite (Fig. 5a), and chalcopyrite with pyrrhotite exsolution lamellae (Fig. 5b). Massive chalcopyrite in equilibrium with pyrrhotite was observed along the contact between the UNA Porphyry and the quartz-feldspar porphyry, and the meta-sedimentary rocks. Sulfide ore along the contact between the meta-sedimentary rocks and the UNA Porphyry contains minor amounts of pyrite, which consists of a framboidal core, likely of sedimentary origin, surrounded by an euhedral rim overgrowth (Fig. 5c). The chalcopyrite grains with pyrrhotite exsolution lamellae are located in massive sulfide veins that crosscut the UNA Porphyry (Figs. 2d and 5b). Both types of chalcopyrite-pyrrhotite mineral pairs contain subhedral inclusions of sphalerite and cassiterite, with trace levels of arsenopyrite. Late arsenopyrite veins crosscut the earlier chalcopyrite veins and massive sulfide mineralization. In the Isabel deposit, chalcopyrite and pyrrhotite mineral pairs occur in thin veins that crosscut sphalerite mineralization (Fig. 5e).

Trace element composition of sulfides

Chalcopyrite, pyrrhotite, and sphalerite

The concentrations of trace element in chalcopyrite from both deposits are similar (Tables 1 and 2), except for Se, Cd, and In, which are distinctly higher at Baal Gammon. Minor variability was observed in the Ti, V, Co, Ga, Ge, Se, and Sb content of chalcopyrite grains from Baal Gammon. The average concentrations are 153 ppm for Ti, 0.4 ppm for V, 2.6 ppm for Co, 3.1 ppm for Ga, 4.5 ppm for Ge, 51 ppm for Se, and 2.1 ppm for Sb (Table 1, Fig. 6a–f). Silver, In, and Sn show large variations (Table 1; Fig. 6), with average concentrations of 609 ppm for Ag, 1194 ppm for In, and 1410 ppm for Sn. The average trace element content of Isabel chalcopyrite is 139 ppm for Ti, 0.7 ppm for V, 0.3 ppm for Co, 4.3 ppm for Ga, 5 ppm for Ge, 17 ppm for Se, 2302 ppm for Ag, 26 ppm for Cd, 725 ppm for In, 1042 ppm for Sn, and 6.6 ppm for Sb (Table 2, Fig. 6g–l).

Fig. 6
figure 6

Box and whiskers plot of trace elements (Ti, Ga, Ge, Se, Co, and Ag) concentrations (ppm) of chalcopyrite, pyrrhotite, sphalerite, and stannite from the Baal Gammon (a–f) and Isabel (g–l) deposits

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The average content of trace elements for pyrrhotite is generally lower when compared to chalcopyrite except for Co (Table 2; Fig. 6). The Ti, V, Co, Ge, and Se contents in pyrrhotite from Baal Gammon show little variability, whereas the Ag, In, and Sn concentrations are highly variable. The average concentrations of these elements in pyrrhotite are 145 ppm Ti, 0.6 ppm V, 118 ppm Co, 6.8 ppm Ge, 30 ppm Se, 46 ppm Ag, 18 ppm Cd, 815 ppm In, and 493 ppm Sn (Table 1, Fig. 6a–f). The average concentration of trace elements in pyrrhotite from the Isabel deposit is 154 ppm Ti, 1.5 ppm V, 20 ppm Co, 7 ppm Ge, 6.8 ppm Se, 79 ppm Ag, 6.3 ppm Cd, 290 ppm In, 1253 ppm Sn, and 30 ppm Sb (Table 2).

The sphalerite grains from both deposits have similar trace element concentrations and show significant enrichment in Cd, In, and Sn. The Baal Gammon sphalerite inclusions in the massive sulfide ores contain, on average, 2702 ppm In, 2717 ppm Cd, and 470 ppm Sn, whereas sphalerite grains from the Isabel deposit contain, on average, 2695 ppm Cd, 1296 ppm In, and 1747 ppm Sn. The average trace element content of Baal Gammon sphalerite is 111 ppm Ti, 0.3 ppm V, 56 ppm Co, 3.4 ppm Ga, 1.1 ppm Ge, 27 ppm Se, 58 ppm Ag, and 6.1 ppm Sb (Table 1, Fig. 6a–f), whereas the average trace element content for sphalerite from Isabel is 104 ppm Ti, 0.8 ppm V, 18 ppm Co, 18 ppm Ga, 1.0 ppm Ge, 4.2 ppm Se, 45 ppm Ag, and 7.4 ppm Sb (Table 2).

Other sulfides

Other sulfide grains that were analyzed for trace element content included arsenopyrite from Baal Gammon and galena and stannite from Isabel. The arsenopyrite (Table 1) is enriched in Co (45–8607 ppm, avg. 1757 ppm) and Sb (538–1168 ppm, avg. 986 ppm). The galena from Isabel (Table 2) is enriched in Ag (389–4920 ppm, avg. 895 ppm) and Sb (416–1447 ppm, avg. 805 ppm). Stannite from Isabel (Table 2) is enriched in Ag (1618–6070 ppm, avg. 2322 ppm) and In (2427–4620 ppm, avg. 3184 ppm).

Sulfur isotope composition of chalcopyrite and pyrrhotite

A summary of the sulfur isotope data for the Baal Gammon and Isabel deposits is presented in Fig. 7. Sulfur isotope analyses for chalcopyrite from the Baal Gammon and Isabel deposits returned δ34S values ranging from + 0.54 to + 1.91‰ (n = 52) and + 0.91 to + 1.45‰ (n = 18), respectively. The 95% uncertainty confidence level for a single spot analysis of chalcopyrite is 0.04‰, and 0.07‰ for pyrrhotite.

Fig. 7
figure 7

Sulfur isotope box and whiskers plot of chalcopyrite (Ccp) and pyrrhotite (Po) from the Baal Gammon and Isabel deposits

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At the Baal Gammon deposit, the δ34S signature of chalcopyrite from massive sulfide located along the contact between the UNA Porphyry and the quartz-feldspar porphyry yields values of 0.99–1.91‰ with an average of 1.71 ± 0.54‰ (n = 16). The chalcopyrite grains from sulfide veins cross cutting the UNA Porphyry, and from the contact zone between meta-sedimentary rocks and the UNA Porphyry yield similar isotope values, which range from 1.46 to 1.74‰ (one outlier at 0.5‰) with an average of 1.63 ± 0.16‰ (n = 21). Chalcopyrite grains from the Isabel deposit yield δ34S values between 0.91 and 1.45‰ with an average of 1.30 ± 0.29‰ (n = 18).

The values of δ34S in pyrrhotite grains from the Baal Gammon and Isabel deposits are similar and range from 1.35 to 2.48‰ (n = 51) for the Baal Gammon deposit, and 1.12 to 2.11‰ (n = 17) for the Isabel deposit. The values of δ34S in pyrrhotite grains from massive sulfide along the contact between the UNA Porphyry and quartz-feldspar porphyry range from 1.68 to 2.21‰ with an average of 1.94 ± 0.26‰ (n = 16). The δ34S values of pyrrhotite from sulfide veins in the deeper section of the UNA Porphyry vary from 1.42 to 2.48‰ with an average of 1.79 ± 0.43‰ (n = 20). The δ34S values in pyrrhotite from the contact zone between meta-sedimentary rocks and the UNA Porphyry range from 1.12 to 2.11‰ with an average of 1.77 ± 0.50‰ (n = 17).

pH, Eh, and fS2 modeling

The results obtained from geochemical modeling are presented in Fig. 8 as metal chloride and ore mineral stability field diagrams. The modeling indicates that chlorine complexes of Cu, Zn, and Fe at 300 °C are stable at pH ranges between 0 and 6, which is in agreement with pH and metal concentration values measured in geothermal waters of New Zealand (Simmons et al. 2016). In addition, In speciation is strongly linked to the pH of the hydrothermal fluids, in which In chlorine complexes were found to be stable only under acidic conditions at pH < 2 (Fig. 8a; Kumar et al. 2023). Modeling of the ore minerals indicates that chalcopyrite formed at a higher fS2 compared to pyrrhotite and sphalerite, which confirms the paragenetic relationships observed at Baal Gammon and Isabel (Fig. 8b).

Fig. 8
figure 8

a Stability field of metal-chloride complexes at 0–300 °C between pH 0 and 8. The metal-Cl decay over temperature and pH was constructed in The Geochemist’s Workbench (Bethke et al. 2022) at conditions: [Cu2+] = 10−5, [Zn2+] = 10−5, [Fe2+] = 10−5, [In3+] = 10−5, [Cl−] = 0.1, pressure = 1 bar and H2O at unity. b Log fS2 vs. pH plot for the key sulfide minerals at the Baal Gammon and Isabel deposits. The fS2 plot was constructed using The Geochemist’s Workbench (Bethke et al. 2022) at conditions: [Cu2+] = 10−5, [Zn2+] = 10−5, [Fe2+] = 10−5, [Co2+] = 10−5, [SO42−] = 10−5, temperature = 300 °C, pressure = 1 bar and H2O at unity. The black, red, and blue boundary lines define Cu, Fe, and Zn species, respectively

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Discussion

The incorporation of trace elements in sulfide ores and temperature of ore formation

Trace element in hydrothermal sulfide grains can reside as solid and fluid inclusions, lattice-bound, and or absorbed constituents in the mineral structure (Fan et al. 2022; Lehrmann et al. 2022). It can be difficult to differentiate the residence state of trace elements in minerals, but inferences can be made from trace element variability. High variability on spot analyses indicates the presence of mineral or fluid inclusions. In contrast, a good element-element correlation indicates lattice-bound substitution. At Baal Gammon, chalcopyrite is the major Cu mineral and contains significant quantities (> 1000 ppm) of Zn (avg. 2161 ppm and SD at 1211 ppm), Ag (avg. 609 ppm and SD at 172 ppm), In (avg. 1194 ppm and SD at 122 ppm), and Sn (avg. 1410 ppm and SD at 350 ppm). There is little compositional variability in these elements (Table 1), which indicates lattice-bound substitution in most analyzed sulfide minerals.

The substitution of Zn, Sn, Ag, and In in chalcopyrite depends upon a number of factors including the ionic radii of the elements and the temperature and physicochemical conditions of the mineralizing fluids. Chalcopyrite from some VMS deposits in Eastern Australia shows substitution of Zn, In, and Sn for Fe, and Ag for Cu (Huston et al. 1995). The ionic radii of Zn, In, and Sn are within the acceptable range for substitution of Fe3+ and Fe2+ in chalcopyrite (George et al. 2018). Moderately strong correlations of Ag, In, and Sn with Fe (Fig. 9a) are likely indicative of coupled substitutions (Fig. 9a). Apart from the valency of Fe (3+ and 2+), the valence state of Cu in chalcopyrite is also debatable and both Cu2+Fe2+(S2-)2 and Cu+Fe3+(S2-)2 can occur (Todd et al. 2003; Pearce et al. 2006), but Li et al. (2013) suggested that intermediate valence states are likely favorable. Chalcopyrite with an intermediate valence state (Klekovkina et al. 2014) is likely present at Baal Gammon, where trivalent In and tetravalent Sn substitutes divalent and trivalent Fe, whereas the monovalent Ag most likely substitutes monovalent and divalent Cu.

Fig. 9
figure 9

Element correlation plots for chalcopyrite from Baal Gammon (a, b) and galena from Isabel (c, d). a Ag+In+Sn vs. Fe; b Zn vs. Fe; c Sb vs. Ag; d Sn vs. Ag; and e an example of time-resolved depth profile showing Zn spike related to sphalerite inclusions in Baal Gammon chalcopyrite. Element concentration on correlation plots is in molality (m)

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The elements, Zn, Ag, In and Sn, when present in high concentrations (> 10,000 ppm), occur mainly as micron-sized inclusions (10–200 μm) of stannite and sphalerite that can be detected by scanning electron microscope (Kumar et al. 2022). The lack of correlation on Zn vs. Fe plot (Fig. 9b, e), highlights the occurrence of Zn as sphalerite inclusions. The average Cd concentration of Baal Gammon sphalerite is 2717 ppm with a low variability (Table 1) indicating a lattice-bound residence. The temperature and redox conditions of Cd substitution is poorly understood, and Cd2+ when substituted for Zn2+ leads to crystallographic deformation of the sphalerite unit cell as Zn2+ has an ionic radius of 0.74 Å, whereas Cd2+ has 0.95 Å (Wright 2009; Babedi et al. 2019, 2021). According to Goldschmidt’s rules, trace element substitution in crystal lattice is preferred for elements with < 15% difference in ionic radius (e.g., George et al. 2018; Zhang et al. 2021), but the difference in ionic radius of Cd2+ and Zn2+ is > 15%, and this substitution pathway may cause deformation in the mineral structure (Babedi et al. 2019). No sphalerite deformation was observed in Baal Gammon massive sulfides, and crystal habits are mostly euhedral or partially terminated (Fig. 5a–c), which indicates a lack of crystallographic deformation, and the substitution of Fe2+ by Cd2+ was favored. Other trace elements that are anomalous in the sulfides are Co and Cd, and they have higher affinity for pyrrhotite and sphalerite, respectively, than for chalcopyrite. The ionic radius of Co (Co2+ at 0.72 Å) is within range for substitution with Fe (Fe2+ at 0.74 Å; Dehaine et al. 2021), but only small amounts of Co (avg. 2.6 ppm) were detected in chalcopyrite (Fig. 6). In contrast, pyrrhotite contains an average of 118 ppm Co, where it likely substitutes for Fe.

At the Isabel deposit, sphalerite grains have high Cd, In, and Sn concentrations (Table 2). The Cd and In show low variability (Table 2) and are likely lattice bound. Tin is both lattice-bound and occurs as solid inclusions of cassiterite and or stannite since it has a higher variability (Table 2). The incorporation of Cd, In, and Sn in the sphalerite structure has been previously reported (Ingles and Mavrogenes 2021; Xu et al. 2021; Zhang et al. 2022), with In3+ substitution suggested to occur coupled with Cu+ (e.g., Cook et al. 2012; Bauer et al. 2019). The galena is cogenetic with sphalerite and formed in the early-stage II sulfide ores and contains high concentrations of Ag and Sb. The Sb vs. Ag plot shows a linear correlation (Fig. 9), indicating the presence of coupled substitution. Coupled substitution of Ag+ and Sb3+ for 2Pb2+ has been confirmed by multiple studies on trace elements in galena (e.g., Sharp and Buseck 1993; George et al. 2015; Liu et al. 2021). A linear trend is also seen on the Sn vs. Ag plot indicating the possibility of Sn incorporation in galena via coupled substitution. Tin substitution in galena has not been reported in the literature but our data indicates a possibility and Sn4+ has a similar ionic radius to Sb3+ (George et al. 2016), which further supports this substitution despite the difference in charge between the two ions. Possible mechanisms for this include coupled substitution of Sn4+ and 2Ag+ for 3Pb2+, or Sn2+ and 2Ag+ for 2Pb2+, or direct substitution of Pb2+ by Sn2+.

Evolution of ore fluids at the Baal Gammon and Isabel deposits

The trace element composition of the major sulfides together with thermodynamic modeling of the Baal Gammon and Isabel deposits can be used to deduce the evolution of the mineralizing fluids. In both deposits, the ore textures indicate a hydrothermal origin for the mineralization and most metals discussed so far are transported as metal chloride complexes in natural hydrothermal systems at very low concentrations (Hedenquist et al. 1993). Our physicochemical modeling suggests that Cu, Zn, and Fe chlorine complexes are favored at pH ≤ 6 between 200 and 300 °C (Fig. 8a), which is consistent with observations from active hydrothermal systems (e.g., Simmons et al. 2016; Grant et al. 2020), whereas In chloride complexes are only stable at pH < 3. The incorporation of large amounts of In in chalcopyrite indicates that, at the time of chalcopyrite formation, the pH of the fluid was low (pH < 3) according to the physicochemical modeling (Fig. 8). Thus, chalcopyrite crystallization condition can be constrained to ~ 300 °C, at pH < 3 and high sulfur fugacity (Fig. 8). At the Baal Gammon deposit, textural evidence suggests that pyrrhotite and sphalerite are cogenetic with chalcopyrite, but their crystallization may have occurred in sequence as the fluid chemistry evolved (Fig. 8b). The sequence of sulfide crystallization likely occurred in this order, from first to last, chalcopyrite followed by pyrrhotite than sphalerite as inclusions, according to ore petrology (Fig. 5a–c) and geochemical modeling (Fig. 8b). The geochemical modeling indicates that during sulfide ore formation, the mineralizing fluid evolved toward a moderate pH and low sulfur fugacity (Fig. 8b). Based on element substitutions and the possibility of existence of multiple valence state of Cu and Fe in chalcopyrite (Cu2+Fe2+(S2-)2 vs. Cu+Fe3+(S2-)2), the redox conditions of the mineralizing fluids likely fluctuated along the redox boundary of Fe2+/Fe3+ and Cu+/Cu2+.

At the Isabel deposit, the mineral paragenesis and cross-cutting relationships indicate that sphalerite and galena formed during early-stage II mineralization and were crosscut by late, thin chalcopyrite and pyrrhotite veins (Figs. 4b and 5e). This suggests that stage II mineralization occurred by two distinct pulses of sulfidic fluid. The first pulse of the fluids was dominated by Zn and Pb that formed the large sphalerite and galena veins and occurred at low sulfur fugacity and moderate pH (Fig. 8b). The second pulse of fluids formed the late-stage II mineralization consisting of thin chalcopyrite and pyrrhotite veins (Figs. 4b and 5e) and records increasing sulfur fugacity and decreasing pH trends (Fig. 8b).

Sulfur and metal sources at the Baal Gammon and Isabel deposits

The texture of the sulfide mineralization (Figs. 2 and 5) together with sulfur isotopes (Figs. 7 and 10) of chalcopyrite and pyrrhotite can be used to make inferences about the source of sulfur during mineralization. At Baal Gammon, the variation in sulfur isotopes in chalcopyrite (+ 1.0 to + 1.9‰) and pyrrhotite (+ 1.4 to + 2.5‰) is small. Chalcopyrite, pyrrhotite, and pyrite from VMS and IOCG deposits show a wide variation in sulfur isotopes (5 to 20‰), which has been attributed to mixing of magmatic, connate, and meteoric fluids (Brueckner et al. 2015; Li et al. 2021). Constrained sulfur isotope signatures were also observed in the cross-cutting chalcopyrite (+ 1 to + 1.5‰) and pyrrhotite (+ 1.1 to + 2.1‰) veins of the Isabel deposit. The Isabel deposit is dominated by sphalerite polymetallic veins and is distally related to the fluid source. The narrow sulfur isotope ranges of chalcopyrite and pyrrhotite from the Baal Gammon and Isabel deposits indicate a magmatic origin of the mineralizing fluids with minimal interaction with connate and meteoric waters. The magmatic source of the sulfur-bearing fluid is further supported by the sulfide mineralization being spatially and temporarily associated to the magmatic activity of the Slaughter Yard Creek Volcanics in the HMF (Kumar et al. 2022).

Fig. 10
figure 10

δ-δ plots for chalcopyrite and pyrrhotite mineral pairs. a The Baal Gammon samples show disequilibrium fractionation of sulfur isotopes in massive chalcopyrite and pyrrhotite (Ccp-Po-massive) and chalcopyrite with pyrrhotite exsolution (Ccp-Po-exsolution). b The Baal Gammon sample containing framboidal pyrite (Ccp-Po-framboidal-Py) shows modified equilibrium fractionation, whereas the Isabel samples of chalcopyrite and pyrrhotite veins (Ccp-Po-vein) plot on the kinetic fractionation line of 300 °C from Kajiwara and Krouse (1971)

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If the fractionation process was in equilibrium and kinetically driven, then the δ34S of sulfide mineral pairs can be used to predict the temperature of formation (Kajiwara and Krouse 1971). An attempt was made (Fig. 10a), but it did not yield expected temperatures for the Baal Gammon deposit. For the Isabel deposit, the temperature predicted by the fractionation of δ34S between chalcopyrite and pyrrhotite is ~ 300 ± 40 °C (Fig. 10b). The temperature of formation for the Baal Gammon deposit (290 ± 12 °C) from stannite-sphalerite thermometer obtained in Kumar et al. (2023) is very similar to the Isabel deposit. This confirms that sulfur isotope fractionation in the Baal Gammon massive sulfides occurred in disequilibrium since they do not plot on the kinetic fractionation line (Fig. 10). This disequilibrium is not surprising, especially in massive sulfide veins from the deeper parts of the UNA Porphyry (Fig. 5b), as chalcopyrite and pyrrhotite did not precipitate simultaneously, but pyrrhotite exsolved along the grain boundaries of chalcopyrite (Fig. 5). Isotopic disequilibrium caused by paragenetic mineral phase differences is widely recognized (Rye 2005) and could explain the sulfur isotope fractionation observed in massive sulfides at the Baal Gammon deposit.

The sulfur isotope pattern from the sedimentary section of the Baal Gammon massive sulfides shows a unique fractionation pattern between chalcopyrite and pyrrhotite (Fig. 10b), which is not compatible with kinetic or mass-dependent fractionation of stable isotopes. Sulfur isotopes can be fractionated by microbial activity to varying degrees, from 2 to 42‰, where the degree of fractionation depends upon the culture of microbes present (Detmers et al. 2001). The presence of framboidal pyrite (Fig. 5c) in the massive sulfides occurring in the meta-sedimentary rocks of the Baal Gammon deposit indicates the presence of microbial activity in the meta-sedimentary rocks. In sedimentary environments, microbe-derived framboidal pyrite has been used to approximate paleo-redox conditions (Chen et al. 2022). In our case, the framboidal pyrite is early, syn-sedimentary, and is overgrown by euhedral pyrite (Fig. 5c), indicating that they may not provide an accurate prediction of redox conditions associated with ore formation, but constitutes a local source of sulfur during the formation of massive sulfide ores at the contact between the UNA Porphyry and meta-sedimentary rocks. Thus, the sulfur isotope fractionation from massive sulfides at the meta-sedimentary rocks contact can be explained by interaction of fluids with sulfur derived from framboidal pyrite.

The low variation in the sulfur isotope data indicates that the bulk of the sulfur was derived from magmatic sources with only a negligible fraction derived from pre-existing syn-sedimentary framboidal pyrite. Magmatic activity that was cogenetic with sulfide mineralization occurred at ca. 288 Ma at Baal Gammon and ca. 292 Ma at Isabel, and regionally, this pulse peaked at ca. 285 Ma (Kumar et al. 2022). In the eastern part of the HMF, this magmatic pulse formed the Koolmoon Volcanic Group (Fig. 1) that consists of the Glen Gordon, Walsh Bluff, and Slaughter Yard Creek Volcanics near the Baal Gammon and Isabel deposits (Figs. 2 and 3). The trace element data from Cheng et al. (2018) indicate considerable amounts of Cu (~ 10 ppm), Zn (~ 10 ppm), W (~ 10 ppm), and Pb (~ 10 ppm) in the Slaughter Yard Creek Volcanics (Fig. 11b). This volcanic group plots above the dividing boundary of the magnetite series granite on the Fe2O3/FeO vs. Rb/Sr plot (Fig. 11a; Blevin et al. 1996; Baker et al. 2005). Therefore, these units were derived from a subvolcanic source with a strongly oxidized signature indicating a high potential for Cu-Mo and Mo mineralization (Blevin et al. 1996; Baker et al. 2005). Oxidized volcanic and felsic rocks associated with Cu mineralization indicate a likely magmatic origin for mineralizing fluids (Hattori and Keith 2001; Richards 2015). The presence of the Baal Gammon and Isabel deposits near oxidized magmatic and metal enriched rocks indicates a genetic association of Cu mineralization to the ca. 290 Ma Slaughter Yard Creek Volcanics unit (Kumar et al. 2022), whereas Sn is linked to the reduced 365–317 Ma O’Briens Creek Supersuite in the region (Fig. 11; Champion 1991; Cheng et al. 2018; Kumar et al. 2022).

Fig. 11
figure 11

a Oxidation and fractionation states of Slaughter Yard Creek Volcanics and O’Briens Creek Supersuite with geochemistry data from Cheng et al. (2018). The plot was adopted from Voudouris et al. (2022) and compiled from Baker et al. (2005), Blevin (2004), and Blevin et al. (1996). b Trace element concentration line plot of Slaughter Yard Creek Volcanics with data from Cheng et al. (2018)

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Possible genetic model of Cu mineralization in the Herberton Mineral Field

Tin-Cu deposits occur in various tectonic settings, for example, the deposits in Cornwall, England, formed during an extensional phase in a collisional orogen, the San Rafael deposit in Peru formed in a subduction-related back-arc setting, the deposits in the Erzgebirge, Germany, formed in a post-collisional extensional zone, and deposits in Yunnan, China, formed in a back-arc, extensional setting (Sillitoe and Lehmann 2022). This suggests that Sn-Cu deposits can form in a variety of tectonic settings with a common presence of coeval magmatism. The latter provides the heat and fluid source and plays an important role in mobilizing metals and transporting metals to form Sn-Cu deposits with multiple heating/hydrothermal episodes recorded in most locations (Romer and Kroner 2016). Due to the complex nature of mineralizing processes resulting from multiple magmatic fluid pulses, most Sn-Cu deposits consist of complex mineral parageneses with likely several generations of mineralization, where Sn, Cu, Zn, and other metals are deposited in multiple stages. This understanding of Sn-Cu deposits allows us to systematically propose a Cu deposition model for the HMF.

The stages of ore paragenesis indicate decoupled Sn and Cu mineralization at the Baal Gammon and Isabel deposits. The Sn episode in the HMF is the earliest record of mineralization in the Kennedy Igneous Association and occurs in fractionated granites of the O’Briens Creek Supersuite (Figs. 1b, 11a and 12a). This episode lasted from 340 to 310 Ma, where the granites were emplaced in a compressional tectonic regime (Cheng et al. 2018; Kumar et al. 2022). Deposition of Sn is facilitated by reduced magmas (e.g., O’Briens Creek Supersuite, Fig. 11a) where Sn species were transported in acidic and reduced hydrothermal fluids (Heinrich 1990). Copper mineralization, in contrast, is promoted by oxidized magmas (Murakami and Ishihara 2013; Sillitoe and Lehmann 2022), which further supports the decoupled deposition of Sn and Cu in the HMF. For these metal pairs to occur together, the reduced magma pool requires fertilization by mafic melts. Zircon geochronology by Kumar et al. (2022) showed that emplacement of granodiorite and other mafic-endmember granitic suites (Alamaden Supersuite; Fig. 1) occurred between 320 and 300 Ma in the HMF. The emplacement of these granodiorite units in the HMF (Fig. 12a), in combination with Nd and Hf isotope anomalies observed in the regional granites, suggests a magmatic recharge between 320 and 300 Ma (Champion and Bultitude 2013; Cheng et al. 2018). This magmatic recharge was likely caused by magma underplating that fertilized the reduced magma pool (Fig. 12a). Magma underplated fertilization of a reduced magma pool has been observed in other Sn-Cu deposits, such as the San Rafael deposit in Peru (Harlaux et al. 2021b), deposits in the Andean Sn belt (Sillitoe and Lehmann 2022), and the Tongshanling deposit in Nanling Range, China (Wu et al. 2021). Following this, an episode of tectonic extension and crustal thinning between 300 and 280 Ma supplied the heat that generated the Slaughter Yard Creek Volcanics in the HMF (Fig. 12a; Kumar et al. 2022). The Slaughter Yard Creek Volcanics were generated from a fertile protolith that formed oxidizing magma and deposited sulfide and critical metals in the HMF.

Fig. 12
figure 12

a Tectonic synthesis of the HMF modified from Kumar et al. (2022). b Simplified genetic model of Cu mineralization in the HMF, showing intrusion of porphyry that caused sulfide mineralization. The metal pathway is indicated with arrows for the hydrothermal system that was likely active. c Schematic illustration of different mineralization styles in fractionated granites adopted from Baker et al. (2005), depicts metal (Cu, Sn and W) mineralization linked to depth

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The intrusion of the Slaughter Yard Creek Volcanics allowed access to the necessary ingredients, such as magmatic sulfur, chlorine, and heat to initiate the hydrothermal cycling of Cu, Zn, In, Ag, Fe, and Pb chlorine ligand complexes that were deposited as magmatic-hydrothermal polymetallic deposits in the HMF (Fig. 12b). The deposition of chalcopyrite in acidic condition caused alteration of the wall rock that is evidenced by silicification at the Baal Gammon and Isabel deposits (Kumar et al. 2022, 2023). Similar alteration and mineralization patterns were observed in the Cornwall ore field (Müller et al. 2006) and the Bolivian Sn-Cu field (Fig. 12c; Torró et al. 2019b, Torró et al. 2019a; Torres et al. 2019). In these styles of Sn deposits that are mostly related to fractionated granites, Cu mineralization is strongly linked to depths above ~ 5 km (Fig. 12c; Baker et al. 2005).

Conclusion

The Baal Gammon and Isabel deposits show that Sn and Cu mineralization in the HMF were derived from different sources. Field evidence and petrology of ore samples indicate that Sn is early, whereas Cu is a late phase of sulfide mineralization. The sulfide mineralization is closely associated with the emplacement of the Slaughter Yard Creek Volcanics, whereas Sn is present in the UNA Porphyry in the Baal Gammon deposit. Constrained low sulfur isotope values indicate that the mineralizing fluids were magmatically derived with little interaction with connate water. The sulfur isotopes and textures of chalcopyrite and pyrrhotite from the UNA Porphyry and the Slaughter Yard Creek Volcanics indicate possible interaction of sulfidic fluid with Sn mineralization. The temperature of the sulfide mineralization was estimated from the sulfur isotopes of the Isabel deposit at 300 °C. The chalcopyrite shows high concentrations of trace element compared to pyrrhotite in the Baal Gammon deposit. The intrusive porphyries of the Slaughter Yard Creek Volcanics played a role in the evolution of the hydrothermal-magmatic event that formed Cu mineralization in the HMF.