Keywords

1 Introduction

Orogenic gold deposits comprise a diverse class of mineral deposits that contain the second largest resource of gold, after the Witwatersrand (Republic of South Africa) paleoplacers (Dubé and Gosselin 2007). In addition to their economic importance, orogenic gold deposits record important geological process active at mid-crustal levels, where the deformation behavior of rocks changes, with increasing depth, from brittle to ductile. The formation of orogenic gold deposits involves advection of significant volumes of crustal fluids, which carry large amounts of dissolved mass and energy at the scale of the crust.

The class of orogenic gold deposits was proposed by Groves et al. (1998) to combine gold deposits that shared numerous characteristics, yet had remained grouped under different deposit types in literature. The orogenic gold deposits were defined to include a range of epigenetic gold styles variably named “mesothermal”, “quartz-carbonate”, “lode gold”, “greenstone gold”, “Mother Lode”, “turbidite-hosted”, “BIF (Banded-Iron Formation)-hosted”, etc. The class of orogenic gold deposits was defined to include deposits where gold is the major economic metal, and which formed in volcanic and/or sedimentary, accretionary, metamorphic terranes in convergent settings of all ages. The geological characteristics of orogenic gold deposits have been reviewed by Goldfarb et al. (2005) and Goldfarb and Groves (2015) amongst others, and are briefly summarized hereafter. Orogenic gold deposits consist of gold-bearing quartz-carbonate veins or disseminated sulfides, commonly representing less than 5 vol.% sulfides, most commonly pyrite and arsenopyrite. The veins display a range of structural features from foliated shear bands to brittle breccia and are characteristically associated to steeply-dipping reverse faults, related to major crustal shear zones typically marking terrane boundaries. The country rocks are varied and include almost all types of volcanic, plutonic, sedimentary rocks and their metamorphosed equivalents. Hydrothermal alteration is characteristically narrow and controlled by the vein structures, replacing peak-metamorphic minerals by white mica, Fe–Mg carbonate minerals and sodic feldspar, being the most common. The veins formed from low-salinity (3–12 wt.% NaCl eq.) carbonic-aqueous hydrothermal fluids with \({\text{X}}_{{\text{CO}}_{2} }\)  < 0.3, at temperatures typically about 350 °C. Orogenic gold deposits formed at depths of less than 20 km, typically near the base of the seismogenic continental crust, under high, to supra-lithostatic, fluctuating fluid pressure. Deposits formed deeper in the crust are characterized by ductile deformation, with gold mineralization mostly in disseminated sulfide minerals or less commonly in foliated veins, and formed at the higher temperature range. Those formed higher in the crust are characterized by more brittle structures, and a lower temperature of formation.

In spite of these well-defined geological characteristics, the origin of hydrothermal fluids that formed orogenic gold deposits remains controversial. Goldfarb and Groves (2015) reviewed sources for fluids in orogenic gold deposits, stating that the stable isotope composition of the minerals is inconsistent between deposits, and leads to equivocal interpretations about the source of fluids and dissolved components. This is particularly intriguing since the auriferous hydrothermal fluids are considered to have a well-defined composition that should enable determination of the source(s) of fluids. Goldfarb and Groves (2015) reiterated a long-held view that fluid-rock exchange, at the site of reaction, could mask part of the original isotopic signature (Ridley and Diamond 2000). If the association of the orogenic gold deposits with a major shear zone is a well-established fact, the role of the shear zone has been contested from being a deep-seated fluid channelling feature (e.g., McCuaig and Kerrich 1998) to a breach into a mid-crustal low permeability barrier (Beaudoin et al. 2006).

The first major compilation of oxygen, H, carbon, sulfur stable isotope data for Au–Ag deposits in metamorphic rocks, now defined largely as orogenic gold deposits, was undertaken by Kerrich (1987). This led to the concept of provinciality in the oxygen isotope composition of orogenic gold deposits, inferred to indicate large and uniform fluid source rock volumes, but different for each orogenic gold vein field (McCuaig and Kerrich 1998). High δ18O values are characteristic of the auriferous fluids forming orogenic gold deposits, with a range of δ18OH2O values of 6 to 11‰ for Precambrian deposits (McCuaig and Kerrich 1998) and of 7 to 13‰ for Phanerozoic deposits (Bierlein and Crowe 2000), and as shown by the range of quartz δ18O values of 6.9 to 25‰ for the 25 largest orogenic gold deposits (Goldfarb et al. 2005).

Goldfarb et al. (2005) compiled δD values ranging from −150 to −26‰ for mica and from −98 to −24‰ for fluid inclusions from the 25 largest orogenic gold deposits. Using the likely temperatures of isotopic equilibrium, the hydrothermal fluids δD values were suggested to range from −80 to −20‰, excluding low values from fluid inclusions of dubious or mixed origins (McCuaig and Kerrich 1998; Goldfarb et al. 2005). Combined with the estimated fluid δ18O values of 6 to 14‰, this has been taken as a strong indication for a metamorphic source for the auriferous fluids (McCuaig and Kerrich 1998).

The carbon isotope composition of hydrothermal carbonates associated with gold mineralization was shown to range mostly between −10 and 0‰, with outlier values as low as −32‰ (Goldfarb et al. 2005). McCuaig and Kerrich (1998) also remarked the narrow range in δ13C values for a given deposit. The isotope composition of carbon has led to interpretation of a number of potential sources, from the mantle to magmas, to crustal devolatilization, mixing and redox reactions (McCuaig and Kerrich 1998). Likewise, the sulfur isotope composition of sulfides are summarized to mostly range from 0 to 9‰, but with a total range from −20 to 25‰, interpreted to reflect local sulfur sources (Goldfarb et al. 2005).

Despite the low variance in geological features, and in the chemical PVTX characteristics of auriferous fluids associated with orogenic gold deposits, the variation of the stable isotope composition of hydrothermal minerals and fluid inclusions has led to various, at times conflicting or equivocal, interpretations. Here, we attempt to address the variations in the stable isotope composition of hydrothermal minerals and fluid inclusions in orogenic gold deposits on the basis of a detailed analysis of an extensive compilation of data published between 1960 and 2010. The study of the database allows to outline major stable isotope compositional features for orogenic gold deposits to yield improved understanding of the sources of fluids, carbon, sulfur, nitrogen, boron and silicon for this deposit type.

2 Methodology

2.1 Database

The database integrates fifty years of stable isotope and orogenic gold deposit research published between 1960 and 2010, since which time data interpretation has been ongoing (Beaudoin 2011). Review of data published since 2010 shows no special features from more recent studies, but specific cases are discussed in relation to pre-2010 data. The first study of stable isotopes in orogenic gold systems was on the sulfur isotope composition of gold quartz veins of the Yellowknife district, Canada (Wanless et al. 1960). Because gold-bearing quartz ± carbonate veins have been described using a range of descriptive names from lode to quartz-carbonate to orogenic gold veins, the search in literature has used all known terms used to identify stable isotope research for this deposit type. The deposits compiled in the database comprise all the major orogenic gold deposits as defined by Goldfarb et al. (2005). Some of the deposits aggregated in the compilation have controversial metallogenic affinity. Nevertheless, the amount of data from deposits of controversial metallogenic affinity is small, such that it should not significantly impact on the overall stable isotope characteristics of orogenic gold deposits, unless otherwise noted. Data were compiled from 193 references, for which tabulated stable isotope values were reported (Supplementary Material Spreadsheet 1). Data reported in figures or as range of values, without individual mineral composition, were not considered in this study. The data were compiled dominantly from peer-reviewed journals, with the exception of a few major deposits for which data were only available in university theses.

The database contains 8580 stable isotope analyses for oxygen (n = 4014), hydrogen (n = 706), carbon (n = 1244), sulfur (n = 2215), nitrogen (n = 254), boron (n = 117) and silicon (n = 30) from 5478 samples (Supplementary Material Spreadsheet 1). Stable isotopic compositions are expressed in ‰, and are reported relative to VSMOW, VPDB, VCDT, AIR, NIST SRM 951, and NBS28 for δ18O and δD, δ13C, δ34S, δ15N, δ11B and δ30Si, respectively. Data reported in literature relative to SMOW, PDB and CDT are considered identical to the newer IAEA Vienna (V-) standards. Where samples have the same name, and several isotope or mineral analyses are reported, the results are deemed to be from the same sample even though textural equilibrium can seldom be verified. Oxygen isotope compositions have been measured in 22 minerals. Two studies report δ18O values from fluid inclusions (FI) hosted by quartz or scheelite (Santosh et al. 1995; Billstrom et al. 2009), which are not considered further because the δ18O values are most likely equilibrated with the host mineral oxygen. The hydrogen isotope composition has been measured in seven minerals and from FI in quartz. Mao et al. (2008) report seven δD values from “K-feldspar”. Because the source of hydrogen in theses analyses is unspecified, but likely to be from FI, these data were not included in the database. Carbon isotopic compositions have been measured in seven minerals and in FI trapped in quartz and scheelite. Sulfur isotopic compositions have been measured in 13 minerals. Nitrogen and silicon isotopic composition have been measured in three and two minerals, respectively. Boron has been measured in tourmaline. The samples are from 564 deposits from all continents with the exception of Antarctica (Fig. 1), and range in age from the Archean to the Cenozoic, including 180 Archean, 68 Proterozoic, 104 Paleozoic, 129 Mesozoic, and 83 Cenozoic deposits. For Archean deposits, most of them are Neo-Archean (n = 164), with fewer Paleo- (n = 2) and Meso-Archean (n = 13). The age of the gold deposits is known with variable accuracy, such that geochronological data yield precise age for some deposits, whereas only age inferences can be made for others.

Fig. 1
figure 1

Location of orogenic gold deposits with stable isotope data. The world geology map is prepared using digital data from Gosselin and Dubé (2005)

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Gold deposits are hosted in a range of country rock types including sedimentary, volcano-sedimentary, or igneous rocks with variable chemical composition from felsic to ultramafic. For deposits hosted by several types of rocks, the dominant lithology/chemical composition is recorded in the database. Sedimentary rock-hosted deposits include unmetamorphosed clastic and/or chemical sedimentary rocks (greywackes, sandstones, shales, limestones, mudstones, argilites, BIF) or their metamorphosed equivalents (paragneiss, schist) which are all reported as “sedimentary” for simplicity. Volcano-sedimentary rock-hosted deposits are mainly represented by Archean and Proterozoic greenstone belts. Igneous rock-hosted deposits include plutonic or volcanic rocks. Volcanic rocks mainly consist of rhyolite to basaltic rocks in greenstone belts, as well as mafic and ultramafic dykes. Plutonic rocks include a wide variety of rocks such as tonalite-trondhjemite-granodiorite (TTG) batholiths as well as granitic—diorite intrusions in various geological environment. Based on the description provided in each publication, geochemical descriptors (felsic, intermediate, mafic and ultramafic) are used to summarize the composition of dominant intrusive and volcanic rocks.

Deposits associated with zeolite, prehnite-pumpellyite or sub-greenschists facies, and/or “low” metamorphic grade rocks, have been combined into a sub-greenschist category. Those hosted by lower, middle, or upper greenschist facies rocks have been classified into a greenschist category. Lower, middle, or upper amphibolite, and granulite metamorphic facies host rocks are categorized into a high-grade metamorphic category. The resulting categories must be used with caution since orogenic gold deposits commonly post-date regional metamorphism. Moreover, in numerous studies, the metamorphic grade is not accurately described or covers a range, such as “lower greenschist to amphibolite”. In such cases, we arbitrarily attributed a deposit to the higher metamorphic grade. Finally, for most of deposits from the North China Craton, a metamorphic grade class is not attributed because the gold deposits formed during the Mesozoic, significantly younger than Precambrian regional granulite grade metamorphism.

2.2 Temperature of Equilibrium

There can be several isotopic fractionations for a pair of minerals (e.g. δquartz–δmineral) (Beaudoin and Therrien 2009) such that the computed temperatures can vary depending on the selected fractionation. Oxygen isotope fractionations derived from previous studies by Vho et al. (2019) are used for temperature estimates from quartz-tourmaline, quartz-chlorite, quartz-biotite, quartz-muscovite, quartz-albite, quartz-plagioclase, quartz-feldspar, quartz-ankerite, quartz-calcite and quartz-dolomite mineral pairs and are reported with a 2σ confidence interval (Supplementary Material, Spreadsheet 2). For quartz-sericite,Footnote 1 no fractionation equation exists in literature and, consequently, we use the quartz-muscovite fractionation equation from Vho et al. (2019). For quartz-scheelite, only one fractionation equation is available (Zheng 1992), but it yields low temperatures (mainly between 110 and 200 °C), inconsistent with the greenschist facies conditions and other geothermometers. Consequently, a combination of two equations, quartz-H2O from Clayton et al. (1972) and scheelite-H2O from Wesolowski and Ohmoto (1986), has been used to estimate the temperature of equilibrium between quartz and scheelite.

For quartz-plagioclase, quartz-feldspar and quartz-carbonate pairs, no information about the composition of the mineral is provided. Consequently, data are not presented in δquartz–δmineral diagrams and temperatures have not been calculated.

For sulfide pairs, pyrite-galena, pyrite-chalcopyrite, pyrite-pyrrothite, pyrite-sphalerite, sphalerite-chalcopyrite and chalcopyrite-galena fractionation data are from Kajiwara and Krouse (1971). The fractionation from Liu et al. (2015) was used for sphalerite-galena. The data are plotted in δmineral–δmineral diagrams overlain by isotherms ranging from 250 to 550 °C.

2.3 Oxygen and Hydrogen Isotope Composition of Water

Oxygen and hydrogen isotope compositions of H2O are calculated from quartz and OH-bearing minerals pairs that yield equilibrium temperatures in the range 250–550 °C. This temperature range is consistent with the accepted range in temperature of formation for orogenic gold deposits (Goldfarb and Groves 2015). Thus, it is presumed that those mineral pairs record O–H isotope equilibrium. The water isotopic composition has been calculated using various hydrogen mineral-H2O fractionations and oxygen quartz-H2O fractionation equation from Vho et al. (2019). The entire dataset is presented in Supplementary Material 2. For tourmaline and chlorite, the fractionations of Kotzer et al. (1993) and Graham et al. (1984) have been used, respectively. For biotite and muscovite (or sericite) the hydrogen fractionations from Suzuoki and Epstein (1976) were used.

2.4 Carbon Isotope Composition of CO2

The carbon isotope composition of CO2 was calculated using calcite-CO2 and dolomite-CO2 fractionations from Ohmoto and Rye (1979) assuming equilibrium temperatures in the range 250–550 °C as determined from quartz-calcite and quartz-dolomite oxygen fractionation data (Vho et al. 2019). Because there are no carbon ankerite-CO2 fractionation in literature, we did not compute δ13C of CO2 from ankerite. The entire dataset is presented in Supplementary Material 2.

3 Stable Isotope Composition of Minerals and Fluid Inclusions

3.1 Oxygen Isotopes

3.1.1 Silicates, Tungstates, Borates and Oxides

The oxygen isotope composition of 2357 quartz samples from orogenic gold deposits ranges from 5.2 to 25.9‰ (Fig. 2). Quartz δ18O values show a broad, bimodal, distribution with one mode near 11.5‰, and the second near 15‰ (Fig. 2).Footnote 2 Quartz from orogenic gold deposits formed during the Archean shows this bimodal distribution, whereas quartz from deposits formed during the Proterozoic are characterized by δ18O values about the lower mode (11.6‰; Fig. 2a). Orogenic gold deposits formed during the Phanerozoic have quartz δ18O values about the higher mode (15‰) and tailing towards the highest δ18O values (Fig. 2a).

Fig. 2
figure 2

Histograms showing δ18O values of quartz forming veins from orogenic gold deposits based on: a the age of deposits; b the craton or orogen hosting the deposits (only craton or orogen with more than 50 δ18O values are shown); c country rock of deposits; and d the geochemistry of igneous rocks hosting the deposits

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Figure 2b shows the distribution of quartz δ18O values for orogenic terranes for which there is more than 50 analyses. Quartz from the Archean Yilgarn Craton (Australia) shows a normal distribution about the 11.5‰ mode, with few higher values. In contrast, quartz from the Archean Superior Craton (Canada) displays a bimodal distribution. Quartz from the Proterozoic Trans-Hudson Orogen (Canada) plots mostly about the 11.5‰ mode. Quartz from the Appalachian/Caledonian and Hercynian orogens (America and Europe) plots mostly near 13.5‰ whereas those from the Tasman Orogen (Australia) are characterized by higher values with a mode near 16.5‰. Mesozoic and Cenozoic orogenic gold deposits from the Cordillera (North America) are characterized by heavier quartz oxygen isotope compositions mainly between 12.5‰ and 19.5‰ (Fig. 2b). Mesozoic orogenic gold deposits from the Daebo Orogen and North China Craton (China, Korea) plot mostly near 13‰.

Figure 2c shows the distribution of all quartz δ18O values as function of the principal country rock type. Quartz hosted in igneous rocks shows the lowest δ18O values with a normal distribution centered at 11.5‰. Deposits hosted in sedimentary country rocks have higher δ18O values with a normal distribution centered at 15‰. The δ18O values for deposits hosted in volcano-sedimentary rocks display a bimodal distribution with modes centered at 12‰ and 15‰. Most of the high δ18O values are from deposits hosted by sedimentary and volcano-sedimentary rocks. The distribution of δ18O values for quartz hosted in igneous rocks, divided by geochemical affinity, is presented on Fig. 2d. No major difference is identified between mafic and felsic country rocks, both showing a unimodal distribution about the major 11.5‰ mode. Vein quartz hosted in ultramafic rocks ranges between 11.5‰ and 15.5‰, without a well-defined mode.

Quartz from orogenic gold districts from the Archean Superior Craton displays distinct oxygen isotopic compositions. The lower δ18O values, between 9‰ and 14‰, with a mode centered on 11‰ are characteristic of the Val-d’Or district, whereas the higher δ18O values between 13‰ and 16‰, with a mode at 15‰ are mostly from the Timmins district (Fig. 3a). There is no grouping of δ18O values for other districts perhaps because of the smaller number of analyses. The difference in δ18O values between the Val-d’Or and the Timmins districts correlates with the distribution of δ18O values as function of the type of country rocks (Figs. 3a, b). Quartz from veins hosted in igneous rocks have δ18O values ranging from 5 to 22‰ with a main mode at 11‰, whereas those hosted in sedimentary displays higher δ18O values, between 5 and 27‰, with a mode at 14.5‰, and those hosted in volcano-sedimentary rocks have values between 7 and 24‰, with modes at 11.5 and 14.5‰. As shown in Fig. 3c, the igneous rocks hosting deposits with high quartz δ18O values are mainly ultramafic.

Fig. 3
figure 3

Histograms showing δ18O values of quartz forming veins from orogenic gold deposits from the Superior Craton based on: a district; b country rocks; c geochemistry of igneous rocks hosting deposits; and d deposits from the Val-d’Or district

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For the Yilgarn Craton, quartz δ18O values range mainly between ~10 and ~14‰, with no distinction between districts (Fig. 4a). Most of these veins are hosted in mafic igneous rocks and the few hosted in volcano-sedimentary rocks have similar δ18O values (Figs. 4b, c). In districts from the Cordilleran Orogen for which there are more than 20 analyses, quartz displays δ18O values mostly between ~12 and ~20‰ (Fig. 4d) higher than those documented from the Superior and Yilgarn cratons. In the Cordilleran Orogen, the Fairview, Klondike and Bridge River districts (12–14‰, 14–16‰ and 18–20‰, respectively) in Canada, the Berners Bay (12 to 16‰) and Mother Lode districts (16–19‰; United States of America—USA) display distinct ranges of δ18O values. In contrast, the δ18O values from the Nome district (USA) define a large range of values (10 to 20‰). All these districts are hosted in volcano-sedimentary and sedimentary rocks but in some of the districts, major deposits are hosted by granite (Fig. 4e). Data do not vary with the chemistry of the igneous rocks (Fig. 4f).

Fig. 4
figure 4

Histograms showing δ18O values of quartz forming veins from orogenic gold deposits from the Yilgarn Craton and Cordilleran Orogen based on: a district (Yilgarn Craton); b country rocks (Yilgarn Craton); c geochemistry of igneous rocks hosting deposits (Yilgarn Craton); d districts (Cordilleran Orogen); e country rocks (Cordilleran Orogen); and f geochemistry of igneous rocks hosting deposits (Cordilleran Orogen)

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At the district scale, quartz from orogenic gold deposits is characterized by narrow ranges of δ18O values (Fig. 3d). For example, in the Val-d’Or vein field, quartz from the Beaufor deposit is characterized by quartz oxygen isotope compositions in the range 9.9–11.1‰, with an average of 10.6 ± 0.2‰ (Beaudoin and Pitre 2005). Few studies allow comparison of several deposits in one district. In the Val-d’Or vein field, Beaudoin and Pitre (2005) and Beaudoin and Chiaradia (2016) showed no systematic variation in composition across vein thickness, along strike, or down-dip. In Val-d’Or, there is a gradual change in the average composition of quartz of a deposit at the scale of the district (Beaudoin and Pitre 2005).

Figure 5 presents the oxygen isotope composition of other silicates, tungstate, borates and oxides. Muscovite (n = 84) and sericite (n = 70) have δ18O values ranging from 7.0 to 17.0‰ and 6.6 to 16.0‰, respectively (Fig. 5a), with one muscovite having a low δ18O value of 0.7‰. Chlorite (n = 87) and biotite (n = 45) have similar δ18O values ranging from 1.0 to 12.0‰ and from 2.0 to 12.6‰, respectively. Two talc samples yield δ18O values of 7.0 and 7.9‰ (Fig. 5a). Scheelite (n = 54) has a small range of δ18O values from 2.0 to 6.6‰, whereas tourmaline (n = 40) displays a wider range of δ18O values from 4.9 to 11.0‰ (Fig. 5b). Actinolite (n = 19) has δ18O values from 5.0 to 11.5‰, whereas two epidote samples yield δ18O values of 7.7 and 9.6‰ (Fig. 5b). Undifferentiated feldspar (n = 20) and plagioclase (n = 20) have δ18O values ranging from 3.6 to 11.3‰ and from 5.5 to 13.4‰, respectively, whereas albite (n = 17) displays δ18O values from 8.3 to 17.8‰ (Fig. 5c). Magnetite (n = 14) and hematite (n = 8) have low δ 18O values from −4.1 to 7.6‰ and from −2.3 to 2.2‰, respectively (Fig. 5d).

Fig. 5
figure 5

Histograms showing δ18O values of silicates and oxides minerals in deposits: a biotite, chlorite, muscovite, sericite and talc; b actinolite, epidote, scheelite, and tourmaline: c albite, feldspar and plagioclase; and d hematite and magnetite

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3.1.2 Carbonates

Figure 6 shows the oxygen isotope composition of carbonate minerals, including calcite (n = 493), dolomite (n = 345), ankerite (n = 207), siderite (n = 53), magnesite (n = 8) and undifferentiated carbonate minerals (n = 42). Carbonate δ18O values show three modes, with two principal modes centered on 10‰ and 12.5‰ (Fig. 6a), and a third mode centered on 24‰. Ankerite, calcite and dolomite have δ18O values ranging from 7.7 to 25.0‰, from 6.3 to 27.5‰, and from 1.2 to 25.3‰, respectively, with the same bimodal distribution for ankerite and calcite. Dolomite displays a bimodal distribution with two modes centered on 12‰ and 16‰. Siderite has δ18O values ranging from 7.7 to 29.8‰, but most of data are around 11‰. The few δ18O values of magnesite and undifferentiated carbonate minerals do not show consistent distribution patterns.

Fig. 6
figure 6

Histograms showing δ18O values of carbonate from veins in orogenic gold deposits according to: a carbonate mineral; b age of deposit; c craton or orogen hosting deposits (only craton or orogen with more than 30 δ18O values are shown); d country rocks of deposits; and e geochemistry of igneous rocks hosting deposits

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Archean deposits display a bimodal distribution of δ18O values with two principal modes centered on 10‰ and 12.5‰ (Fig. 6b). Proterozoic deposits do not display a specific distribution of δ18O values, whereas Paleozoic deposits display a bimodal distribution with two principal modes centered on 12.5‰ and 14.5‰. Mesozoic deposits have two main modes of δ18O values at 14.5‰ and 16‰. Values higher than ~20‰, are mostly from Paleozoic deposits with only few Archean, Proterozoic and Mesozoic analyses. Figure 6c shows that δ18O values for Archean cratons (Kaapvaal, Slave, Superior, Yilgarn) are mostly lower than ~15‰. The distribution of the δ18O values from Yilgarn is centered on 10‰, whereas data from the Kaapvaal, Slave and Superior cratons are mostly centered on 12.5‰. For the Appalachian/Caledonian Orogen, δ18O values are scattered about a large mode near 13‰. For others cratons/orogens, data are less numerous and scattered. Figure 6d displays carbonate δ18O values as function of the country rock type. Deposits hosted in igneous rocks have a bimodal distribution with two modes about 10‰ and 12.5‰. Values for deposits hosted in sedimentary rocks display a normal distribution centered on ~14‰, whereas those hosted in volcano-sedimentary environment display three modes centered on ~10‰, ~12‰ and ~14‰. Figure 6e shows carbonate δ18O values as function of the chemical composition of igneous country rocks. Deposits hosted in felsic to mafic rocks have two modes about 10‰ and 12.5‰. Deposits hosted in ultramafic rocks have a normal distribution of carbonate δ18O values centered on 12.5‰.

3.1.3 Hydrogen Isotopes in Silicates

The hydrogen isotope composition (Fig. 7a) has been measured in muscovite (n = 86), sericite (n = 51), biotite (n = 32), chlorite (n = 42), actinolite (n = 9), epidote (n = 5), and tourmaline (n = 13). The range in δD values for vein minerals is remarkably wide, from −187 to −4‰ (Fig. 7a). The distribution of vein mineral δD values is bimodal, with a principal mode near -58‰, and a second mode near −85‰, with an asymmetric tail to low δD values. Epidote has the highest δD values (−30 to −4‰). Muscovite and sericite are characterized by δD values that spread mainly between −80 and −26‰, with a mode near −55‰, but with a wide spread to low δD values (−186‰; Fig. 7a). Chlorite displays a main mode at slightly lower values compared to muscovite and sericite (between −90‰ and −70‰), but also yields low δD values (−187‰). Biotite displays a relatively small range of δD values from −117 to −51‰ (Fig. 7a). Tourmaline and actinolite have δD values that span the range of the two main modes of hydrogen isotope compositions.

Fig. 7
figure 7

Histograms showing δD values in OH-bearing minerals from veins from orogenic gold deposits based on: a nature of mineral; b age of deposit; c craton or orogen hosting deposits; d country rocks of deposits; and e geochemistry of igneous rocks hosting deposits

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In Fig. 7b, δD values of minerals are reported relative to the age of the deposits. The bimodal distribution about the principal (−58‰) and the second (−85‰) modes encompasses δD values of all ages. However, Mesozoic and Cenozoic deposit δD values display a unimodal distribution centered around -58‰, with an asymmetric tail to the lowest δD values for deposits from the Cordillera Orogen. Deposits from Trans-Hudson Orogen display a unimodal distribution with values ranging between ~−120 and ~−30‰, and a mode at ~−90‰. For other cratons/orogens, data are less numerous and scattered.

Figures 7d and e show the distribution of δD values as function of country rock and of chemical composition of igneous rocks, respectively. No trends were identified, but the lowest values documented for Mesozoic and Cenozoic deposits from the Cordillera Orogen are from deposits hosted in sedimentary and volcano-sedimentary rocks.

3.1.4 Hydrogen Isotopes in Fluid Inclusions

Figure 8 presents the δD values of fluid inclusions extracted from quartz (n = 468). The range in δD values for fluid inclusions is from −179 to −7‰, with a wide and asymmetric distribution towards low δD values (Fig. 8a). The mode is centered on −70‰. Hydrogen isotope compositions of fluid inclusions from Archean, Proterozoic, Paleozoic and Cenozoic orogenic gold deposits are typically heavier, from −7 to ~−110‰, with few outlying values to ~−150‰ (Fig. 8a). Fluid inclusions δD values for Mesozoic orogenic gold deposits typically have values lower than −40‰, with a bimodal distribution with a mode near −70‰ and a second mode centered on −160‰ (Fig. 8a).

Fig. 8
figure 8

Histograms showing δD values from fluid inclusion (FI) trapped in quartz from veins from orogenic gold deposits based on: a age of deposit; b craton or orogen hosting deposits (only craton or orogen with more than 20 δD values are shown); c country rocks of deposits; and d geochemistry of igneous rocks hosting deposits

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Figure 8b shows the distribution of δD values from fluid inclusions as function of craton/orogen. The main mode near −70‰ is dominated by deposits from the Daebo and Tianshan orogens, and from the North China Craton Craton. Deposits from the Himalayan Orogen display a wide range of δD values, with a secondary mode at ~−110‰ from three deposits in Tibet (Supplementary material, Spreadsheet 1). The low δD values (< −120‰) are documented mainly for deposits hosted in the Cordillera Orogen, with a few measurements from deposits hosted in the Baltic Shield and the Trans-Hudson Orogen (Supplementary material, Spreadsheet 1).

Figure 8c shows δD values as function of the type of country rocks. The lowest δD values are characteristic of volcano-sedimentary rocks hosting deposits in the Mesozoic Cordillera Orogen. The dominant country rocks of deposits with δD values higher than −120‰ are igneous and sedimentary with only few analyses for deposits hosted in volcano-sedimentary rocks. In Fig. 8d, δD values between −120‰ and −60‰ are mainly from deposits hosted by felsic to intermediate igneous country rocks, whereas mafic and ultramafic hosted deposits have fluid inclusions with δD values between −60‰ and −20‰.

3.1.5 Carbon Isotopes in Minerals and Fluid Inclusions

Figure 9a shows the distribution of the carbon isotopic compositions for calcite (n = 513), dolomite (n = 344), ankerite (n = 214), siderite (n = 44), magnesite (n = 7), undifferentiated carbonate minerals (n = 47), and graphite (n = 2). Most δ13C values range from −15 to 3‰ with a wide mode at −4‰, a secondary mode around −14‰ and a third one at −22‰. The δ13C values from fluid inclusions trapped in quartz and scheelite show a unimodal distribution centered near −6‰ with rare values as low as −26‰ (Fig. 10a).

Fig. 9
figure 9

Histograms showing δ13C values of carbonate minerals and graphite from orogenic gold deposits based on: a nature of minerals; b age of deposit: c craton or orogen hosting deposits (only craton or orogen with more than 30 δ13C values are shown); d country rocks of deposits; and e geochemistry of igneous rocks hosting deposits

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There is no significant difference of δ13C values as function of the mineral or fluid inclusions (Figs. 9a and 10a). In Archean deposits, carbonates have δ13C between −8 and 3‰, whereas Proterozoic, Paleozoic and Mesozoic deposits have similar δ13C values distributed from −15 to 3‰ with a mode at −5‰ (Fig. 9b). However, low δ13C values are documented for Paleozoic deposits of the Meguma district hosted in sedimentary rocks (Appalachian/Caledonian Orogen; Fig. 9b, c, d).

Fig. 10
figure 10

δ13C values of CO2 from fluid inclusion (in quartz and scheelite) and CO2 in equilibrium with calcite and dolomite using temperatures of equilibrium calculated using O isotope fractionation of Vho et al. (2019) on quartz-carbonate minerals pairs (see Fig. 13 and Supplementary material, spreadsheet 2). Histograms showing \({\updelta }^{{13}} {\text{C}}_{{\text{CO}}_{2} }\) values are presented based on: a mineral; b age of deposit; and c country rocks of deposit. d shows \({\updelta }^{{13}} {\text{C}}_{{\text{CO}}_{2} }\) values of various geological reservoirs (Taylor 1986; Kerrich 1989; Trull et al. 1993)

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Carbonates from Archean Slave, Superior, Kaapvaal and Yilgarn cratons and from the Mesozoic North China Craton Craton display δ13C values between −7 and 0‰, whereas the range of δ13C values from the West African Craton is centered on −13‰ (Fig. 9c).

The carbonate δ13C values from deposits hosted by igneous rocks show a unimodal distribution with a mode near −4‰ (Fig. 9d). Volcano-sedimentary rock-hosted deposits display a narrow range of δ13C values, between −10 and 0‰. Sedimentary rock hosted deposits show a multimodal distribution with a main mode at −6‰, and three other modes around −3, −13 and −22‰. Felsic igneous rocks hosted deposits show a unimodal distribution with a main mode at −4‰. The carbonate δ13C values from deposits hosted in mafic and ultramafic rocks, mainly from −9 to 0‰, each show a distribution with a main mode near −3‰ (Fig. 9e).

3.1.6 Sulfur in Sulfides and Sulfates

The sulfur isotope composition of sulfides ranges from −30 to 17‰, with a mode at ~ 3‰ and a tail to low δ34S values. (Fig. 11a). Anhydrite (n = 3) and barite (n = 29) δ34S values range from −2.9 to 18.2‰ with a mode at ~8‰. Most δ34S values are from pyrite (n = 1305), arsenopyrite (n = 258) and galena (n = 199). No variation of δ34S in relation to sulfide species is identified. The δ34S values are similar for deposits formed during the Archean and Proterozoic (Fig. 11b). However, Paleozoic age deposits show a bimodal distribution of δ34S values with a main mode at 0‰ and a second mode at 3‰. Mesozoic and Cenozoic age deposits display a distribution of δ34S values similar to that of Archean and Proterozoic deposits, although with a wider range to low and high δ34S values.

Fig. 11
figure 11

Histograms showing δ34S values of vein sulfide minerals from orogenic gold deposits based on: a mineral; b age of deposit; c craton or orogen hosting deposit (only craton or orogen with more than 50 δ34S values are shown); d country rocks of deposits; e geochemistry of igneous rocks hosting deposits; and f age of deposits in sedimentary country rocks. g Variations in δ34S values for various geological reservoirs (Machel et al. 1995; Wortmann et al. 2001; Seal 2006; Johnston et al. 2007; Labidi et al. 2013). \(\Delta^{{34}} {\text{S}}_{{\text{SO}}_{4} - {\text{H}}_{2} {\text{S}}}^{2 - }\) due to thermochemical sulfate reduction (TSR) calculateed at 500 °C and 100 °C (Eldridge et al. 2016)

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The distribution of δ34S values is homogeneous for all cratons/orogens, and most of the data are spread along the entire range as illustrated for the Slave, Superior and Daebo cratons/orogen (Fig. 11c). The range of δ34S values from the West African Craton is narrower, from −12 to −5‰. However, these data are from six deposits and have been measured mainly on arsenopyrite (Fig. 11a). Similarly, the Trans-Hudson Orogen and Dharwar Craton display a narrow distribution from 2 to 10‰. No significant difference has been identified for δ34S values from deposits hosted by igneous, sedimentary or volcano-sedimentary rocks (Fig. 11d). However, sulfide minerals from deposits hosted in sedimentary rocks display δ34S variations with age (Fig. 11e). Archean deposits have δ34S values with a unimodal distribution at 2.5‰, whereas Proterozoic deposits display a unimodal distribution at 4‰. Paleozoic deposits display a bimodal distribution with two main modes at 0‰ and 2.5‰ and with a number of values around 9‰. Mesozoic deposits display scattered δ34S values with several data about 2.5‰ and tailing to low values. There are not enough data from Cenozoic age deposits to describe the distribution. The δ34S values of minerals hosted by mafic to ultramafic igneous rocks show a bimodal distribution, with a main mode near 1.5‰ and a second mode at −5‰ (Fig. 11e). The range of δ34S values for deposits hosted by felsic igneous rocks is slightly higher than for mafic rocks, between −5‰ and 15‰ with a mode at 3‰.

3.1.7 Nitrogen in Silicates

The N isotope composition of biotite (n = 15), muscovite (n = 208) and K-feldspar (n = 31) range from 1 and 24‰ (Fig. 12a). Archean deposits have mostly δ15N values higher than 10‰. Proterozoic deposits display a range between 7 to 12‰ whereas Phanerozoic deposits have mostly δ15N values below 7‰ (Fig. 12b). Figure 12c shows that δ15N values labelled by craton/orogen mimics the distribution as a function of the age of the deposits. Figure 12d shows that low δ15N values are mostly documented from deposits hosted in volcano-sedimentary and sedimentary rocks whereas higher values are mainly from igneous rock hosted deposits.

Fig. 12
figure 12

Histograms showing δ15N values of minerals from veins from orogenic gold deposits based on: a mineral; b age of deposit; c craton or orogen hosting deposit; and d country rock of deposits. e δ15N values of various geological reservoirs are (Peters et al. 1978; Haendel et al. 1986; Minagawa and Wada 1986; Bebout and Fogel 1992; Compton et al. 1992; Williams et al. 1995; Wu et al. 1997; Ader et al. 1998; Jia and Kerrich 2000; Sephton et al. 2002; Gu 2009)

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3.1.8 Boron in Tourmaline

The B isotope composition of tourmaline (n = 117) ranges from −21.6 to 9‰ (Fig. 13a). The δ11B values for Proterozoic deposits range from −16 to −9‰ (Figs. 13a, b). Archean deposits hosted in the Dharwar Craton have δ11B values that spread mainly between −5 and 9‰, with a bimodal distribution with a main mode about −2.5‰ and a second mode about 4‰. The δ11B values for the Archean Yilgarn Craton are the lowest and range between −24 and −18‰. These data are similar to δ11B values of tourmaline compiled by Trumbull et al. (2020) from orogenic gold deposits worldwide, ranging between −24.8 and 19.8, with two modes at −15 and −2‰.

Fig. 13
figure 13

Histograms showing δ11B values of tourmaline from veins from orogenic gold deposits based on: a the age of the deposit; and b the craton or orogen hosting deposits. c δ11B values of various geological reservoirs (Palmer and Swihart 1996; van Hinsberg et al. 2011; Lambert-Smith et al. 2016; Marschall et al. 2017; Trumbull and Slack 2018)

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3.1.9 Silicon in Silicates

The silicon isotope composition of quartz (n = 19) and sericite (n = 11) range from −0.5 to 0.8‰ and from −0.3 to 0.4‰, respectively (Fig. 14a). The limited data is from two deposits, Hemlo (Superior Craton) of controversial origin (Ding et al. 1996), and Sawaya’erdun (Tianshan Orogen) (Liu et al. 2007), such that the representativeness of the range of δ30Si values is limited. No difference in composition is identified for δ30Si values in relation to host mineral and deposit age (Fig. 14b).

Fig. 14
figure 14

Histograms showing δ30Si values of minerals from vein from orogenic gold deposits based on: a mineral; and b age of deposit. c δ30Si values of various geological reservoirs (Poitrasson 2017 and reference therein)

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4 Isotope Equilibrium and Temperature of Formation

The isotope composition of a mineral depends on intrinsic and extrinsic factors. Intrinsic factors are related to the crystal structure, mass of the atoms forming molecular bounds in minerals, and isotope diffusion coefficients. These intrinsic factors determine partitioning of light and heavy isotopes of an element between two coexisting phases, the isotopic fractionation. The extrinsic factors comprise the temperature of the system, the pressure, particularly for gaseous phases, the bulk isotopic composition of the system and its components, and the isotope exchange reaction rate which, if fast, enables close system isotopic equilibrium to be reached or, if slow, will yield a kinetic open system in isotopic disequilibrium. The extrinsic factors determine the magnitude of the fractionation between isotopic species, temperature being the most important.

In order to make interpretations from the isotopic composition of minerals, it is thus essential to verify the assumptions about the state of equilibrium of the system in order to use theoretical or experimental fractionations to derive the temperature of equilibrium, or the bulk isotopic composition of the system and its components. Verification of equilibrium is best achieved comparing the composition of isotopic phases using δ–δ diagrams (Criss et al. 1987). In δ–δ diagrams, two coexisting phases in a closed system, minerals for example, will plot on the isotherm of the temperature of equilibrium, which is determined by the fractionation between the two species. If mineral pairs formed at the same temperature, but in isotopic systems with different bulk compositions, the data will plot along the same isotherm. If minerals formed at different temperatures in a system with a bulk constant isotopic composition, data will plot in an array orthogonal to isotherms.

δ–δ diagrams are drawn for quartz-silicate/borate/tungstate (Fig. 15), quartz-carbonate (Fig. 16), and sulfide-sulfide (Fig. 17). Quartz-muscovite mineral pairs from the same sample form a linear array that is parallel to isotherms (Fig. 15a). The data array is centered on the 350 °C isotherm calculated using Vho et al. (2019), which indicates that most quartz-muscovite pairs are approaching isotopic equilibrium near 350 °C. Few mineral pairs plot outside the array, which indicate either isotopic equilibrium at higher/lower temperatures, or isotopic disequilibrium. Using the fractionation of Vho et al. (2019), 60 of the 68 samples plot between the 250 and 550 °C isotherms. An apparent lower temperature of equilibrium is documented for Mesozoic age deposits (Fig. 15a), but this trend is related to the fact that five of the seven quartz-muscovite with temperatures below 350 °C are from the Alleghany district, Cordilleran Orogen (Böhlke and Kistler 1986). Quartz-sericite pairs show a similar distribution (Fig. 15b), but data are mostly centered along the 450 °C quartz-muscovite isotherm. Quartz-biotite pairs (Fig. 15c) show more scatter than quartz-muscovite, but are mostly centered on the 350 °C isotherm. Quartz-chlorite pairs (Fig. 15d) show a similar distribution than that of quartz-biotite pairs, although a number of samples yield large Δqz-chl, up to 12‰, which clearly indicates these mineral pairs were not in isotopic equilibrium. Quartz-albite (Fig. 15e) pairs plot mainly parallel to 250 °C isotherm. Quartz-tourmaline pairs (Fig. 15f) form an array centered on the 450 °C isotherm using Vho et al. (2019), but would be around the 350 °C isotherm using the fractionation of Kotzer et al. (1993), which has been defined using Proterozoic and Archean orogenic gold-bearing veins. Several samples have small (~1‰) or large (< 7‰) Δqz-tur values that indicate that these samples are not in isotope equilibrium. Finally, quartz-scheelite pairs scatter in the δ–δ diagram (Fig. 15 g), with several pairs with small (~3‰) or large (< 12‰) Δqz-sch values that indicate that these samples are not in isotope equilibrium. For all quartz-silicate/borate/tungstate pairs, there is no obvious trends with the composition of the country rocks, or age of the deposit.

Fig. 15
figure 15

δ18Osilicates versus δ18Oquartz of coexisting vein minerals with symbols indicating the age and the type of country rocks. Isotherms drawn using oxygen isotope fractionation equations from Vho et al. (2019) for a, b, c, d, e, g and h and from Clayton et al. (1972) and Wesolowski and Ohmoto (1986) for f

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Fig. 16
figure 16

δ18Ocarbonate versus δ18Oquartz of coexisting vein minerals with symbols indicating the age and the type of country rocks. Isotherms drawn using O isotope fractionation equations from Vho et al. (2019) for a, and c, and from Clayton et al. (1972) and Wesolowski and Ohmoto (1986) for b

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Fig. 17
figure 17

δ34SSulfide versus δ34SSulfide of coexisting vein sulphide minerals with symbols indicating the age and the type of country rocks. Isotherms drawn using S isotope fractionation equations from Kajiwara and Krouse (1971) for a, b, d and f, from Li and Liu (2006) for c and e, and from Liu et al. (2015) for g

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Quartz-ankerite mineral pairs form a broad array parallel to isotherms (Fig. 16a), but with a large proportion of samples with fractionations too small or too large to record isotopic equilibrium at geologically reasonable temperatures (250–550 °C; (Goldfarb et al. 2005). Both quartz-calcite (Fig. 16b) and quartz-dolomite (Fig. 16c) also form broad arrays along isotherms, but data distribution also shows a trend of high δ18O values for calcite and dolomite (< 25‰), that indicates isotope disequilibrium with coexisting quartz in these samples.

Because the magnitude of Δ34S between sulfide minerals at equilibrium temperatures between 250 and 550 °C is small (e.g., 1.6 to 0.7‰ for pyrite-chalcopyrite), small Δ34S variations yield large variations of the calculated temperature of equilibrium, such that temperature estimation using sulfur isotopes is less accurate. Pyrite-chalcopyrite (Fig. 17a) pyrite-sphalerite (Fig. 17b) and pyrite-galena (Fig. 17c) pairs similarly plot as broad arrays parallel to isotherms. For these mineral pairs, only a small proportion of samples show small or large fractionations that indicate sulfur isotope disequilibrium. Most data plot close to the 350 °C isotherm indicating isotope equilibrium. A few pyrite-pyrrhotite pairs scatter in the δ-δ diagram (Fig. 17d), indicating sulfur isotope disequilibrium. Sphalerite-chalcopyrite (Fig. 17e) and sphalerite-galena (Fig. 17f) pairs, similar to most pyrite-sulfide pairs, plot in broad arrays centered along the 350 °C isotherm. Finally, the galena-chalcopyrite δ–δ diagram (Fig. 17 g) shows a broad array parallel to isotherms, but at low calculated temperatures (circa 150 °C).

In summary, quartz-silicate, quartz-carbonate, and sulfide-sulfide mineral-pairs that display evidence for isotopic equilibrium between the 250–550 °C yield an average temperature of 360 ± 76 °C (σ = 1, n = 332) without evidence for secular change with the age of formation of then orogenic gold deposits.

5 Discussion

5.1 Secular Variations in Mineral and Fluid Inclusion Compositions

Figure 2a shows that δ18O values of quartz from Archean (8–16‰) and Proterozoic (8–16‰) age deposits broadly overlap with Paleozoic (9–17.5‰), Mesozoic (10–22.5‰) and Cenozoic (10–17.5‰) age deposits, yet Phanerozoic age deposits yield slightly higher δ18O values (9–22.5‰) than Precambrian age deposits (8–16‰). Figure 6b, likewise shows that carbonate minerals from Archean and Proterozoic age deposits typically have lower δ18O values (7–15‰) than that of Paleozoic (9–25‰) and Mesozoic (8–20‰) age deposits. The change in quartz and carbonate oxygen isotope composition with age of deposit (Figs. 2a and 6b), is not continuous with a reversal to lower δ18O values for Cenozoic age deposits, casting doubt on a secular evolution of the composition of hydrothermal fluids in orogenic gold deposits, as suggested by Goldfarb and Groves (2015). As shown in Fig. 2c, the lower Archean quartz δ18O values are dominantly for orogenic gold deposits hosted by igneous rocks, whereas Phanerozoic age deposit higher δ18O values are characteristic for deposits hosted by sedimentary and volcano-sedimentary country rocks. Figure 6d shows a similar pattern where carbonate with lower and higher δ18O values correspond to Archean and Proterozoic age deposits dominantly hosted, respectively, by igneous or sedimentary and volcano-sedimentary country rocks.

Less abundant data for other silicate minerals do not allow analysis by age of deposit. In contrast, carbonate mineral δ13C values do not display a distribution by age of deposit, with the exception of the low δ13C values from deposits of the Appalachian Orogen, which indicates this a provincial feature (Fig. 9b). Similarly, silicate (Fig. 7b) and inclusion fluidFootnote 3 (Fig. 8b) δD values do not display different distributions based on the age of the deposits, with the exception of the low values from the Mesozoic Cordillera Orogen deposits. Sulfide sulfur in deposits hosted by igneous and volcano-sedimentary rocks does not display compositional variations related to the deposit age (Fig. 11). In contrast, deposits hosted in sedimentary rocks show different but overlapping ranges of δ34S values with age. Chang et al. (2008) showed that Phanerozoic orogenic gold deposits hosted in sedimentary rocks have sulfide δ34S values that closely track that of seawater sulfate at the age of the host rocks. This result was interpreted to record leaching of reduced seawater sulfate from the sedimentary country rocks. Likewise, Goldfarb et al. (1997) showed that the Juneau district Cenozoic orogenic gold deposits have δ34S values that follow those of their Phanerozoic sedimentary host rocks. Thus, the sulfur isotope composition of orogenic gold deposits hosted in sedimentary rocks vary with the age of the host rocks following the secular variation of seawater sulfate, but with no evidence of secular variation related to the age of the deposits.

The less abundant data for N (Fig. 12b), B (Fig. 13a) and silicon (Fig. 14b) allows for less definitive assessments of secular variations in composition. The N isotope composition is characteristically higher for Archean (11–24‰) and Proterozoic (7–12.5‰) compared to Paleozoic- (2.5–14‰) and Mesozoic- to Cenozoic-aged deposits (5–7‰: Fig. 12b). The decrease in δ15N values with decreasing deposit age has been ascribed to progressive mantle degassing (δ15N = −5‰ + /−2‰) into the atmosphere (δ15N = 0‰), and the gradual sequestration of heavier atmospheric N2 into sedimentary rock micas (Jia and Kerrich 2004; Kerrich et al. 2006). However, other studies suggest that atmospheric δ15N values were constant through time and that isotopic variations in sedimentary rocks reflect isotopic fractionation related to fluid-rock interaction and/or biogenic activity (Cartigny and Marty 2013).

The compilation shows that Archean age deposits have mostly higher δ11B values than Proterozic age deposits (Fig. 13a). However, recent studies show that several Archean deposits also have low tourmaline δ11B values, such as in Val-d’Or (Beaudoin et al. 2013; Daver et al. 2020) and the Hattu schist belt (Molnár et al. 2016). Finally, the limited amount of silicon isotope compositions does not display a trend with age (Fig. 14b).

5.2 Temperature Variations

5.2.1 Isotopic Disequilibrium

Most minerals pairs plot between the 250–550 °C isotherms, temperatures typical for the formation of orogenic gold deposits, thus suggesting widespread isotopic equilibrium for quartz-silicates/borate/tungstate and sulfide-sulfide pairs. In contrast, quartz-carbonate pair δ18O values are scattered with only few pairs along 250–550 °C isotherms suggesting common isotopic disequilibrium.

Diachronous formation of quartz and carbonate could explain such disequilibrium, although quartz and carbonate are commonly paragenetically coeval in orogenic gold deposits. Isotopic exchange can also occur between mineral-pairs during cooling by volume diffusion (Giletti 1986). As shown by Sharp and Kirschner (1994), the degree of retrograde oxygen diffusion between quartz and calcite depends of the closure temperature of quartz which, in turn, depends of its grain size, oxygen diffusion coefficient and cooling rate. They showed that for quartz grains in a large reservoir of calcite, resetting below 400 °C will only occur for small quartz grain sizes. Consequently, retrograde diffusion exchange between quartz and calcite in veins is unlikely to explain scattered data in δQtz–δCarb space (Fig. 16). Late infiltration of low temperature fluids can significantly alter the δ18O value of calcite and disturb the initial isotopic equilibrium between co-genetic minerals (Sharp and Kirschner 1994). This low temperature resetting has been proposed to explain the common isotopic disequilibrium documented for carbonates with high δ18O values compared to quartz (Kontak and Kerrich 1997; Beaudoin and Pitre 2005): Fig. 16). Because veins commonly experienced successive deformation events (as folded or sheared veins), such late fluid circulation can easily be channelized in newly formed fractures.

5.2.2 Secular Variations in Temperature of Formation of Orogenic Gold Deposits

Goldfarb et al. (2005) suggested that Phanerozoic and Paleoproterozoic orogenic gold deposits formed at lower temperatures (250–350 °C) that older Archean deposits (325–400 °C). However, they acknowledged that each age group of orogenic gold deposits displayed large, and overlapping, ranges of temperature of formation. For mineral pairs showing isotope equilibrium, there is no secular trend of changing equilibrium temperature with age of deposit (Figs. 15, 16 and 17). This observation is independent of the fractionation used to compute the temperature, as the different fractionations will only shift the nominal value of the temperature, not the data distribution along isotherms. Thus, the stable isotope composition of vein minerals does not support the interpretation of a lower temperature of formation for Proterozoic and Phanerozoic orogenic gold deposits, compared to higher temperature Archean age orogenic gold deposits.

5.3 Composition of Hydrothermal Fluids and Dissolved Elements

In the following sections, the δ18O and δD values of H2O and δ13C values of CO2 are calculated from minerals isotopic composition at temperatures of 250–550 °C for mineral pairs that show evidence of isotope equilibrium. δ34S values of H2S have not been calculated because isotope fractionation between sulfides and H2S are small (< 1‰) between 250 and 550 °C (Li and Liu 2006). The source of H2S is discussed based on δ34S values of sulfides.

5.3.1 Water

Metamorphic versus magmatic water. The two most common sources of fluids proposed for orogenic gold deposits are magmatic and metamorphic water. As reviewed by Goldfarb and Groves (2015), the inconsistent timing between magmatism and orogenic gold deposit formation, and lack of a specific magmatic association, both argue against the ubiquity of magmatic fluids in the formation of orogenic gold deposits.

Most of the calculated water compositions plot in the field for metamorphic water, but a small proportion of analyses also plot in the field for magmatic water and a few plot outside both fields (Figs. 18a, b; Sheppard 1986). Magmatic water has been argued in some cases (Li et al. 2012; Zeng et al. 2014; Deng et al. 2015), but this interpretation is based on fluid δ18O and δD values between −0.2 and 6‰ and −96 and −52‰, respectively, that plot largely outside of the magmatic water field, and are too low for the inferred granitic magma source. Thus, metamorphic water is the more common fluid source, and in most cases the only likely fluid source for the formation of orogenic gold deposits. The variation of oxygen and hydrogen isotopes composition of metamorphic water(s) documented on Fig. 18 probably reflect that fluids were sourced from metamorphism of variable proportions of sedimentary and igneous rocks in the crust.

Fig. 18
figure 18

Diagrams showing the variation of δ18O and δD values of H2O calculated from quartz and OH-bearing minerals. a δDH2O versus δ18OH2O with symbols indicating the age of deposits, and b the type of country rocks. The high δD values are interpreted to reflect the effect of evaporation–condensation cycles, whereas higher δ18O reflect O isotope buffering by fluid-rock interactions. The Global Meteoric Water Line is derived from Craig (1961), whereas the metamorphic and magmatic water boxes are from Sheppard (1986)

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Evidence of meteoric water in fluid inclusions. Fluid inclusions are trapped either during crystal growth (primary) or in cracks during late fluid circulation (secondary). Caution must be exercised using isotope composition of fluid inclusions because the mechanical or thermal decrepitation leach methods extract all populations of fluid inclusions, some of which not associated with gold formation. It has been shown by Faure (2003) than thermal decrepitation yield inaccurate δD values, possibly because of uncontrolled reactions between released H2O and Si–OH bounds in quartz, for example. Likewise, Gaboury (2013) showed that fluid inclusions have different hydrocarbon compositions as a function of the temperature of decrepitation. As an example, Foley et al. (1989) showed bulk extract contained two types of fluid inclusions with distinct δD values at Creede (Colorado, USA). This is particularly likely in orogenic gold deposits where veins are commonly deformed and thus may have trapped various fluids during and/or after mineralization. For example, in laminated veins, most fluid inclusions are secondary (Ridley and Diamond 2000). Pickthorn et al. (1987) reviewed evidence for discrepancy between fluid inclusion and mica δD values and concluded to mixing of various generations of fluid inclusions, including late meteoric water trapped in secondary inclusions in the host minerals. This was contested by Nesbitt et al. (1987) on the basis that various mixtures of fluid inclusions should yield a mixing line between two end members, contrary to the low variance of δD values.

As shown on Fig. 19, micas are commonly out of isotopic equilibrium with fluid inclusions trapped in quartz from the same vein. Isotopic disequilibrium is well documented for chlorite and some sericite/muscovite displaying high δD values (~-60‰) compared to the δD values of coexisting fluid inclusion (<-60‰). Such disequilibrium has been interpreted to record infiltration of late, low δD meteoric water or a mixture of several types of fluid inclusions with different hydrogen isotope compositions (Goldfarb et al. 1991). Fluid inclusions (Figs. 8a) and OH-bearing minerals (Fig. 7b) low δD values are mainly documented from Mesozoic and Cenozoic deposits, and are commonly interpreted to result from late meteoric water infiltration disconnected with the formation of the deposits (Nesbitt et al. 1989; Zhang et al. 1989; Madu et al. 1990; Goldfarb et al. 1991; Shaw et al. 1991; Rushton et al. 1993; Apodaca 1994; Jia et al. 2003). However, similar low δD values are also documented for older deposits. In deposits from Trans-Hudson terrane (Figs. 8a, b), Liu (1992) considered that meteoric water circulated in the vein structure, but after the formation of the deposit. In contrast, Billstrom et al. (2009) suggest that low δD values in Baltic Shield Paleoproterozoic age deposits could reflect infiltration of surface water during gold deposition.

Fig. 19
figure 19

δDMineral versus δDFI coexisting OH-bearing mineral and fluid inclusions from a same sample. Isotherms have been drawn using mineral-H2O H isotope fractionations of Suzuoki and Epstein (1976) for muscovite and biotite and from Graham et al. (1987) for chlorite. The grey area represents the uncertainty associated to the chlorite-H2O H isotope fractionation of Graham et al. (1987) between −40‰ and −30‰ between 200 and 500 °C

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Fluid Mixing. Figures 20a, b, c and d show the variation of δ18OH2O and δDH2O values as a function of the temperature of equilibrium between pairs of minerals. For both isotopes, data is spread along a linear trend between low δ18O (~2‰), high δD (~10‰) values at low temperatures (~250 °C), and high δ18O (~12‰), low δD (~-100‰) values at high temperatures (~550 °C) of equilibrium. This trend is similar to that documented by Beaudoin and Chiaradia (2016) for quartz and tourmaline from orogenic veins from the Val-d’Or vein field, which was interpreted to result from mixing between a low temperature upper crustal fluid and a high temperature deep-seated crustal fluid. Data for orogenic gold deposits (Figs. 20a,b,c,d) show that the same mixing trend between similar fluid end-members is recorded for deposits of all ages (Archean to Cenozoic), hosted in different country rock types, suggesting that fluid mixing between two common reservoirs may be an important process documented in orogenic gold deposits of all ages, worldwide, yet poorly documented in literature. The high δ18O—low δD—high T deep-seated crustal fluid is likely metamorphic in origin, whereas the low δ18O—high δD—low T upper crustal fluid is water of surficial origin with a long history of water–rock reactions, perhaps not unlike long residence water in cratons, as reviewed by Warr et al. (2021). The ultimate origin of the surficial water, obscured by water–rock exchange, may be difficult to decipher in several deposits. High δDH2O values, up to 0‰, in the Wiluna deposit (Yilgarn) have been interpreted to result from mixing of surface water, at shallow levels, with metamorphic-magmatic fluids (Hagemann et al. 1994). Boiron et al. (2003) showed that the deep-seated fluid was the dominant reservoir involved in the first step of the vein formation followed by progressive infiltration of surficial water during basement uplift. It is known that crustal rocks porosity contains water deep in the crust (Smithson et al. 2000). This water may have been trapped since rock formation or may have infiltrated from surface (seawater or meteoric) slowly reacting with country rocks along the fluid pathways. If fluid mixing has not been commonly identified in orogenic gold districts, the data compilation indicates that most deposits, in which we can interpret a temperature of equilibrium and a H–O fluid composition from mineral pairs, plot along a broad array that we interpret as a mixing line between two common reservoirs. Beaudoin et al. (2006) showed that the regional change in oxygen isotope composition of orogenic gold deposits of the Val-d’Or vein field can be reproduced by simulating fluid flow and oxygen isotope transport and reaction of the deep-seated fluid in rocks saturated by the upper crustal fluid end-members. Infiltration of the deep-seated fluids along the higher permeability shears and fractures of the orogenic gold deposits resulted in mixing with the upper crustal fluids filling the host rocks porosity, thus yielding the mixing lines shown in Fig. 20. Mixing of various proportions of fluids from different reservoirs also explains the gradual change in composition between deposits in one district, as shown in the Val-d’Or (Beaudoin and Pitre 2005) and Victoria (Gray et al. 1991) vein fields.

Fig. 20
figure 20

Diagrams showing the variation of δ18O of H2O calculated from quartz and OH-bearing minerals as function of temperature calculated from mineral pairs based on: a the age of deposits; and b the type of country rocks. Diagrams showing the variation of δD of H2O calculated from quartz and OH-bearing minerals as function of temperature calculated from mineral pairs based on: c the age of deposits and d the type of country rocks. The broad linear arrays in a, b, c and d suggest mixing between a low temperature, low δ18O, high δD (interpreted as an upper crustal) and a high temperature, high δ18O, low δD (interpreted as a deep crustal) fluids. Diagrams showing the variation of δ18O of H2O calculated from quartz and OH-bearing minerals as function of temperature calculated from mineral pairs based on: e the district; and f the type of country rocks. The shift in δ18O values between different districts could suggest provinciality on the fluid, that is that the deep-seated fluid reservoir varies slightly in composition between orogenic gold vein districts

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Fluid buffering. As shown in Figs. 2c, 6d, the highest quartz and carbonate δ18O values are documented in deposits hosted dominantly in volcano-sedimentary and sedimentary rocks, consistent with the interpretation that isotopic oxygen exchange between fluids and sedimentary rocks leads to 18O-enriched quartz (Böhlke and Kistler 1986; Goldfarb et al. 1991, 2004; Boer et al. 1995; Zhang et al. 2006) and carbonates (Böhlke and Kistler 1986; Steed and Morris 1997). For carbonates, some values enriched in 18O also likely reflect a late, retrograde, and low-temperature re-equilibration with a surface-derived fluid.

Figures 18b and 20b shows that most δ18OH2O values higher than those along the mixing trend, are from deposits hosted in sedimentary or volcano-sedimentary rocks. This suggests oxygen isotope buffering of the fluids by sedimentary and volcano-sedimentary country rocks, which is detected in sedimentary rocks because of their heavier oxygen isotopic composition (Sheppard 1986).

Provinciality. Figures 20e and f present δ18OH2O values as function of temperature for the Archean Yilgarn (Kalgoorlie and Coolgardie districts) and Superior (Val-d’Or, Timmins, Larder Lake, and Kirkland Lake districts) cratons. Data displays two trends between a common low δ18OH2O—high δD—low temperature upper crustal fluid endmember, and two slightly different deep-crustal fluid endmembers distinguished by their oxygen isotope composition. Deep crustal fluid #1 is characteristic of the Kirkland Lake, Larder Lake, and Timmins Superior Craton districts, and has a δ18OH2O about 4‰ higher than deep crustal fluid #2, which is defined by deposits of Superior Craton Val-d’Or, and Yilgarn Craton Kalgoorlie and Coolgardie, districts. This suggests that the deep-seated fluids were sourced from rocks at depth with a slightly different oxygen isotope composition, likely reflecting different proportions of igneous and sedimentary rocks in the crustal segment undergoing prograde metamorphism, such that the deep-seated fluid reservoir varies slightly in composition between orogenic gold vein districts. The difference in end-member composition cannot be an artifact of water–rock reactions because deposits from both trends are hosted dominantly by igneous rocks (Fig. 20f).

Origin of high δD fluids. The formation of the gold-bearing quartz vein is commonly associated with the precipitation of micas in the veins and host rock, or micas formed by hydration of host-rock minerals (Groves et al. 1998; Ridley and Diamond 2000; Eilu and Groves 2001; Craw et al. 2009). Muscovite and chlorite are the two main micaceous minerals in orogenic gold deposit. Here, we model the evolution of the H–O isotopic composition of water from which micas precipitate progressively in an open system following Rayleigh distillation to determine if mica precipitation can explain the deuterium enrichment, up to 20‰, of the hydrothermal fluids in orogenic system (Fig. 18e, f). Isotopic exchange between fluid and hostrock is not considered. Equations 1 and 2 are modified from Sharp (2017):

$${\updelta }_{{\text{Liq}}} = [{\updelta }_{{\text{Liq}}} + {1}000]{\text{F}}^{\left( {{\upalpha }_{{\text{Chl}}/{\text{Ms}} - {\text{Liq}}}{{\text{ }} - 1} } \right)} -{1}000,$$
(1)

and

$${\updelta }_{{\text{Chl}}/{\text{Ms}}} = {\upalpha }_{{\text{Chl}}/{\text{Ms}} - {\text{Liq}}} \left( {{\updelta }_{{\text{Liq}}} + {1}000} \right)-{1}000,$$
(2)

where F is the remaining water fraction, at 350 °C and using 1000ln αchlorite-H2O = −35‰ (Graham et al. 1987) for H, 1000ln αchlorite-H2O = –0.7‰ (Wenner and Taylor 1971) for O, 1000ln αmuscovite-H2O = –37.8‰ (Suzuoki and Epstein 1976) for H, and 1000ln αmuscovite-H2O = 1.3‰ (Vho et al. 2019) for oxygen.

The composition of water in equilibrium with chlorite and/or muscovite during progressive crystallization of micas follows a D-enrichment trend toward high δD values (Fig. 21c). A 20‰ enrichment in deuterium is reached after consumption of ~40–45% hydrogen to form chlorite or muscovite. During chlorite precipitation, there is a ~0.4‰ 18O enrichment of the residual fluid, for a δD enrichment of ~20‰, whereas muscovite precipitation causes a decrease of fluid δ18O values of about −0.7‰ for the same deuterium enrichment. In a δD–δ18O space, muscovitization yields a trend with a steep negative slope whereas chlorite precipitation yields a steep positive slope. The change in δD of fluids during muscovite precipitation could contribute to explain high δD values of the hydrothermal fluid as shown in Fig. 18a.

Fig. 21
figure 21

a Conceptual model of the effect of chloritization/sericitization of country rock on the ore fluid. b Conceptual model of the successive boiling-condensation cycles related to the fault valve model (Robert et al. 1995). Vapor formation from an over-pressured fluid (F0) can occur during seismic events. Vapor (V1) is expected to migrate upward, faster than residual liquid and then to condense entirely (F1) during re-sealing of fault. This process can be repeated numerous (n) times because of the successive re-opening of the fault during seismic cycles. c Quantification of the vector of enrichment in deuterium and depletion in 18O of the evolved fluid (Fn) computed for the different models

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Seismic rupture of fault veins occurs after a pre-seismic stage during which fluid pressure increases progressively to reach and overcome the lithostatic pressure (Sibson et al. 1988). Seismic rupture occurs once the increasing shear stress reaches a critical value, which then triggers progressive drainage of the overpressured fluid upward along the fault, and an abrupt fluid pressure decrease. This leads to increased effective stress on the fault plane that seals the fault, during which vein mineral precipitates, which reduces permeability of the fault and yield a new cycle of increasing fluid pressure. This cycle can be repeated numerous times, as attested by the typical crack and seal texture of orogenic veins, in the so-called fault-valve model (Sibson et al. 1988; Robert et al. 1995; Cox 2005). It has been shown that these pressure fluctuations can induce water vapor phase separation (Wilkinson and Johnston 1996; Ridley and Diamond 2000). It is likely that lower density vapor will migrate upward along a high permeability structure, separating from the residual liquid in a vein opening. Upon sealing of the vein, the vapor will condense under increasing pressure. The boiling-condensation cycles can be summarized as follow:

$${\text{F}}_0 \to {\text{V}}_{1} \to {\text{F}}_{1} \to {\text{V}}_{2} \to {\text{F}}_{2} \to \ldots {\text{V}}_{\text{n}} \to {\text{F}}_{\text{n}} ,$$
(3)

where F0 is the initial fluid, Vn the composition of vapor separated from Fn − 1, which then condenses in a new fluid Fn and n the number of boiling-condensation cycles (Fig. 21b).

We model the isotope effects during boiling-condensation open-system Rayleigh isotopic fractionation for vapor formation from a fraction of the liquid and assume complete condensation of vapor. This means that the new fluid Fn will have the same isotopic composition as the vapor Vn using Eqs. 4 and 5, modified after Sharp (2017):

$$\rm{\delta }_{\text{Liq}} = \left[ {\rm{\delta }_{{\text{Liq}},{\text{i}}} + 1000} \right]{\text{F}}^{\left( {{\upalpha }_{{\text{Vap}} - {\text{Liq}}}{ {\text{ }}- 1} } \right)} - 1000,$$
(4)

and

$${\updelta }_{{\text{Vap}}} = {\upalpha }_{{\text{Vap}} - {\text{Liq}}} \left( {{\updelta }_{{\text{Liq}}} + {1}000} \right)-{1}000.$$
(5)

Using isothermal conditions at 350 °C, considering 10% or 40% of vapor formation from liquid for each successive vapor separation episode, and oxygen and hydrogen isotope fractionation between vapor and liquid from Horita and Wesolowski (1994).

The two scenarios (10 and 40% vapor separation) yield similar linear trends with higher δD and lower δ18O values for the condensed water, along a vector with a slope of −10.5 (Fig. 21c). The proportion of water evaporation for each cycle (10% or 40%) only changes the magnitude of the isotopic shift (Fig. 21c), and consequently the number of cycles required to reach 20‰ enrichment in deuterium of residual water, and a corresponding depletion of −2‰ in 18O, similar to the maximum fluid δD values in orogenic gold deposits (Figs. 18a, b). Using 10% of evaporation, ~11 cycles are required to reach the δD value of 20‰, whereas ~20 cycles are needed with 40% evaporation steps (Fig. 21c). The modeled boiling-condensation trend is consistent with positive δD values (~20‰) combined with smaller 18O depletion (~-4‰) of the fluid recorded by equilibration fractionation with hydrogen bearing minerals (Fig. 18a). The boiling-condensation trend is consistent with data with δDH2O > 0‰ from the Val-d’Or vein field (Beaudoin and Chiaradia 2016) showing similar deuterium enrichment (δD from 0‰ to 29‰) and 18O depletion (δ18O from 5.9‰ to 2.9‰) of mineralizing fluid.

5.3.2 Source(s) of CO2

CO2 plays an important role in the formation of orogenic gold deposits as testified by (i) the common carbonation of country rocks, (ii) the association of gold with quartz-carbonate veins, and (iii) the occurrence of CO2-rich fluid inclusions trapped in orogenic quartz veins (Smith et al. 1984; Ho et al. 1992). CO2 has the capacity to buffer the pH of the fluid in a range where the gold complexation with reduced sulfur is increased (Phillips and Evans 2004). Nevertheless, the origin of CO2 remains debated. Four main carbon reservoirs have been proposed to explain the wide variation of δ13C values of CO2; (i) reduced organic carbon in sedimentary rocks (such as graphite, organic matter) with low δ13C values −26‰, Kerrich (1989); (ii) deep-seated mantle-derived carbon with δ13C ranging from −10.8 to −1.6‰ (Trull et al. 1993), (iii) magmatic CO2 (δ13C from −5 to −2‰; (Taylor 1986), and (iv) seawater-derived carbonate with a δ13C value near 0‰.

Figure 10 shows the distribution of δ13C values CO2 either from fluid inclusion trapped in quartz and scheelite or calculated from quartz-carbonate pairs with equilibrium temperatures in the range 250–550 °C. The δ13C values of CO2 display a similar range of values, between −25 and 6‰, as well as a similar distribution, mainly between −12 and 0‰ (Fig. 10a). No trend between δ13C values and the age of deposit is documented, although all Archean age deposits have δ13CCO2 values above −6‰.

The positive δ13C values are commonly interpreted as metamorphic CO2 released during decarbonation reaction of a source with δ13C near 0‰, such as sedimentary carbonate rocks (Oberthuer et al. 1996; Chen et al. 2008; Scheffer et al. 2017). The low δ13CCO2 values (< −8‰) are mainly documented from deposits hosted in sedimentary rocks (Fig. 10c). This is particularly well illustrated by the Meguma district from the Appalachian/Caledonian Orogen (Fig. 9c), where carbonate minerals have negative δ13C values (−25.9‰ < δ13C < −18.2‰), which was interpreted by Kontak and Kerrich (1997) to represent the formation of 13C-depleted CO2 by oxidation of graphite from the sedimentary country rocks (Fig. 9d). Most authors have argued that CO2 with negative δ13CCO2 values could formed by oxidation/hydrolysis of reduced carbon in sedimentary rocks during metamorphism (Oberthuer et al. 1996; Kontak and Kerrich 1997; Klein et al. 2008; Billstrom et al. 2009).

Nevertheless, some low δ13CCO2 values are also documented from deposits hosted by igneous rocks. One value at −24.1‰ is from fluid inclusions in quartz from the Proterozoic Chega Tudo gold deposit (Brazil), where Klein et al. (2008) argued that the low δ13CCO2 values reflect organic carbon from the carbonaceous schist hosting the porphyritic rocks. The two others values (−25.0‰ and −18.8‰) have been measured from fluid inclusions in scheelite from the Proterozoic intrusive-hosted Björkdal gold deposit, Sweden, (Billstrom et al. 2009), where they argued that these values indicate organic carbon derived from the greywacke-dominated sedimentary rocks that host the intrusion.

The most debated δ13CCO2 values are those between −12 and −1‰ (Fig. 10). Several authors argued that these signatures indicate a deep-seated carbon source, such as sub-continental mantle or an igneous carbon source (Rye and Rye 1974; Shelton et al. 1988; Zhang et al. 1989; Santosh et al. 1995; Oberthuer et al. 1996; Haeberlin 2002; So et al. 2002; Pandalai et al. 2003; Klein et al. 2007; Billstrom et al. 2009; Sun et al. 2009). Klein et al. (2007) concluded that the wide range of δ13CCO2 values (−10.7 to −3.9‰) do not allow to discriminate a particular source and could reflect leaching of carbon of the country rocks along the fluid pathway. Alternatively, some authors argued that these signatures could also be derived from prograde metamorphism of reduced carbon-bearing sediments in the middle crust (Zhang et al. 1989; Madu et al. 1990; Oberthuer et al. 1996).

5.3.3 Source of H2S

The sulfide minerals in orogenic gold deposits (Fig. 11) have δ34S mostly from ~0‰ to 10‰ with a significant number of data tailing to values as low as −30‰. Mantle H2S has a narrow range of compositions with δ34S between ~−1.5‰ and ~1.0‰ (Labidi et al. 2013) whereas magmatic H2S has δ34S values typically ranging between ~−3.7‰ and ~16‰ (Seal 2006). Ancient marine evaporites record the marine sulfate reservoir with positive δ34S values mainly between ~10 and ~35‰ (Claypool et al. 1980; Seal 2006). Finally, sedimentary pyrite defines a large range of δ34S with low values, as low as −50‰, and positive values up to ~16‰. This large range of values reflects variable sulfur isotope fractionation during bacterial and thermochemical sulfate reduction (BSR and TSR). Sulfur isotope fractionation during TSR yields sulfides with 11–43‰ lower than sulfate δ34S values at 500 °C and 100 °C respectively (Eldridge et al. 2016). During BSR, depletion of 34S is up to 72‰ depending of environmental conditions (Machel et al. 1995; Wortmann et al. 2001; Johnston et al. 2007). In orogenic gold deposits, the range in H2S sulfur isotope composition can result from mixing of different sulfur reservoirs or sulfur leaching during fluid/rock exchange along the fluid pathway, and changes in physico-chemical conditions (pressure, temperature, fO2, pH) of the fluid (Schwarcz and Rees 1985; Fedorowich et al. 1991; Thode et al. 1991; Couture and Pilote 1993; Neumayr et al. 2008). Commonly, low δ34S values (< 10‰) are attributed to BSR sources of reduced H2S such as from sedimentary rocks (Nie and Bjorlykke 1994; Oberthuer et al. 1996; Nie 1998; Bierlein et al. 2004; Goldfarb et al. 2004; Chen et al. 2008), whereas higher δ34S values (> 10‰) are attributed to TSR sources of H2S (Kontak and Smith 1989; Tornos et al. 1997) or igneous sulfur (Hattori and Cameron 1986).

Investigation of mass independent fractionation of sulfur (MIF-S; Δ33S) in hydrothermal systems has been used as a complementary tracer of the source of H2S. Whereas δ34S values are sensitive to physico-chemical processes which can modify its isotopic composition along the fluid pathway or at the deposition site, MIF-S is a chemically conservative tracer. By combining δ34S and Δ33S, several studies of Archean orogenic gold deposits from Yilgarn, Barberton and Abitibi have shown that H2S is partly derived from Archean sedimentary rocks (Helt et al. 2014; Agangi et al. 2016; Selvaraja et al. 2017; LaFlamme et al. 2018; Petrella et al. 2020). The low variance of Δ33S values in a deposit has been interpreted to indicate a homogeneous sedimentary H2S reservoir, such that δ34S variations in a deposit record variations of the oxidation state of the hydrothermal fluid (LaFlamme et al. 2018; Petrella et al. 2020) or redox reactions with country rocks, as suggested in Neumayr et al. (2008) for the St. Yves gold camp (Yilgarn). On the other hand, δ34S variations at the deposit or district scales can record the incorporation of H2S from the local country rocks as documented by Goldfarb et al. (1991); Goldfarb et al. (1997)

5.3.4 Nitrogen

Similar secular changes of the N isotope composition of mica and K-feldspar from orogenic gold veins and of sedimentary rocks suggest that hydrothermal fluids were dominantly derived from metamorphic dehydration of sedimentary rocks, consistent with the oxygen and hydrogen isotopic composition (Jia and Kerrich 1999, 2004; Jia et al. 2001, 2003), or that were equilibrated with sedimentary rocks along the fluid pathway (Kreuzer 2005). Nitrogen is incorporated into sedimentary rocks by biological activity and is thus mainly bound to organic matter. Organic matter displays a wide range of δ15N values (−4–15‰) (Minagawa and Wada 1986; Wu et al. 1997; Gu 2009). Kerogen in sedimentary rocks displays a more limited range of δ15N values (2–6‰) (Williams et al. 1995; Ader et al. 1998; Sephton et al. 2002), whereas organic N in marine sediments ranges between 0‰ and 10‰ (Peters et al. 1978; Compton et al. 1992; Williams et al. 1995). During diagenesis and metamorphism, the maturation of organic matter releases ammonium (NH4+) that substitutes for K+ in mica and K-feldspar. Metamorphic rocks display δ15N values between 2 and 17‰ (Haendel et al. 1986; Bebout and Fogel 1992; Jia and Kerrich 2000). Some studies argued that metamorphism decreases the N concentration, and increase δ15N values of the metamorphosed rocks (Haendel et al. 1986; Bebout and Fogel 1992), but the N isotope fractionation between source rock, hydrothermal fluid and newly formed minerals is small (Busigny et al. 2003; Jia et al. 2003; Jia and Kerrich 2004; Pitcairn et al. 2005; Kerrich et al. 2006).

5.3.5 Boron

The B isotope composition of tourmaline from orogenic gold deposits displays a wide range of values from −22‰ and 9‰ (Fig. 13). At temperatures between 250 and 550 °C, the tourmaline-fluid fractionation ranges from −4.5 and −1.6‰, (Meyer et al. 2008). Consequently, the large variation of 30‰ for δ11B values cannot be the result of variation in the temperature of formation of orogenic gold deposits. The B isotope composition of the marine seawater reservoir has values near 40‰. In contrast, δ11B values of others terrestrial reservoirs mostly overlap: continental crust (−6 to −14‰), I-type and S-type magmas (−7 to 3‰ and −7 to −15‰, respectively), MORB and mantle rocks (−8 to −6‰), marine sediments (−14 to −2‰), or marine carbonate (−12 to 26‰) (Fig. 13c) (Palmer and Swihart 1996; van Hinsberg et al. 2011; Lambert-Smith et al. 2016; Marschall et al. 2017; Trumbull and Slack 2018).

At the Paleoproterozoic Palokas gold deposit (Fennoscandian Shield), δ11B values in the range of −4.5—1.0‰ are interpreted to result from a single magmatic-hydrothermal event (Ranta et al. 2017). In the Paleozoic Passagem de Mariana gold deposit (Brazil), Trumbull et al. (2019) reported δ11B values between −11.5 and 7.1‰ for orogenic gold veins, and argued in favor to a crustal B source derived from the sedimentary rocks hosting the deposit. A similar hypothesis was discussed by Büttner et al. (2016) and Jiang et al. (2002) to explain the range of δ11B values (−18.4 to −8.9‰ and −21.4 to −17.8‰, respectively) in Proterozoic gold deposits from the Twangiza-Namoya gold belt (Democratic Republic of Congo) and in the Archean Mount Gibson deposit (Yilgarn Craton), respectively.

The contribution of at least two different sources of B is invoked to explain the wide range of δ11B recorded in some deposits. At Hira Buddini (India), Krienitz et al. (2008) documented tourmaline with δ11B values between −4 and 9‰, interpreted to result from mixing between a metamorphic fluid (δ11B ~ 0‰), generated by devolatilization of volcanic rocks, and a magmatic fluid (δ11B ~ 10‰) exsolved from I-type granitic intrusions. A similar interpretation was advanced by Molnár et al. (2016) to explain the wide range of δ11B values (−24.1 to −9.6‰) in the Archean Hattu schist belt (Finland), where the range in values was interpreted to reflect mixing between a heavy magmatic B with a lighter B reservoir from metamorphic devolatilization of sedimentary rocks.

Tourmaline from orogenic gold quartz-tourmaline veins from the Proterozoic Tapera Grande and Quartzito deposits (Brazil) shows δ11B values ranging from −15.7 to −5.0‰ (Garda et al. 2009). The wide range of value is interpreted to result from mixing B from multiple sources, derived from volcano-sedimentary rocks (higher δ11B), and from granitic or sedimentary rocks with lower δ11B values. Similarly, tourmaline from the Proterozoic Bhukia gold deposit (India) has δ11B values in the range of −12.6 to −9.2‰, that have been interpreted to originate from two fluid sources, a granite-derived hydrothermal and a metapelite-derived metamorphic fluids (Hazarika et al. 2019). Lambert-Smith et al. (2016) also argued that the wide range of δ11B values, from 3.5 to 19.8‰, in tourmaline from gold-rich veins in the Loulo mining district (Mali), resulted from mixing marine evaporite with a sedimentary-derived B. In the Val-d’Or district, Beaudoin et al. (2013) and Daver et al. (2020) concluded that tourmaline δ11B values, between −15.6 and −7.7‰, are a result of mixing of a low δ11B prograde metamorphic fluids released from sedimentary and/or volcanic rocks with a high δ11B seawater-derived pore fluid. Trumbull et al. (2020) show a bimodal distribution of the δ11B values for orogenic gold deposit tourmaline, consistent with two common B reservoirs, sourced from metamorphic/igneous rocks (low δ11B) or from seawater-altered volcano-sedimentary rocks (high δ11B).

In summary, B in orogenic gold deposits generally derived from multiple sources. It is generally accepted that B in orogenic deposits is derived from the devolatilization of volcano-sedimentary and/or sedimentary sequences during metamorphism (Jiang et al. 2002; Krienitz et al. 2008; Garda et al. 2009; Beaudoin et al. 2013; Büttner et al. 2016; Lambert-Smith et al. 2016; Molnár et al. 2016; Hazarika et al. 2019; Trumbull et al. 2019; Daver et al. 2020). The wide range of B isotopic values suggest mixing with other B source(s) that remains matter of debate. B-rich seawater-derived fluid (Beaudoin et al. 2013; Lambert-Smith et al. 2016; Daver et al. 2020) has been proposed as possible source, consistent with the upper crustal fluid identified using oxygen and hydrogen isotopes. B-rich magmatic-derived fluid are also suggested in rare cases (Krienitz et al. 2008; Molnár et al. 2016) but are unlikely to be a common boron reservoir.

5.3.6 Silicon

The silicon isotope composition of quartz is poorly documented in orogenic gold deposits. Only 30 δ30Si values are compiled (Fig. 14a,b), from the Hemlo (Superior Craton) and the Sawaya’erdun (Tianshan Orogen, China) deposits (Ding et al. 1996; Liu et al. 2007). Natural variation of the δ30Si values between various terrestrial reservoirs is small (Fig. 14c). Mantle peridotite, granulitic lower continental crust, igneous continental rocks, oceanic crust (MORB + OIB), and sedimentary rocks display overlapping ranges of δ30Si values between −0.8 and 0‰ (Poitrasson 2017). Only chert-BIFs, and seawater stand out with values between −3.6–5.0‰ and 0.6–4.4‰, respectively (Poitrasson 2017). Silicon does not fractionate significantly during fluid-rock interaction  at temperatures above 50 °C (Douthitt 1982; Ding et al. 1988; Geilert et al. 2014; Poitrasson 2017), and thus does not provide information on temperature and processes during quartz precipitation. For orogenic gold systems, δ30Si is mainly used to discuss the source of silicon by comparing the δ30Si values of the veins with those of the local country-rocks. Hence, (Liu et al. 2007) argued that the silica (δ30Si = −0.5–0.5‰) associated with the formation of the Sawaya’erdun orogenic gold deposit was derived from sedimentary (δ30Si = −0.5–1.9‰) and volcanic (δ30Si = 0.1‰) country rocks.

6 Conclusions

The stable isotope geochemistry of veins minerals has been widely used to constrain various parameters related to fluid systems involved in the formation of orogenic gold deposits. In this review, we compare and integrate multi-isotopes data from numerous deposits worldwide of various ages and locations to identify global trends shared by most of orogenic gold deposits.

6.1 Temperature

Temperature estimates using quartz-silicate, quartz-carbonate and sulfide-sulfide mineral-pairs display consistent temperature ranges of 360 ± 76 °C (σ = 1, n = 332) with no secular changes, which are in agreement with temperatures expected for deposits formed at temperatures typical for the greenschist to lower amphibolite facies of the country rocks. Quartz-silicate pairs are mostly in isotopic equilibrium, suggesting that retrograde isotopic exchange between minerals and fluids did not disturb the bulk isotopic composition. In contrast, a significant proportion of quartz-carbonate pairs show isotopic disequilibrium. This could result from quartz and carbonate forming at different times, but later oxygen isotopic exchange of carbonate with a low-temperature fluid is more consistent with the paragenetic sequence of orogenic gold deposits. Sulfide-sulfide mineral-pairs are mostly in isotopic equilibrium.

6.2 Fluid Reservoirs

The oxygen and hydrogen isotope composition of fluids indicate that fluid mixing is a process documented in many orogenic gold deposits. The mixing endmembers are a low T—high δD—low δ18O upper crustal and a high T—low δD—high δ18O deep-seated fluid reservoirs.

The nature of the upper crustal fluid is surficial water of various origins having experienced long lived exchange with country rocks. Meteoric water is rarely documented and is commonly related to late infiltration disconnected from the formation of the gold deposits. This is likely the case for most low δD values (< −120‰) documented in OH-bearing mineral and fluid inclusion, mainly from Mesozoic deposits.

The deep-seated fluid endmember is likely metamorphic in origin as attested by hydrogen and oxygen isotope compositions of water plotting mostly in the field for metamorphic water, even if a few data plot in the overlapping field for magmatic and metamorphic waters. The δ11B values also point a metamorphic source for the fluid that mixed with B-rich seawater-derived fluid, consistent with O–H isotope. The δ15N values of orogenic gold vein minerals also favor a deep-seated fluid reservoir derived from metamorphic dehydration of sedimentary rocks. The contribution of magmatic water exsolved from magma is documented in some specific cases but is not an essential component for the formation of orogenic gold deposits.

Spatial variations of oxygen and hydrogen isotope compositions of deep-seated fluid between districts indicate variable proportions of igneous and sedimentary rocks in the crustal segment released water during prograde metamorphism. Such provinciality is documented for δ18O values of quartz between cratons and districts.

Finally, a significant number of high δD values (up to 20‰) for water result of (i) muscovite precipitation during wallrock alteration and (ii) successive phase-separation of water associated to cyclic seismic rupture of fault-veins that displays trends of enrichment in deuterium that fit well the data.

6.2.1 Carbon Isotopes

No secular variations of mineral isotope composition are identified for carbon isotopes (δ13C). The δ13C values in orogenic gold deposits suggest that CO2 can be derived from the oxidation of reduced carbon and/or from decarbonation reaction of ~0‰ inorganic sedimentary rocks during metamorphism, or derived from a deep-seated carbon source such as the mantle. Wide ranges of δ13C are commonly documented at the deposit scale that suggests a carbon contribution from multiple sources, or that fluid acquired its δ13C isotopic signature along the fluid pathway through exchange with country rocks.

6.2.2 Sulfur Isotope

δ34S values of sulfide minerals do not display secular variation or provinciality, but display wide ranges of values. Mixing of sulfur from different reservoirs and/or changes in physico-chemical conditions of the fluid at the site of deposition have been proposed to explain the δ34S variation in a deposit. Commonly, low δ34S values are attributed to BSR sulfur from sedimentary rocks, whereas positive δ34S values are attributed to TSR sulfur or igneous sulfur. Orogenic gold deposits hosted by sedimentary rocks contain sulfur leached from the local country rocks. At the deposit scale, δ34S values record changes in the oxidation state of the hydrothermal fluid related to redox reactions with country rocks along the fluid pathway or at the site of deposition.