Abstract
Mineralized shear zones in the Archean Missanabie-Renabie gold district (~ 1.1 Moz Au; Wawa, Ontario, Canada) locally define composite orebodies that record three hydrothermal events: (1) a pre-orogenic Au1 event (pre-D1 and pre-prograde-metamorphic); (2) a syn-orogenic, post-peak-metamorphic Retrograde event (syn-D3); and (3) a late syn-to post-orogenic Au2 event (late syn- to post-D4). Genetic considerations indicate the orebodies are hybrids with early intrusion-related (Au1) and later orogenic (Retrograde + Au2) events. Pearson product-moment correlation coefficients (log10) of whole-rock and LA-ICP-MS pyrite trace metal datasets distinguish Au1 from Au2 mineralization by Au-Ag, Au-Bi, and Au-Te correlations > 0.7 (p < 0.05) in the former, irrespective of sample medium and analytical method. An Au-Mo correlation in whole rock data (0.58–0.76; p < 0.05) further distinguishes Au1 from Au2 and supports an independently inferred intrusion-related origin for Au1. Sulfur isotope data is similar for both Au1 and Au2 pyrite with average δ34S values of − 5.5‰ ± 0.2‰ (1σ) and − 3.5‰ ± 0.3‰ (1σ) and average Δ33S values of 0.4‰ ± 0.1‰ (1σ) and − 0.3‰ ± 0.2‰ (1σ), respectively. SIMS δ18Oquartz values for the Au1, Retrograde, and Au2 events largely overlap and, like δ18Ocarbonate values of previous studies, tend to be lower than values typical of Archean gold deposits. The results of this study suggest that correlation coefficients in trace metal datasets are useful in discriminating and characterizing different gold events. Caution is emphasized with the use of S- and O-isotope datasets for these purposes. The presence of low δ18O values in vein quartz and carbonate is best explained by an 18O-depleted fluid formed during the Retrograde hydrothermal event. The latter is inferred at 2580 ± 21 Ma based on U-Pb geochronlogy of hydrothermal titanite, and relates to deformation and metamorphism in the nearby, amphibolite- to granulite-grade Kapuskasing metamorphic belt. Geochronological and geochemical evidence suggest that the 18O-depleted fluid may have formed via the devolatilization of biotite-bearing granitoids during deep-crustal metamorphism.
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Introduction
Geochemical studies of gold deposits in metamorphic terranes aim to constrain the nature of mineralizing fluids, metal reservoirs, and ore-forming processes (Goldfarb and Groves 2015). However, using geochemical data to achieve this is not always straightforward. For example, mass-dependent sulfur isotope fractionation is influenced by intensive fluid properties (T, fO2, and pH; Ohmoto and Rye 1979; Seal 2006), which can produce δ34S values in ore-stage sulfides that are non-characteristic of their source. A recent example of this is Schlegel et al.’s (2017) study of the Prominent Hill IOCG deposit where δ34Spyrite varies from − 34 to + 30‰ despite the original fluid having an inferred δ34S of about 4‰. This effect in part can account for some of the significant inter- (e.g., McCuaig and Kerrich 1998) and intra- (e.g., Godefroy-Rodríguez et al. 2020) deposit variance in δ34S sulfide values not uncommon in orogenic gold deposits. Thus, geochemical tracers sensitive to changes in the chemical and physical properties of a fluid are limited in uniquely defining a link between a fluid’s source and the site of ore formation. Trace metal associations are another parameter commonly considered, but are often characterized by a similar Au-Ag ± As ± B ± Bi ± Sb ± Te ± W association for geologically dissimilar syenite-associated, orogenic, and intrusion-related gold deposits (Robert 2001; Groves et al. 2003; Goldfarb et al. 2005; Hart 2007). Such apparently similar metal associations in different deposit types largely obscure their genetic significance. Laser ablation inductively coupled plasma-mass spectrometry (LA-ICP-MS) elemental mapping of sulfides has also introduced cautionary aspects, such as the realization of a protracted element paragenesis and likely remobilization and decoupling of Au from other elements (Neyedley et al. 2017; Gourcerol et al. 2018a, b; Kerr et al. 2018), which appears to be the norm rather than exception in orogenic gold systems (Kontak et al. 2018).
Vein quartz and carbonate δ18O values are considered to be relatively consistent among orogenic gold deposits (e.g., Kerrich 1989; McCuaig and Kerrich 1998; Goldfarb and Groves 2015). However, compilations suggest fluid types and/or vein-forming conditions (e.g., temperature, fluid mixing) are in fact variable, as it has been shown that significant δ18Oquartz variation occurs in both gold districts (e.g., ~ 5‰ in Val-d’Or; Beaudoin and Pitre 2005) and within the orogenic deposit class in general (e.g., ~ 20‰; Beaudoin 2011). Furthermore, the calculated δ18OH2O values generally overlap fields for magmatic and metamorphic waters (Kerrich 1989; McCuaig and Kerrich 1998), which causes uncertainty in fluid source identification. Overprinting hydrothermal events may present further difficulties, as the traditional use of bulk isotopic analysis values may reflect multiple fluid types.
The above discussion highlights how geochemical data for gold deposits in metamorphic terranes can be equivocal (Groves et al. 2003; Goldfarb et al. 2005; Goldfarb and Groves 2015). Here, we provide further insight into this phenomenon by using multiple geochemical methods to characterize three geologically well-constrained hydrothermal events in the ca. 1.1 Moz Au Missanabie-Renabie gold district (Wawa, Ontario, Canada; Fig. 1). These hydrothermal events are well-documented by recent studies (McDivitt et al. 2017, 2018): (1) pre-regional metamorphic, pre-D1, gold-bearing laminated quartz veins with phyllic alteration halos that account for the majority of gold endowment (the Au1 event); (2) post-peak metamorphic, syn-D3, gold-barren quartz ± epidote ± chlorite ± K-feldspar ± hematite veins related to extensive zones of hematite-bearing alteration (the Retrograde event); and (3) late-stage, post-peak metamorphic, syn- to post-D4 gold-bearing pyrite veins (the Au2 event). Whereas the study of McDivitt et al. (2017) constrains the different hydrothermal events within the structural framework of the Missanabie-Renabie gold district, the study of McDivitt et al. (2018) assesses the hydrothermal events from the perspective of fluid chemistry, alteration characteristics, and metamorphic timing. Results of these earlier studies illustrate that the Au1 event records the ingress of a H2O-CO2, Au-Ag-Bi-Cu-Mo-Pb-Te-W-Zn-S and large ion lithophile element bearing fluid into early (i.e., pre-D1), brittle, pre-metamorphic structures related to potassium-poor granitoid magmatism. In contrast, the Retrograde event reflects a post-peak metamorphic, Na-S-H2O fluid event that reactivated earlier brittle structures during later D3 deformation. Lastly, the Au2 event is represented by gold-bearing pyrite veins that overprint phases of the Retrograde event during or after a D4 reactivation of D3 structures. Because the different hydrothermal events contrast in their metal signatures and fluid chemistries, as well as their structural and metamorphic timing, they are interpreted to record different genetic processes. For example, because the Au1 event is localized to pre-metamorphic and pre-orogenic brittle structures related to potassium-poor granitoids, an intrusion-related origin is inferred (McDivitt et al. 2018). In contrast, D3 and D4 are both post-peak metamorphism and late-stage brittle-ductile deformation events thereby favoring an orogenic affiliation for the Retrograde and Au2 events (McDivitt et al. 2018) Here, we geochemically characterize these three different and contrasting hydrothermal systems using: (1) δ34S and Δ33S values of pyrite; (2) trace-metal signatures using whole-rock analysis and LA-ICP-MS elemental mapping of pyrite; and (3) in situ secondary ion mass spectrometry (SIMS) for δ18Oquartz. We also report in situ LA ICP-MS U-Pb geochronology results for titanite related to the Retrograde event to constrain its age.
A corollary but related intent of this study is to use our δ18O data to further assess the similarities among deposits of the Wawa gold camp highlighted in McDivitt et al. (2017, 2018). Particularly relevant based on earlier studies (Callan 1991; Samson et al. 1997; Haroldson 2014) is that vein δ18Ocarbonate values across the camp tend to be low when compared to gold deposits of the Abitibi greenstone belt (Kerrich 1990a). To provide insight into these δ18O values, we evaluate the results in the context of our in situ SIMS quartz analysis and fluid inclusion δD data from quartz to test potential origins of this 18O-depleted hydrothermal fluid.
Geological setting and gold mineralization
From 1941 to 1991, ~ 1.1 million ounces of gold were produced from shear zones containing laminated quartz veins in the Renabie mine area (Fig. 2a, b; Turek et al. 1996; Callan and Spooner 1998). This historic production was sourced from the Renabie underground mine and the C-zone, Nudulama, and Braminco open pits (Fig. 2b), which are hosted in 2720.8 ± 1.4 Ma biotite tonalite of the Missinaibi Lake batholith (MLB; Kamo 2015; Robichaud et al. 2015). The batholith is a plutonic component of the Wawa gneiss domain, which is a predominantly amphibolite-facies region of felsic gneisses, granitoid plutons, and minor supracrustal rocks. The granitoid plutons and supracrustal rocks range in age from ca. 2.92 to 2.66 Ga, with deformational/metamorphic events having occurred between ca. 2.70 to 2.64 Ga (Moser 1989, 1993, 1994; Percival and West 1994). The Wawa gneiss domain terminates to the east at the subprovince-bounding Kapuskasing structural zone, which is a linear, amphibolite- to granulite-grade metamorphic belt exposed in a Neoarchean to Paleoproterozoic intracratonic uplift (Percival and West 1994). The Kapuskasing structural zone records a protracted deformational and metamorphic history from ca. 2.70 to 2.50 Ga, with metamorphic U-Pb zircon ages representing granulitization from ca. 2.60 to 2.58 Ga, and U-Pb titanite ages recording progressive cooling from ca. 2.60 to 2.50 Ga (Moser 1989, 1993, 1994; Percival and West 1994). The transition from the Kapuskasing structural zone across the Wawa gneiss domain into the Michipicoten greenstone belt (MGB; Fig. 1) has been considered an oblique crustal cross section (Percival and Card 1985), with the Kapuskasing structural zone representing the deep-crustal equivalent of the MGB.
The MGB is dominated by greenschist-facies metavolcanic rocks formed during three principal bimodal volcanic cycles (ca. 2.90, 2.75, and 2.70 Ga; Sage 1986; Turek et al. 1982, 1984; Sage and Heather 1991; Turek et al. 1992; Heather and Arias 1992). Granitoid plutons of various ages intruded the MGB: (1) the 2745 ± 3 Ma dioritic- to granodioritic Jubilee stock (Sullivan et al. 1985); (2) the 2720.8 ± 1.4 Ma biotite tonalite of the MLB; and (3) the compositionally variable granitoids of a ca. 2.68 to 2.66 Ga intrusive suite (Turek et al. 1996; Robichaud et al. 2015).
Other granitoid-hosted deposits of the Wawa gold camp are in plutons which intrude the MGB. These include the Surluga, Jubilee, and Minto deposits hosted in the Jubilee dioritic stock (Samson et al. 1997), and the Magino deposit hosted by the undated Webb Lake trondhjemitic stock (Fig. 1; Haroldson 2014). Features of these mineralized sites are considered to be similar to those in the Missanabie-Renabie district (McDivitt et al. 2017, 2018): (1) gold-bearing, pre-orogenic quartz which host low- to moderate-salinity H2O-CO2 fluid inclusions and their unmixed counterparts; (2) these are overprinted by syn-orogenic, gold-barren breccia veins and hematite-bearing alteration assemblages characterized by CO2-poor aqueous fluids of variable salinities; and (3) fracture-hosted, auriferous pyrite cross-cuts the earlier hydrothermal events in the Missanabie-Renabie gold district and at Magino. In the Missanabie-Renabie gold district, the pre-orogenic, syn-orogenic, and late-stage, pyrite-associated mineralization styles correspond to the Au1, Retrograde, and Au2 events, respectively. Correlative hydrothermal events at the Jubilee stock and Magino deposits are considered to represent regional equivalents.
Analytical methods
Polished thin sections and epoxy resin mounts containing quartz and/or pyrite from the different vein types were examined in reflected and transmitted light. These samples were further characterized on a JEOL 6400 scanning electron microscope (SEM) coupled to an Oxford Sight energy dispersive X-ray spectrometer (EDS) at the Micro-Analytical Centre (MAC), Laurentian University. Analytical conditions employed an accelerating voltage of 20 kV, a 1.005 nA beam current, a working distance of 15 mm, and an acquisition time of 5 s. Cathodoluminescence (CL) imaging on a single quartz vein sample from the Retrograde event was completed using a Gatan ChromaCL™ detector attached to it with a linear array of photomultiplier tubes with 16 separate photocathodes.
Representative gold mineralized samples were analyzed for trace metals at the Geoscience Laboratories of the Ontario Geological Survey, Sudbury, Ontario. Sample location information and analytical data is provided in McDivitt (2016). Samples comprised ~ 100 cm3 of rock material and consisted of Au1 laminated veins and phyllic alteration (analyzed separately) and Au2 pyrite veins. Both vein and alteration material was separated from adjoining wall rock. Samples were pulverized and dissolved in an open-vessel agua regia digestion and concentrations were determined by ICP-MS (IML-100 technique; McDivitt 2016). For the purpose of plotting, element concentrations reported below analytical are shown as one-half the detection limit value, which are as follows (in ppm): Au (0.002), Ag (0.01), As (0.8), Bi (0.01), Cu (0.7), Mo (0.03), Pb (0.2), Sb (0.06) Te (0.01), U (0.01), W (0.02), and Zn (3).
One representative pyrite grain from each of the Au1 and Au2 events was analyzed using LA-ICP-MS in traverse mode to generate element maps and assess gold-trace metal associations. Elements analyzed were Ag, As, Au, Bi, Cd, Co, Fe, In, Mo, Ni, Pb, Sb, Sc, Te, U, W, and Zn. This work was done at the Geochemical Fingerprinting Lab of Laurentian University following the details outlined in the Electronic Supplementary Materials (ESM Table 1) and elsewhere (Neyedley et al. 2017; Gourcerol 2018a, b; Kerr et al. 2018). In addition to generating element maps, the LA-ICP-MS traverse data was broken down into time-slice domains or TSD which allow for data representing element concentrations over specific time intervals (i.e., 0.23 s for one complete analysis) to be displayed graphically on scatter plots (Gourcerol et al. 2018a, b). As discussed elsewhere (Gourcerol et al. 2018a, b), the error associated with the TSD approach is estimated to be < 20%. Regions of stepped data occur at low elemental concentrations when the TSD data is plotted on scatter plots (e.g., Gourcerol et al. 2018a, b). These regions of stepped data define domains in which the data is close to or below the lower limits of detection. The stepped nature of low-concentration data is due to the pulse counting detector and low backgrounds combined with variable intensities from the pyrite. The TSDs that returned negative concentration values were assigned a value of 0.1 ppm for the purpose of plotting and excluded from the statistical analysis described below.
Pyrite samples were powdered using a Dremel™ drill and analyzed at the Manitoba Isotope Research Facility, University of Manitoba, Winnipeg, Canada. Samples were analyzed by a Costech™ 4010 Element Analyzer coupled to a Thermo Finnigan™ Delta V Plus isotope-ratio mass-spectrometer via an open-split interface. Each sample was analyzed three times non-consecutively as a test of reproducibility. Isotopic ratios are reported as δ34S and δ33S (‰) relative Vienna Cañon Diablo Troilite (VCDT). The international standards IAEA-S1, IAEA-S2, and IAEA-S3 were used as calibration standards and analyzed at the beginning, middle, and end of each run. Least-squares linear regression of the known and measured standard values was used to produce a calibration line. Anderson pyrrhotite was used as an internal standard and repeated analysis of this standard yielded uncertainties (1σ) of 0.2‰ for δ34S and 0.3‰ for δ33S. The uncertainty for Δ33S is estimated to be 0.1‰ by the standard error of the mean Δ33S value calculated for the internal standard.
Vein quartz was crushed into mm-size fragments and hand sorted under a binocular microscope to avoid impurities. Select grains were mounted in epoxy resin with ~25 grains per mount. Spot locations selected for SIMS O-isotope analysis were chosen based on areas that lacked fractures and mineral inclusions. In addition, a euhedral crystal of quartz from the Retrograde event with primary fluid inclusions was cut perpendicular to its C-crystallographic axis and investigated along a detailed SIMS O-isotope traverse following its CL imaging. The SIMS δ18O analysis was completed at the University of Manitoba Isotope Research Facility using a CAMECA 7f ion microprobe with a 3 nA cesium (Cs+) primary beam accelerated to 10 kV and focused to a 20 μm spot size. An offset voltage of 300 V was used to eliminate molecular ion interferences. Ions were detected with a Balzers SEV 1217 electron multiplier coupled with an ion-counting system using a dead time of 52 ns. Both 18O and 16O were detected by switching the magnetic field. The results are presented using the δ-notation (‰) relative to Vienna Standard Mean Ocean Water (VSMOW). Internal standards used for calibration were analyzed at multiple times during the analysis; the spot reproducibility on the standards ranged from 0.4 to 0.5‰. The 1σ error value associated with unknown measurements is estimated to be 1.2‰.
Two samples of quartz from the Retrograde event at Nudulama East were analyzed at the Facility for Isotope Research at Queen’s University for H isotopic composition. Fluid inclusions from the same samples were studied by McDivitt et al. (2018). Fluid inclusions within the quartz comprise two-phase (liquid-vapor) H2O inclusions with ~ 10–20 vol% vapor and occur as primary, pseudosecondary, secondary, and indeterminate fluid inclusion assemblages. Samples were weighed into silver capsules, degassed at 100 °C for 1 h, and then loaded into a zero-blank auto sampler. The hydrogen isotopic composition was measured using a Thermo-Finnigan thermo-combustion element analyzer coupled to a Thermo-Finnigan Deltaplus XP Continuous-Flow Isotope-Ratio Mass Spectrometer. Values are reported relative to VSMOW with a precision of 3‰.
Measurements for in situ U-Pb geochronology of titanite (n = 86) within hematite-bearing Retrograde event alteration (Sample NUE-28C) was conducted using LA-ICP-MS. Prior to analysis, the grains were imaged in BSE mode on the SEM. The LA-ICP-MS analysis was completed at the Geochemical Fingerprinting Lab of Laurentian University using the operating conditions and data processing procedures given in ESM Table 2. Titanite OLT-1 (1014.8 ± 2 Ma; Kennedy et al. 2010) was used as a calibration standard and titanites BLR-1 (1047.1 ± 0.4 Ma; Aleinikoff et al. 2007) and FCT (28.53 ± 0.05 Ma; Schmitz and Bowring 2001) were used as secondary standards. The secondary standards were analyzed 6 time each throughout the run and returned ages of 1033 ± 18 Ma (BLR-1) and 29.6 ± 1.2 Ma (FCT).
Vein-hosted quartz, pyrite, and alteration-related titanite: defining features and textural characteristics
The analyzed materials were collected from mineralized outcrops in the MGB (Pileggi No. 1; Fig. 2a) and the MLB (Baltic D, C-zone, Nudulama, and Nudulama East; Fig. 2b). The different quartz and pyrite generations are displayed in the context of hydrothermal events and the camp’s deformational framework in Fig. 3 (McDivitt et al. 2017, 2018).
Au1 laminated quartz veins
Quartz (Qtz1) crystals range from equidimensional, coarser crystals (~ 1–2 mm) with well-defined triple junction boundaries to smaller (~ 0.1–0.2 mm), irregular, amoeboid-shaped crystals with lobate to cuspate boundaries. The smaller crystals occur in association with sericite domains or lithons, which give Au1 veins their characteristic laminated texture (Fig. 4a, b). Pyrite (Py1) within these veins is fine- to medium-grained, euhedral to subhedral, and commonly hematite altered on its margins (Fig. 4c). Magnetite occurs in textural equilibrium with the late-stage hematite. Quartz is the dominant mineral in Au1 veins (~ 80%), with the lesser sericite, carbonate, and accessory titanite, allanite, epidote, tellurides, and sulfides (pyrite, chalcopyrite, molybdenite, galena, sphalerite). These veins are surrounded by phyllic alteration halos (quartz (~ 30%)-sericite (~ 65%)-pyrite (~ 2–5%)) containing the same accessory mineralogy as the laminated veins.
Retrograde hematite-associated alteration veins
The Retrograde event veins are associated with hematite-bearing alteration assemblages and comprise both discrete-planar and breccia veins that overprint and include phases of the Au1 event (Fig. 4d, e; McDivitt et al. 2017). Vein fill consists of quartz (Qtz2), chlorite, epidote, K-feldspar, and earthy-red hematite (Fig. 4f). These phases are commonly euhedral and locally define comb-like textures, with quartz and epidote generally medium- to coarse-grained whereas chlorite is a late-stage, fine-grained infill. The modal proportions of infill minerals vary, with hematite and K-feldspar the least abundant (typically < 20%), and quartz, chlorite, and epidote the most abundant (locally to 100%). These veins are associated with zoned alteration halos comprising proximal K-feldspar-hematite alteration and distal chlorite-titanite-albite-hematite alteration where titanite forms fine-grained aggregates in chloritized biotite of the MLB (Fig. 5a).
Au2 massive pyrite veins and quartz tension gashes
The Au2 veins consist of fine- to medium-grained, anhedral pyrite (Py2), which commonly displays jagged- to serrated-grain boundaries (Fig. 4g). These veins lack alteration halos, but pyrite is commonly disseminated in the adjacent wall rocks. Quartz (Qtz3) is the majority infill phase (90–95%) of NW-striking quartz-tension gashes at the Pileggi No. 1 outcrop (Fig. 4h), with lesser, brown, granular oxide and fine-grained pyrite accounting for the remainder. Quartz crystals are typically elongate in thin section (~ 1–5 mm), display undulatory extinction, and have irregular- to jagged-grain boundaries. Micro-fracturing is abundant in some of the quartz, in particular where zones of cataclasite are present. The Qtz3 tension gashes do not carry significant gold grade (avg. 0.017 ppm Au; n = 10), but they are similar in orientation and structural timing to the Au2 pyrite veins (McDivitt et al. 2017). Therefore, for the purpose of this study Qtz3 is assigned to the Au2 hydrothermal event (Fig. 3).
Analytical results
Titanite U-Pb geochronology
The results of 86 U-Pb titanite spot analyses are shown on a Tera-Wasserburg concordia diagram in Fig. 5b and given in the ESM (Table 3). Calculations and plots were completed in Isoplot v. 3.76 (Ludwig 2003). The overall dataset displays a high degree of scatter, and discordance due to both Pb loss and common Pb incorporation (Storey et al. 2006; Kirkland et al. 2017). The dataset was filtered on the basis of percentage discordance (< 0% and < 20% excluded) and uranium concentration (< 9 ppm U excluded). Two points were excluded on the basis of large 2σ error values. Five additional points were excluded on the basis of spot overlap with adjacent mineral phases. An unanchored regression of the reduced dataset (n = 38) yields an age of 2577 ± 51 Ma (95% confidence; MSWD = 1.8) with a 207Pb/206Pb intercept of 0.96 ± 0.25 (95% confidence). The 207Pb/206Pb intercept is in agreement with a common 207Pb/206Pb value of 1.0795 inferred from host rock age (2720.8 ± 1.4 Ma biotite tonalite) and the terrestrial Pb evolution model of Stacey and Kramers (1975). The latter is used in an anchored regression (Fig. 5c) and in a 207Pb-corrected 238U/206Pb weighted mean calculation (Fig. 5d), which return respective ages of (95% confidence) of 2586 ± 21 Ma (MSWD = 1.7) and 2580 ± 21 Ma (MSWD = 1.8).
Whole rock and LA-ICP-MS trace metal associations
Samples analyzed for trace metals are plotted in Fig. 6. In general, the data indicate (1) significant range in Au from 0.004 to 27.05 ppm for Au1 samples and 0.04–3.99 ppm for Au2 samples (Fig. 6a); (2) similar range in Ag from 0.01–11.83 ppm for Au1 samples and 0.53–4.34 ppm for Au2 samples (Fig. 6a); (3) Au1 and Au2 samples are typically enriched in Bi (up to ~ 27 ppm; Fig. 6b) relative to Au1 veins from Pileggi No. 1 (max = 0.07 ppm); (4) a large range in Te and Pb ( ~0.01/0.2 to ~ 40 ppm; Fig. 6c, d) and a restricted range for U (0.01 to 2.53 ppm; Fig. 6e), with Au1 veins from Pileggi No. 1 relatively depleted in Te (max = 0.33 ppm; Fig. 6c), Pb (max = 0.9 ppm; Fig. 6d), and U (max = 0.05 ppm; Fig. 6e); (5) select samples of the Au1 event are markedly enriched in Mo (to > 350 ppm upper detection limit; Fig. 6f); and (6) largely overlapping values among different sample populations for Cu, W, and Zn, but with W generally low (most < 1 ppm), and most Zn and Cu values range from 1 to 100 ppm (Fig. 6g, h). The Pileggi No. 1 samples are enriched in Cu (to 233 ppm; Fig. 6h) relative to most other samples. In addition, Sb (not plotted) is below detection (0.06 ppm) in all samples and As (not plotted) is either below detection (0.8 ppm) or present only slightly above detection (max = 4.5 ppm). It is therefore apparent that for these samples the Au-Sb-As association, so common in other gold systems, is lacking.
Pearson product-moment correlation coefficients and associated p values were calculated for different sample populations and illustrate the following (Table 1): (1) well-defined (0.58–0.99), significant (p < 0.05) correlations between Au and Ag, Bi, Pb, Te, and Mo in the Au1 vein and phyllic alteration samples hosted in the MLB; (2) lower, non-significant (p > 0.05) correlation coefficients (0.04–0.54) between Au and Ag, Bi, Mo, Te, and Pb in Au2 samples; and (3) a strong Au-Ag (0.97; p < 0.05) in Au1 veins from the Pileggi No. 1 outcrop, but correlation coefficients between Au and Te, Pb, Mo, and Bi are lower (0.07–0.81) and nonsignificant (p > 0.05).
Element maps and corresponding trivariate scatter plots of TSD data for representative samples of Py1 and Py2 are shown in Figs. 7 and 8, respectively. The element map for Py1 (Fig. 7a) shows a central area relatively enriched in Ni (Fig. 7b) with a rim enriched in Co (Fig. 7c) but with obvious primary growth zones lacking. The remaining elements shown all lack spatial correlation with Ni and Co and instead define local or spotty enrichment zones (Fig. 7d–j). It is likely that at least some of the Au, Ag, Bi, Te, and Pb enrichment records the presence of micro-inclusions of Au-bearing tellurides with associated Ag, Bi, and Pb, as shown by SEM-EDS analysis (Fig. 7k, l). In addition to Au, other elements (Ag, Bi, Te, Pb, U, W) also define narrow, linear zones of enrichment (Fig. 7d–j). Similar relationships are evident in the trivariate scatter plots of Fig. 7, which show positive covariation trends between Au and Ag (Fig. 7m–o), Bi (Fig. 7p), and Te (Fig. 7q); in contrast, Pb displays a weak but positive covariation with Au (Fig. 7r) and neither U nor W show an association with Au (Fig. 7s, t).
The element map for Py2 (Fig. 8a) contrasts with Py1 in that there is a subtle zoning defined by Co and Ni (Fig. 8b, c). In addition, there is noted enrichment in the central part of the pyrite crystal by Au (Fig. 8d) and other elements (Ag, Bi, Te, Pb, U, and W; also Zn but not shown; Fig. 8e–j). The trivariate plots also document Au-Ag (Fig. 8k–m) and Au-Bi-Te (Fig. 8n, o) associations in the form of strong and moderate- to weak-, positive covariations, respectively. In contrast, Pb, U, and W lack covariation with Au (Fig. 8p–r).
Pearson product-moment correlation coefficients were calculated from the TSD data using log10 values in the same manner as for whole-rock data (Table 1); a graphical comparison of the results is shown in Fig. 9. Significant (p < 0.05) correlation coefficients include values > 0.7 defined by Au-Ag, Au-Bi, and Au-Te in Py1, with the Au-Pb correlation lower (0.37; p < 0.05) and no Au-Mo correlation (− 0.03; p < 0.05). In Py2 there is a significant Au-Ag correlation (0.66; p < 0.05), and Au-Bi, Au-Pb, and Au-Te correlations are lower (0.36–0.45; p < 0.05), with no Au-Mo correlation (0.08; p < 0.05).
Stable isotopes
Sulfur isotope analysis (Table 2) for Py1 and Py2 returned similarly negative δ34S average values (Py1 = − 5.5‰ ± 0.2‰ (1σ) n = 3; Py2 = − 3.5‰ ± 0.3‰ (1σ) n = 12) and δ33S average values (Py1 = − 3.2‰ ± 0.2‰ (1σ) n = 3; Py2 = − 2.1‰ ± 0.3‰ (1σ) n = 12). The δ34S values are consistent with Callan and Spooner (1998) for pyrite from the Renabie mine (δ34SAvg = − 5.7‰ ± 0.9‰ (1σ) n = 8). Average Δ33S values are similar for both Py1 (− 0.4‰ ± 0.1‰ (1σ) n = 3) and Py2 (− 0.3‰ ± 0.2‰ (1σ) n = 12).
The results of SIMS δ18Oquartz analysis for Qtz1, Qtz2, and Qtz3 are summarized in Fig. 10 (a, b, and c, respectively; ESM Table 4). Whereas the δ18O values of Qtz1 range from 3.9 to 13.4‰ and approximate a standard normal distribution (avg. = 8.4‰; Fig. 10a), the data for Qtz2 (avg. = 6.7‰; Fig. 10b) define a negatively skewed distribution with a larger proportion of values lower than 6‰, with minimum and maximum values of 1.2‰ and 9.7‰, respectively. The δ18O values in Qtz3 are similar in value and distribution to Qtz1 and range from 5.8‰ to 10.9‰ (avg. = 8.0‰; Fig. 10c). For the SIMS transect of a Qtz2 crystal (Fig. 10e, f), the data are summarized in both a histogram (Fig. 10d) and superimposed on the grain image (Fig. 10f). Results for the transect indicate a range for δ18O values of 5.4‰ to 12.6‰ (Fig. 10d) with a general trend of lower values in the core and higher values towards the rims (Fig. 10e, f). There is no apparent relationship between the areas of markedly different CL response and δ18O values. Hydrogen isotope analysis from the two samples of Qtz2 (Table 2; Fig. 10e)) yielded identical δD values of − 37‰.
Discussion
A geochemical comparison of the hydrothermal events
When the trace-metal characteristics of Au1 samples from the MGB are compared to those from the MLB, geochemical contrasts are evident that may represent local host-rock effects on fluid chemistry (Polito et al. 2001; Evans et al. 2006; Kontak et al. 2011; Gourcerol et al. 2018b). For example, the Au-Bi and Au-Te plots (Fig. 6b, c) show that laminated vein samples hosted in the MGB (Pileggi No. 1 samples) are lower in Bi and Te (below 35th percentile for Au1 samples), and define a trend discordant to the Au1 samples from the MLB. Similarly, the MGB laminated vein samples are enriched in Cu (above 70th percentile for Au1 samples; Fig. 6h). When comparing correlation values between laminated veins in the MLB and the MGB, the Pileggi No. 1 veins only show a significant (p < 0.05) Au-Ag correlation and lack the Au-Bi, Au-Pb, Au-Te, and Au-Mo correlations shown by the laminated veins in the MLB (Table 1). When correlation coefficients are used to compare whole-rock samples of Au1 and Au2 from the MLB (Fig. 9), Au1 samples show significant (p < 0.05) and consistent Au-Ag, Au-Bi, Au-Te, Au-Pb, and Au-Mo correlations that are not defined by Au2 samples. When LA-ICP-MS pyrite data is considered (Fig. 9), Au1 pyrite stands out with higher (> 0.7) Au-Ag, Au-Bi, and Au-Te correlations than Au2 pyrite. Overall, the correlation coefficient patterns defined by both whole-rock data and pyrite from Au2 are similar. The correlation coefficient pattern defined by Au1 pyrite shows similar Au-Ag, Au-Bi, and Au-Te correlations in comparison to the whole-rock data; however, the Au-Pb and Au-Mo correlations are markedly lower in the pyrite data (Fig. 9). This suggests that the Au-Pb and Au-Mo correlations in whole-rock data reflect the presence of other phases such as galena and molybdenite in the whole-rock samples. The correlation coefficient data shows clear differences in the trace-metal signatures of the Au1 and Au2 mineralization styles, and the results emphasize that correlation coefficients in trace metal datasets are useful ways to characterize and discriminate among different hydrothermal gold events. In the context of using these trace-metal associations to provide insight into ore deposit genesis, the Au-Ag-Bi-Te-Pb associations are not indicative of a certain deposit type, with the same associations displayed by syenite-associated, orogenic and intrusion-related deposits (Robert 2001; Groves et al. 2003; Goldfarb et al. 2005; Hart 2007). The Au-Mo correlation and elevated Mo concentrations in Au1 samples (Fig. 6f) are perhaps the most insightful as to deposit affinity, as these are not characteristic of orogenic gold deposits (Groves et al. 2003; Goldfarb et al. 2005; Goldfarb and Groves 2015), and these are features commonly documented in magmatic-hydrothermal systems, including intrusion-related deposits (Thompson et al. 1999), syenite-associated deposits (Robert 2001), gold deposits with anomalous metal associations (Groves et al. 2003), and porphyry deposits (Seedorff et al. 2005). Thus, it is permissible to infer that the Au1-Mo association and Mo enrichment in Au1 samples reflect an intrusion-related origin, as previously suggested based on both geological (McDivitt et al. 2017) and geochemical (McDivitt et al. 2018) evidence.
The sulfur isotopic signatures of Py1 and Py2 are similar (Table 2), although there is a 2‰ difference between the mean δ34Spyrite values of − 5.5‰ and − 3.5‰, respectively. Taking into account analytical precision (0.1‰ for Δ33S), the Δ33S values indicate components of mass-independently-fractionated sulfur in both Py1 and Py2 (Farquhar and Wing 2003; LaFlamme et al. 2018). The slightly-negative Δ33S values differ from the slightly-positive Δ33S common in gold deposits of the Yilgarn craton (Selvaraja et al. 2017; Godefroy-Rodriguez et al. 2020; LaFlamme et al. 2018; Groves et al. 2019), and in examples of Archean intrusion-related gold mineralization (Helt et al. 2014). The Δ33S values are more similar to those reported for Archean magmatic Ni sulfide and VMS mineralization (Xue et al. 2013; Sharman et al. 2015; Selvaraja et al. 2017; Ripley and Li 2017). The similar δ34S and Δ33S values for Au1 and Au2 pyrites suggests that isotopically similar reservoirs have been utilized throughout successive hydrothermal events in the district. Unlike the trace-metal data, the sulfur isotope data does not discriminate between the Au1 and Au2 events.
The δ18O values of Qtz1, Qtz2, and, Qtz3, which show considerable overlap, are also not a good discriminant of the different hydrothermal events. In addition, the δ18O values in all quartz generations tend to be lower than those typical of Archean lode gold deposits (Fig. 11a). The mean δ18O values of Qtz1, Qtz2, and, Qtz3 are similar to the mean δ18O values derived from equivalent carbonate generations from earlier studies (Fig. 11b). Whereas, the mean δ18O values from Au1 and Au2 quartz/carbonate are similar, the mean δ18O values from Retrograde quartz/carbonate are consistently lower.
Origin of low δ18O values in vein quartz and carbonate
Herein, we explore different models to account for the low δ18O values in vein quartz and carbonate throughout the camp. Retrograde breccia vein quartz from Nudulama East (Qtz2) is used to constrain the origin of the low δ18O values for the following reasons: (1) it returned the lowest mean δ18O values in both local and regional datasets (Fig. 11b); (2) δD analysis of primary inclusions (Fig. 10e), and the reproducible δD results (Table 2) agree with both petrographic observations and microthermometric data (McDivitt et al. 2018) indicating a single fluid is present in Qtz2; (3) the quartz is generally euhedral, coarse-grained, and characterized by primary growth textures suggesting it is likely to preserve a primary isotopic signature; and (4) U-Pb titanite geochronology infers the timing of the Retrograde event and quartz formation at 2580 ± 21 Ma, thus allowing it to be interpreted in the context of a regional temporal framework. We consider the temperature range of 200–400 °C based on the following (ESM Fig. 1): (1) a minimum T constrained from the Th values of aqueous fluid inclusions in Retrograde breccia vein quartz (McDivitt et al. 2018); and (2) a maximum T constrained by the upper limit of biotite chloritization reactions in granitic rock bodies (Yuguchi et al. 2015), as such reactions define alteration associated with the Retrograde quartz (McDivitt et al. 2018). Additional maximum to intermediary T constraints include the following: (1) the T range of retrograde silica solubility (Fournier 1985), as inferred by dissolution-reprecipitation textures in the cores of quartz crystals (Fig. 10f); (2) the T range for brittle-ductile deformation of quartz (Smith and Bruhn 1984; Sibson 2001), as brittle-ductile deformation of quartz occurred during the Retrograde event (McDivitt et al. 2017); and (3) similar temperatures of Archean orogenic quartz vein formation (Goldfarb et al. 2005).
Low δ18O values in vein quartz from porphyry-epithermal and epizonal Archean lode gold settings are interpreted to fingerprint the involvement of surficial fluids (Taylor 1979; Hagemann et al. 1994). In order to evaluate the presence of surficial fluids, we utilize the δ18O and corresponding δD values of different fluid types (Taylor 1974) in a fluid-mixing/cooling model. A fluid of δ18OH2O = 10‰, typical for Archean lode gold systems (McCuaig and Kerrich 1998), is modeled to mix with a high-latitude meteoric water (δ18OH2O = − 15‰) from 200 to 400 °C (Fig. 11c). A high-latitude meteoric water is used in the modeling since for any given volume it would produce a larger isotopic shift than a low-latitude meteoric water or seawater; thus, it is used as a conservative constraint. The δ18O values of Qtz2 crystals can be explained if the endmember fluid with δ18OH2O = − 15‰ is present in proportions ranging from ~ 25 to 50%. Corresponding δD values in the mixed system should be approx. − 43 to − 65‰ although this assumes the upper limit of δD reported in Taylor (1974) where δ18OH2O = 10‰ for a metamorphic fluid (δD approx. − 20‰). If the lower limit (δD approx. − 60‰) is considered, the δD values in the mixed fluid should range from approx. − 72 to − 85‰. The reproducible δD values returned from Qtz2 of − 37‰, which is typical of metamorphic and magmatic fluids, do not indicate the involvement of a high-latitude meteoric water. While some studies illustrate that δD values from quartz provide δDH2O values concordant with those calculated from hydrothermal muscovite (Ojala et al. 1995), others conclude that δD values from quartz represent both trapped fluid inclusions and structurally bound water and may not reflect the original δDH2O values (Simon 2001). In this latter case, δD quartz values may be lower than the inferred δDH2O values (Simon 2001). In consideration of this, the δD in Qtz2 of − 37‰ still do indicate the involvement of a high-latitude meteoric water, as the δD values in Qtz2 are too high. It is possible that the δD values in Qtz2 reflect a low-latitude meteoric water or seawater, but a low-latitude meteoric water (δ18O = − 5‰) would need to be present in proportions from ~40–90% to explain the Qtz2 δ18O values in the mixing/cooling model. Seawater would need to be present in proportions from ~ 60–100%. As discussed by Pickthorn et al. (1987), (Kerrich 1990b), and Kyser and Kerrich (1990), the transport of these large volumes of surficial fluid to crustal depths where orogenic gold mineralization forms in an isotopically unmodified state presents issues. This is not our favored interpretation for the low δ18O values in veins quartz and carbonate in the camp.
Fluid-rock interaction provides another means to modify δ18OH2O values (Kontak et al. 2011). However, the shear zone-hosted nature of the breccia at Nudulama East (McDivitt et al. 2017) favors a high fluid/rock ratio system typical of orogenic mineralization (McCuaig and Kerrich 1998). Furthermore, the formation of a fluid with low δ18OH2O values via fluid-rock interaction invokes low δ18O host rock values (< 5.5‰), such as those formed by the interaction of magmas with meteoric water, or by assimilation of altered country rock with low δ18O values (Bindeman and Valley 2001; Hammerli et al. 2018). This situation is not supported by the whole rock δ18O values of 6.4 to 9.5‰ for granitoids in the Wawa gneiss domain (Li et al. 1991).
Rayleigh distillation has been considered in explaining the isotopic data from Archean gold deposits (Kerrich 1989, 1990a; Samson et al. 1997; Kontak et al. 2016; Neyedley et al. 2017). When the δ18O Qtz2 range by is assessed by closed-system Rayleigh distillation using an initial δ18OH2O = 10‰ over a temperature range of 200–400 °C, the δ18Oquartz range requires significant fluid distillation with f values to approx. ≥ 0.2 where f is the amount of original fluid remaining. The SIMS δ18Oquartz values from the CL-imaged crystal show increasing values from core to rim (Fig. 10f), and contradict a model of progressive Rayleigh fluid distillation. In contrast, the isopleth-parallel distribution of the δ18Oquartz values over a 200 to 400 °C range (Fig. 11c) suggests a progressive cooling model associated with crystal growth. In the CL image, it is also evident that some of the lowest δ18O values (i.e., 5.4‰) occur juxtaposed against higher values (i.e., 9.4‰). In this case, analytical error is an important consideration (1σ = 1.2‰), but the juxtaposition of contrasting values in a growth zone of quartz with uniform CL indicates that any difference in δ18O may not reflect T since variation of T is commonly reflected in CL zoning (Alan and Yardley 2007; Rusk et al. 2008; Mao et al. 2017). It is possible that the local, small-scale δ18O variance in zones where CL is uniform records Rayleigh distillation effects, but overall the distribution of δ18Oquartz values is best explained by a cooling model with the initial fluid having a low δ18O value.
A low δ18O fluid as the product of a devolatilized Granitoid source?
An initial δ18OH2O value of ~ 4‰ is indicated for the Retrograde fluid if the average δ18O value from the core (Fig. 10f; 7.9‰) is considered to have formed at 400 °C. Geochronological constraints on the timing of the Retrograde fluid infer its circulation at 2580 ± 21 Ma, which post-dates local granitoid magmatism (ca. 2.72–2.66 Ga; Turek et al. 1996; Kamo 2015), but overlaps metamorphism and deformation in the amphibolite- to granulite-grade Kapuskasing structural zone (Fig. 1; Moser 1994; Percival and West 1994). Devolatilization of greenstone belt lithologies during regional metamorphism is often considered a fluid-generating mechanism responsible for veining and alteration in Archean gold districts (Groves et al. 1987; Phillips et al. 1987; Goldfarb and Groves 2015). The presence of Retrograde event phases in the MLB (Fig. 2), which is external to the MGB, renders this mechanism of fluid generation unlikely due to a lack of greenstone fluid sources at depth. If a metamorphic fluid source at depth is to be considered, then it is likely represented by granitoids and gneisses of the Wawa gneiss domain and Kapuskasing structural zone. Mineral δ18O values reported by Li et al. (1991) from granitoids of the Wawa gneiss domain (~ 9 to 11‰ for quartz; ~ 6 to 8‰ for feldspars; and ~ 2 to 5‰ for biotite) illustrate that devolatilization of such granitoids, in particular the biotite component, could generate an 18O-depleted metamorphic fluid, as δ18OBt-δ18OH2O fractionation at temperatures typical of amphibolite grade and higher is low (approx. − 2.5‰; Zheng 1993). Another consideration is the fractionation between aluminosilicate O and hydroxyl O, as the latter may be depleted in 18O relative to the former (δ18OBt-δ18OOH ~ 10–3‰ at 300–800 °C; Zheng 1993); hence, if the dehydroxylation of biotite is the dominant mechanism in metamorphic fluid production, the δ18OH2O of the fluid may be lower than the δ18O of the precursor biotite. In orogenic gold settings where crustal zones are characterized by high permeability, brittle-ductile deformation, and transitions from lithostatic to hydrostatic fluid pressure (Sibson 2001), open-system metamorphic devolatilization (Valley 1986) may preclude fluid-rock equilibrium and preserve the 18O-depleted nature of a fluid generated from the metamorphic devolatilization of biotite. Such a deep-crustal metamorphosed granitiod source potentially explains the lack of a significant carbonic component and gold mineralization associated with the Retrograde event in the Missanabie-Renabie district (McDivitt et al. 2018), as carbonic fluids and gold mineralization are often attributed to the devolatilization of carbonated greenstone belt lithologies (Powell et al. 1991; Elmer et al. 2006), or the involvement of mantle sources (Hronsky et al. 2012).The similarities of mean values in both our SIMS δ18O data and the compiled regional data (Fig. 11b) suggests the low δ18O values for the Retrograde fluid was a regional-scale phenomenon. Because the Retrograde fluid event overprints earlier Au1 mineralization, the low δ18O values for Au1 quartz may represent dissolution of Qtz1 and precipitation of Qtz2 in Au1 veins during the Retrograde event as is inferred by fluid inclusion evidence (McDivitt et al. 2018). In contrast the Au2 equivalent quartz and carbonate returns to a higher mean δ18O value (Fig. 11b), suggesting the event may record metamorphic fluids affected by a greater degree of country-rock equilibration (Fig. 11d).
Conclusions
A geochemical comparison of three hydrothermal events (Au1, Retrograde, and Au2) in the Missanabie-Renabie gold district using whole-rock and LA-ICP-MS pyrite trace metal analysis in addition to sulfur and oxygen isotope data highlights the following:
-
(1)
Pearson product-moment correlation coefficients (log10) of whole-rock and LA-ICP-MS pyrite trace-metal datasets show good suitability in discriminating and characterizing different gold systems and events.
-
(2)
From a genetic perspective, Au-Mo correlations and elevated Mo concentrations in the Au1 trace-metal dataset are the most significant and support an intrusion-related affinity for the Au1 event.
-
(3)
Sulfur (δ34S/Δ33S in pyrite) and oxygen (δ18O in quartz) isotope data shows poor suitability in discriminating the different hydrothermal events and in characterizing them from a genetic perspective; hence, caution is warranted in using such data for the foregoing purposes.
The involvement of a late-stage, 18O-depleted metamorphic fluid is considered to be the best explanation for the low δ18O quartz and carbonate values documented across the Wawa gold camp. Geological, geochemical, and geochronological data suggest that this fluid was generated during the 2580 ± 21 Ma Retrograde event due to devolatilization associated with the prograde metamorphism of biotite-bearing granitoids at crustal levels represented by the Wawa gneiss domain and Kapuskasing structural zone. The association of this fluid to low- to moderate-salinity H2O fluid inclusions suggests that metamorphically-devolatilized granitoid rocks may be under-recognized source components of orogenic gold deposits.
References
Alan M, Yardley B (2007) Tracking meteoric infiltration into a magmatic-hydrothermal system: a cathodoluminescence, oxygen isotope and trace element study of quartz from Mt. Leyshon, Australia. Chem Geol 240:343–360
Aleinikoff JN, Wintsch RP, Tollo RP, Unruh DM, Fanning CM, Schmitz MD (2007) Ages and origins of rocks of the Killingworth dome, south-Central Connecticut: implications for the tectonic evolution of southern New England. J Sci 307:63–118
Beaudoin G (2011) The stable isotope geochemistry of orogenic gold deposits. In: Barra F, Reich M, Campos E, Tornos F (eds) Proceedings of the Eleventh Biennial SGA Meeting, Antofagasta, pp 556–558
Beaudoin G, Pitre D (2005) Stable isotope geochemistry of the Archean Val-d’Or (Canada) orogenic gold vein field. Mineral Deposita 40:59–75
Bennett G (1972) Precambrian geology of Stover and Brackin townships. Ontario Geological Survey preliminary map P.2380
Bindeman I, Valley JW (2001) Low-δ18O rhyolites from Yellowstone: magmatic evolution based on analyses of zircons and individual phenocrysts. J Petrol:1491–1517
Bruce EL, Horwood HC (1942) Rennie–Leeson area, district of Sudbury, Ontario Ontario Geological Survey: Annual Report Map 51G
Callan NJ (1991) Syn-deformational shear zone-hosted Au-quartz vein mineralization in TTG host rocks, Renabie Mine area, North Ontario: structural analysis, microstructural characteristics and vein paragenesis. In: Ontario Geological Survey Open File Report 5759, 194 p
Callan NJ, Spooner ETC (1998) Repetitive hydraulic fracturing and shear zone inflation in an Archean granitoid-hosted, ribbon banded, au-quartz vein system, Renabie area, Ontario, Canada. Ore Geol Rev 12:237–266
Elmer FL, White RW, Powell R (2006) Devolatilization of metabasic rocks during greenschist-amphibolite facies metamorphism. J Metamorph Geol 24:497–513
Evans KA, Phillips GN, Powell R (2006) Rock-buffering of auriferous fluids in altered rocks associated with the Golden mile-style mineralization, Kalgoorlie gold field, Western Australia. Econ Geol 101:805–817
Farquhar J, Wing B (2003) Multiple sulfur isotopes and the evolution of the atmosphere. Earth Planet Sci Lett 213:1–13
Ferguson SA (1968) Renabie Mines, Nudulama Mines and adjoining properties, surface geology, parts of Renabie, Leeson and Brackin townships, District of Sudbury: Ontario Department of Mines preliminary geological map P.492
Fournier RO (1985) Theb behavior of silica in hydrothermal solutions. In: Berger BR, Bethke PM (eds) Geology and geochemistry of epithermal systems, vol 2. Reviews in economic geology, pp 45–61
Godefroy-Rodríguez M, Hagemann S, LaFlamme C, Fiorentini M (2020) The multiple sulfur isotope architecture of the Golden Mile and Mount Charlotte deposits, Western Australia. Miner Deposita 55:797–822. https://doi.org/10.1007/s00126-018-0828-y
Goldfarb RJ, Groves DI (2015) Orogenic gold: common or evolving fluid and metal sources through time. Lithos 233:2–26
Goldfarb RJ, Baker T, Dubé B, Groves DI, Hart CJR, Gosselin P (2005) Distribution, character, and enesis of gold deposits in metamorphic terranes. In: Hedenquist JW, Thompson JFH, Goldfarb RJ, Richards JP (eds) Economic Geology One Hundredth Anniversary Volume, pp. 407–450
Gourcerol B, Kontak DJ, Thurston P, Petrus J (2018a) Application of LA-ICP-MS sulfide analysis and methodology to deciphering elemental paragenesis and associations in addition to multi-stage processes in metamorphic gold settings. Can Mineral 56:1–18
Gourcerol B, Kontak DJ, Thurston P, Petrus J (2018b) LA-ICP-MS mapping, elemental parageneses and trace metal associations in sulfides from auriferous Algoma-type BIFs: implications for nature of mineralizing fluids, metal sources and deposit models. Mineral Deposita 53:871–894
Groves DI, Phillips GN, Ho SE, Houston SM, Standing CA (1987) Craton-scale distribution of Archean greenstone gold deposits: predictive capacity of the metamorphic model. Econ Geol 82:2045–2058
Groves DI, Goldfarb RJ, Robert F, Hart CJR (2003) Gold deposits in metamorphic belts: overview of current understanding, outstanding problems, future research, and exploration significance. Econ Geol 98:1–29
Groves DI, Santosh M, Deng J, Wang Q, Yang L, Zhang L (2019) A holistic model for the origin of orogenic gold deposits and its implications for exploration. Mineral Deposita 55:275–292. https://doi.org/10.1007/s00126-019-00877-5
Hagemann SG, Mariam-Gebre M, Groves DI (1994) Surface-water influx in shallow-level Archean lode-gold deposits in Western Australia. Geology 22:1067–1070
Hammerli J, Kemp AIS, Jeon H (2018) An Archean Yellowstone? Evidence from extremely low δ18O in zircons preserved in granulites of the Yilgarn Craton, Western Australia. Geology 46(5):411–414
Haroldson EL (2014) Fluid inclusions and stable isotope study of Magino: a magmatic-related Archean gold deposit. Unpublished M.Sc. thesis, The University of Wisconsin-Madison, 81 p
Hart CJR (2007) Reduced intrusion-related gold systems: In Goodfellow WD (ed) Mineral deposits of Canada: a synthesis of major deposit types, District Metallogeny, the Evolution of Geological Provinces, and Exploration Methods: Geological Association of Canada, Mineral Deposits Division, Special Publication No 5, pp 95–112
Heather KB, Arias Z (1992) Geological and structural setting of gold mineralization in the Goudreau–Lochalsh area, Wawa gold camp. Ontario Geological Survey, Open File Report 5832, 159 p
Helt KM, Williams-Jones AE, Clark JR, Wing BA, Wares RP (2014) Constraints on the genesis of the Archean oxidized, intrusion-related Canadian Malartic gold deposit, Quebec, Canada. Econ Geol 109:713–735
Hronsky JA, Groves DI, Loucks RR, Begg GC (2012) A unified model for gold mineralization in accretionary orogens and implications for regional-scale exploration targeting methods. Miner Deposita 47:339–358
Kamo SL (2015) Report on U-Pb CA-ID-TIMS geochronology on volcanic and plutonic rocks, superior and Grenville provinces, Ontario. Internal report for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, p 48
Kennedy AK, Kamo SL, Nasdala L, Timms NE (2010) Grenville skarn titanite: potential reference material for SIMS U-Th-Pb analysis. Can Mineral 48:1423–1443
Kerr MJ, Hanley JJ, Kontak DJ, Morrison GG, Petrus J, Sharpe T, Fayek M (2018) Evidence of upgrading of gold tenor in an auriferous orogenic quartz-carbonate vein system by late magmatic-hydrothermal fluids at the Madrid deposit, Hope Bay Greenstone Belt, Nunavut, Canada. Geochim Cosmochim Acta 241:180–218
Kerrich R (1989) Geochemical evidence on the sources of fluids and solutes for shear zone hosted mesothermal au deposits: In Bursnall JT (ed) Mineralization and shear zones: geological Association of Canada, Short Course Notes 6, pp. 129–197
Kerrich R (1990a) Carbon-isotope systematics of Archean Au-Ag vein deposits in the Superior province. Can J Earth Sci 27:40–56
Kerrich R (1990b) Mesothermal gold deposits: A critique of genetic hypotheses: In: Robert F, Sheahan PA, Green SB (eds) Greenstone gold and crustal evolution, NUNA Conference Volume, pp 13–27
Kirkland CL, Hollis J, Danisik M, Peterson J, Evans NJ, McDonald BJ (2017) Apatite and titanite from the Karrat Group, Greenland; implications for charting the thermal evolution of crust from the U-Pb geochronology of common Pb bearing phases. Precambrian Res 300:107–120
Kontak DJ, Horne RJ, Kyser K (2011) An oxygen isotope study of two contrasting orogenic gold vein systems in the Meguma Terrane, Nova Scotia, Canada, with implications for fluid sources and genetic models. Mineral Deposita 46:289–304
Kontak DJ, Hanley JB, Fayek M (2016) A Rayleigh distillation model to explain large variations in oxygen isotope data for orogenic gold veins. Abstract, Geological Association of Canada-Mineralogical Association of Canada
Kontak DJ, Hanley JB, Gourceol B, Petrus JA, Kelly C, Kerr M, Letourneau M, Malcolm K, McDivitt J, Neyedley K, Tokaryk S (2018) Protracted and complex fluid histories the norm in orogenic-type gold deposits as revealed by LA ICP-MS sulfide mapping. Abstract, resources for future generations (RFG) meeting, Vancouver
Kyser TK, Kerrich R (1990) Geochemistry of fluids in tectonically active crustal regions: In: Nesbitt BE (ed) Short course on fluids in tectonically active regimes of the continental crust, Mineralogical Association of Canada Short Course Handbook, 18, pp. 133–230
LaFlamme C, Jamieson JW, Fiorentini ML, Thebaud N, Caruso S, Selvaraja V (2018) Investigating sulfur pathways through the lithosphere by tracing mass independent fractionation of sulfur to the Lady Bountiful orogenic gold deposit, Yilgarn Craton. Gondwana Res 58:27–38
Li H, Schwarez HP, Shaw DM (1991) Deep-crustal oxygen isotope variations: the Wawa-Kapuskasing crustal transect, Ontario. Contrib Mineral Petrol 107:448–458
Ludwig KR (2003) Isoplot 3.00 a geochronological toolkit for Microsoft Excel. Berkeley Geochronology Center
Mao W, Rusk B, Yang F, Zhang M (2017) Physical and chemical evolution of the Dabaoshan porphyry Mo deposit, South China: insights from fluid inclusions, cathodoluminescence, and trace elements in quartz. Econ Geol 112:889–918
Matsuhisa Y, Goldsmith JR, Clayton RN (1979) Oxygen isotopic fractionation in the system quartz-albite-anorthite-water. Geochim Cosmochim Acta 43:1131–1140
McCuaig TC, Kerrich R (1998) P-T-t-deformation-fluid characteristics of lode gold deposits: evidence from alteration systematics. Ore Geol Rev 12:381–454
McDivitt JA (2016) Gold mineralization in the Missanabie-Renabie District of the Wawa subprovince: geochemical data and photographs. Ontario Geological Survey, Miscellaneous Release—Data 339
McDivitt JA, Lafrance B, Kontak DJ, Robichaud, L (2015) Characterization of gold mineralization in the Missanabie-Renabie District of the Wawa gold camp. Summary of Field Work and Other Activities, Open File Report 6313, pp. 6–1 to 6–8
McDivitt JA, Lafrance B, Kontak DJ, Robichaud L (2017) The structural evolution of the Missanabie-Renabie gold district: pre-orogenic veins in an orogenic gold setting and their influence on the formation of hybrid deposits. Econ Geol 112:1959–1975
McDivitt JA, Kontak DJ, Lafrance B, Robichaud L (2018) Contrasting fluid chemistries, alteration characteristics, and metamorphic timing relationships recorded in hybridized orebodies of the Missanabie-Renabie gold district, Archean Wawa subprovince, Ontatio, Canada. Econ Geol 113:397–420
Moser D (1989) Mid-crustal structures of the Wawa gneiss terrane near Chapleau, Ontario. Current Research, Part C. Geological Survey of Canada, Paper 88-lC, pp. 93–99
Moser D (1993) A geological, structural and geochronological study of the Central Wawa Gneiss Domain: Implications for the development of different crustal levels of the Archean Abitibi-Wawa Orogen of the Southern Superior Province, Canadian Shield. Unpublished Ph.D. thesis, Queen's University, 182 p
Moser D (1994) The geology and structure of the mid-crustal Wawa gneiss domain: a key to understanding tectonic variation with depth and time in the late Archean Abitibi-Wawa orogen. Can J Earth Sci 31:1064–1080
Neyedley K, Hanley K, Fayek M, Kontak DJ (2017) Textural, fluid inclusions, and stable O isotope constraints on vein formation and gold precipitation, 007 deposit, Bissett, Manitoba, Canada. Econ Geol 112:629–660
Ohmoto H, Rye RO (1979) Isotopes of sulfur and carbon: In Barnes HL (ed) Geochemistry of Hydrothermal Ore Deposits, 2nd edition, pp. 509–567
Ojala VJ, Groves DI, Ridley JR (1995) Hydrogen isotope fractionation factors between hydrous mineral and ore fluid at low temperatures: evidence from the Granny Smith gold deposit, Western Australia. Mineral Deposita 30:328–331
Percival JA, Card KD (1985) Structure and evolution of Archean crust in central Superior province, Canada. In: Ayres LD, Thurston PC, Card KD, Weber W (eds) Evolution of Archean Supracrustal Sequences, vol 28. Geological Association of Canada Special Paper, pp 179–192
Percival JA, West GF (1994) The Kapuskasing uplift: a geological and geophysical synthesis. Can J Earth Sci 31:1256–1286
Phillips GN, Groves DI, Brown IJ (1987) Source requirements for the Golden Mile Kalgoorlie: significance to the metamorphic replacement model for Archean gold deposits. Can J Earth Sci 24:1642–1651
Pickthorn WJ, Goldfarb RJ, Leach DL (1987) Dual origins of lode gold deposits in the Canadian Cordillera—discussion. Geology 15:471–473
Polito PA, Bone Y, Clarke JD, Mernagh TP (2001) Compositional zoning of fluid inclusions in the Archaean Junction gold deposit, Western Australia: a process of fluid—wall-rock interaction? Aus J Earth Sci 48:833–855
Powell R, Will T, Phillips G (1991) Metamorphism of Archean greenstone belts—calculated fluid compositions and implications for gold mineralization. J Metamorph Geol 9:141–150
Riley RA (1971) Precambrian geology of Glascow, Meath and Rennie townships. Ontario Geological Survey, preliminary map P.2210
Ripley EM, Li C (2017) A review of the application of multiple S isotopes to magmatic Ni-Cu-PGE deposits and the significance of spatially variable Δ33S values. Econ Geol 112:983–991
Robert F (2001) Syenite-associated disseminated gold deposits in the Abitibi greenstone belt, Canada. Min Deposita 36:503–516. https://doi.org/10.1007/s001260100186
Robichaud L, McDivitt JA, Trevisan BE (2015) Geology and Mineral Potential of Rennie and Leeson Townships, Michipicoten Greenstone Belt. Summary of Field Work and Other Activities 2015, Ontario Geological Survey, Open File Report 6313, pp.5–1 to 5–11
Rusk BG, Lowers HA, Reed MH (2008) Trace elements in hydrothermal quartz: relationships to cathodoluminescent textures and insights into vein formation. Geology 36:547–550
Sage RP (1986) Stratigraphic correlation in the Wawa area. Volcanology and Mineral Deposits, Ontario Geological Survey, Miscellaneous Paper 129, pp. 62–68
Sage RP, Heather KB (1991) The structure, stratigraphy and mineral deposits of the Wawa area. Geological Association of Canada–Mineralogical Association of Canada–Society of Economic Geologists, Joint Annual Meeting, Toronto 1991, Field Trip A6, 118p
Samson IM, Bluent B, Holm EP (1997) Hydrothermal evolution of auriferous shear zones, Wawa, Ontario. Econ Geol 92:325–342
Schlegel TU, Wagner T, Boyce A, Heinrich CA (2017) A magmatic source of hydrothermal sulfur for the Prominent Hill deposit and associated prospects in the Olympic iron oxide copper-gold (IOCG) province of South Australia. Ore Geol Rev 89:1058–1090
Schmitz MD, Bowring, SA (2001) U-Pb zircon titanite systematics of the Fish Canyon Tuff: an assessment of high-precision U-Pb geochronology and its application to young volcanic rocks. Geochim Cosmochim Acta 65, No. 15: 2571–2587
Seal RR (2006) Sulfur isotope geochemistry of sulfide minerals. Rev Mineral Geochem 61:633–677
Seedorff E, Dilles JH, Phoffett Jr JM, Einaudi MT, Zurcher L, Stavast WJA, Johnson DA, Barton MD (2005) Porphyry deposits: characteristics and Origin of Hypogene Features: In: Hedenquist JW, Thompson, JFH, Goldfarb, RJ, Richards JP (eds), Economic Geology One Hundredth Anniversary Volume, pp 251–298
Selvaraja V, Caruso S, Fiorentini ML, LaFlamme CK, Bui T-H (2017) Atmospheric sulfur in orogenic gold deposits of the Archean Yilgarn Craton, Australia. Geology 45(8):691–694
Sharman ER, Taylor BE, Minarik WG, Dubé B, Wing BA (2015) Sulfur isotope and trace element data from ore sulfides in the Noranda district (Abitibi, Canada): implications for volcanogenic massive sulfide deposit genesis. Mineral Deposita 50:591–606
Sibson RH (2001) Seismogenic framework for hydrothermal transport and ore deposition: in: JP Richards, RM Tosdal (eds) Structural controls on ore genesis, Reviews in Economic Geology, v 14, pp 25–50
Simon K (2001) Does δD from fluid inclusion in quartz reflect the original hydrothermal fluid? Chem Geol 177:483–495
Smith RB, Bruhn RL (1984) Interplate extensional tectonics of the eastern Basin-Range: Inferences on structural style from seismic reflection data, regional tectonics, and thermal mechanical models of brittle-ductile deformation. J Geophys Res 89:5733–5762
Stacey JC, Kramers JD (1975) Approximation of terrestrial lead isotope evolution by a two-stage model. Earth Planet Sci Lett 26(2):207–221
Storey CD, Jeffries TE, Smith M (2006) Common lead-corrected laser ablation ICP-MS U-Pb systematics and geochronology of titanite. Chem Geol 227(1–2):37–52
Stott GM, Corkery MT, Percival JA, Simard M, Goutier J (2010) A revised terrane subdivision of the Superior Province. Summary of Field Work and Other Activities 2010, Ontario Geological Survey, Open File Report 6260, pp 20–1 to 20–10
Sullivan RW, Sage RP, Card KD (1985) U-Pb zircon age of the Jubilee Stock in the Michipicoten Greenstone Belt near Wawa, Ontario. Current Research, Part B, Geological Survey of Canada Paper 85-1B, pp 361–365
Taylor HP Jr (1974) The application of oxygen and hydrogen isotope studies to problems of hydrothermal alteration and ore deposition. Econ Geol 69:843–883
Taylor HP Jr (1979) Oxygen and hydrogen isotope relationships in hydrothermal mineral deposits. In: Barnes HL (ed) Geochemistry of Hydrothermal Ore Deposits, 2nd edition, pp 236–277
Thompson JFH, Sillitoe RH, Baker T, Lang JR, Mortensen JK (1999) Intrusion-related gold deposits associated with tungsten-tin provinces. Mineral Deposita 34:323–334
Turek A, Smith PE, Van Schmus WR (1982) Rb-Sr and U-Pb ages of volcanism and granite emplacement in the Michipicoten Belt—Wawa, Ontario. Can J Earth Sci 19:1608–1626
Turek A, Smith PE, Van Schmus WR (1984) U-Pb zircon ages and the evolution of the Michipicoten plutonic-volcanic terrane of the Superior Province, Ontario. Can J Earth Sci 21:457–464
Turek A, Sage RP, Van Schmus WR (1992) Advances in the U-Pb zircon geochronology of the Michipicoten greenstone belt, Superior Province, Ontario. Can J Earth Sci 29:1154–1165
Turek A, Heather KB, Sage RP, Van Schmus WR (1996) U/Pb zircon ages for the Missanabie–Renabie area and their relation to the rest of the Michipicoten greenstone belt, Superior Province, Ontario, Canada. Precambrian Res 76:191–211
Valley JW (1986) Stable isotope geochemistry of metamorphic rocks. In: Valley JW, Taylor HP, O’Neil JR (eds) Stable isotopes in high temperature geological processes, vol 16. Rev Mineral Geochem, pp 445–490
Williams HR, Stott GM, Heather KB, Muir TL, Sage RP (1991) Wawa subprovince. In: Thurston PC, Williams HR, Sutcliffe RH, Stott GM (eds) Geology of Ontario, special volume 4 (part 1). Ministry of Northern Development and Mines, Ontario Geological Survey, Sudbury, pp 485–538
Xue Y, Campbell I, Ireland TR, Holden P, Armstrong R (2013) No mass-independent sulfur isotope fractionation in auriferous fluids supports a magmatic origin for Archean gold deposits. Geology 41:791–794
Yuguchi T, Sasao E, Ishibashi M, Nishiyama T (2015) Hydrothermal chloritization process from biotite in Toki granite, central Japan: temporal variations of the compositions of hydrothermal fluids associated with chloritization. Am Min 100:1134–1152
Zheng Y (1993) Calculation of oxygen isotope fractionation in hydroxyl-bearing silicates. Earth Plan Sci Lett 120:247–263
Acknowledgments
GoldTrain Resources Inc. is thanked for allowing access to drill core and stripped outcrops. Ryan Sharpe, with the SIMS laboratory at the University of Manitoba, is thanked for his role in the oxygen isotope analysis. We sincerely acknowledge the constructive feedback provided by Georges Beaudoin, Iain Samson, Peter Hollings, and an anonymous reviewer.
Funding
This study is financially and logistically supported by the Ontario Geological Survey. This manuscript stems from a M.Sc. thesis completed by JM at Laurentian University that was financially supported through the Ontario Geological Survey-Laurentian University Mapping School Agreement. Additional financial contributions to JM are provided by the Natural Sciences and Engineering Research Council (NSERC), the Society of Economic Geologists Foundation, and the Goodman School of Mines.
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McDivitt, J.A., Kontak, D.J., Lafrance, B. et al. A trace metal, stable isotope (H, O, S), and geochronological (U-Pb titanite) characterization of hybridized gold orebodies in the Missanabie-Renabie district, Wawa subprovince (Canada). Miner Deposita 56, 561–582 (2021). https://doi.org/10.1007/s00126-020-00983-9
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DOI: https://doi.org/10.1007/s00126-020-00983-9