Porphyry Deposits and Oxidized Magmas_2015

Ore Geology Reviews 65 (2015) 97–131

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Ore Geology Reviews journal homepage: www.elsevier.com/locate/oregeorev

Review

Porphyry deposits and oxidized magmas Weidong Sun a,⁎, Rui-fang Huang b,c, He Li a, Yong-bin Hu a,c, Chan-chan Zhang a,c, Sai-jun Sun a,c, Li-peng Zhang a,c, Xing Ding b, Cong-ying Li a, Robert E. Zartman a, Ming-xing Ling b a b c

CAS Key Laboratory of Mineralogy and Metallogeny, Guangzhou Institute of Geochemistry, Chinese Academy of Sciences, 511 Kehua Street, Wushan, Guangzhou 510640, China State Key Laboratory of Isotope Geochemistry, Guangzhou Institute of Geochemistry, Chinese Academy of Sciences, 511 Kehua Street, Wushan, Guangzhou 510640, China University of the Chinese Academy of Sciences, Beijing 100049, China

a r t i c l e

i n f o

Article history: Received 17 June 2014 Received in revised form 19 August 2014 Accepted 2 September 2014 Available online 16 September 2014 Keywords: Porphyry deposit Oxidized magmas Oxygen fugacity Adakite Slab melts Arc magmas Plate subduction Ridge subduction

⁎ Corresponding author. E-mail address: [email protected] (W. Sun).

http://dx.doi.org/10.1016/j.oregeorev.2014.09.004 0169-1368/© 2014 Elsevier B.V. All rights reserved.

a b s t r a c t Porphyry deposits supply most of the world's Cu and Mo resources. Over 90% of the porphyry deposits are found at convergent margins, especially above active subduction zones, with much fewer occurrences at postcollisional or other tectonic settings. Porphyry Cu–(Mo)–(Au) deposits are essentially magmatic–hydrothermal systems, which are generally initiated by injection of oxidized magmas saturated with metal-rich aqueous fluids, i.e., the parental magmas need to be water rich and oxidized with most of the sulfur appearing as sulfate in the magma. Sulfur is the most important geosolvent that controls the behavior of Cu and other chalcophile elements, due to high partition coefficients of chalcophile elements between sulfide and silicate melts. Small amount of residual sulfides can hold a large amount of Cu. Therefore, it is essential to eliminate residual sulfides to get high Cu contents in magmas for the formation of porphyry deposits. Sulfate (SO2− 4 ) is over 10 times more soluble than sulfide (S2−), and thus the solubility of sulfur depends strongly on sulfur speciation, which in turn depends on oxygen fugacities. The magic number of oxygen fugacity is log fO2 N FMQ + 2 (i.e., ΔFMQ + 2), where FMQ is the fayalite–magnetite–quartz oxygen buffer. Most of the sulfur in magmas is present as sulfate at oxygen fugacities higher than ΔFMQ + 2. Correspondingly, the solubility of sulfur increases from ~1000 ppm up to N 1 wt.%. Oxidation promotes the destruction of sulfides in the magma source, and thereby increases initial chalcophile element concentrations, forming sulfur-undersaturated magmas that can further assimilate sulfides during ascent. Copper, Mo and Au act as incompatible elements in sulfide undersaturated magmas, leading to high chalcophile element concentrations in evolved magmas. The final porphyry mineralization is controlled by sulfate reduction, which is usually initiated by magnetite crystallization, accompanied by decreasing pH and correspondingly increasing oxidation potential of sulfate. Hematite forms once sulfate reduction lowers the pH sufficiently, driving the oxidation potential of sulfate up to the hematite–magnetite oxygen fugacity (HM) buffer, which is ~ΔFMQ + 4. Given that ferrous iron is the most important reductant that is responsible for sulfate reduction during porphyry mineralization, the highest oxygen fugacity favorable for porphyry mineralization is the HM buffer. In addition to the oxidation of ferrous iron during the crystallization of magnetite and hematite, reducing wallrock may also contribute to sulfate reduction and mineralization. Nevertheless, porphyry deposits are usually mineralized in the whole upper portion of the pluton, whereas interactions with country rocks are generally restricted at the interface, therefore assimilation of reducing sediments is not likely to be a decisive controlling process. Degassing of oxidized gases has also been proposed as a major process that is responsible for sulfate reduction. Degassing, however, is not likely to be a main process in porphyry mineralization that occurs at 2–4 km depths in the upper crust. Sulfide formed during sulfate reduction is efficiently scavenged by aqueous fluids, which transports metals to shallower depths, i.e., the top of the porphyry and superjacent wallrock. According to traditional views, sulfide saturation and segregation during magma evolution is not favorable for the formation of porphyry Cu ± Au ± Mo deposits. This is the main difference between porphyry deposits and Ni–Cu sulfide deposits. Nevertheless, in places with thick sections of reducing sediments, e.g., the western North America, sulfide saturation and segregation may occur during evolution of the magma, forming Cu-rich cumulates at the base of plutons. These Cu-rich sulfides may evolve into porphyry mineralization or even control the ore-forming process. Their contribution depends heavily on subsequent oxidation, i.e., a major contribution can be expected only when the sulfide cumulates are oxidized to sulfate, liberating the chalcophile elements. Sulfate reduction and ferrous Fe oxidation form H+, which dramatically lowers the pH values of ore-forming fluids and causes pervasive alteration zones in porphyry Cu deposits. The amount of H+ released during mineralization and the alkali content in the porphyry together control the intensity of alterations. In principle, H2 and methane

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form during the final mineralization process of porphyry deposits, but are mostly oxidized by ferric Fe during subsequent processes. Some of the reduced gases, however, may survive the highly oxidizing environment to escape from the system, or even to get trapped in fluid inclusions. Therefore, small amount of reduced gases in fluid inclusions cannot argue against the oxidized feature of the magmas. Reduced magmas are not favorable for porphyry mineralization. Reduced porphyry deposits so far reported are just mineralization that has either been reduced in host rock away from the causative porphyry or through assimilation of reducing components during emplacement. © 2014 Elsevier B.V. All rights reserved.

Contents 1. 2.

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Brief introduction of major oxygen buffers . . . . . . . . . . . . . . . . . . . . 2.1. Fayalite–magnetite–quartz (FMQ) oxygen buffer . . . . . . . . . . . . . 2.2. Hematite–magnetite (HM) oxygen buffer . . . . . . . . . . . . . . . . . 2.3. Ni–NiO(NNO) oxygen buffer . . . . . . . . . . . . . . . . . . . . . . 2.4. Pyrite + pyrrhotite + magnetite (PPM) oxygen buffer . . . . . . . . . . . 3. The association of porphyry deposits with oxidized magmas . . . . . . . . . . . 3.1. Large porphyry deposits . . . . . . . . . . . . . . . . . . . . . . . . 3.1.1. Porphyry Cu and Au deposits . . . . . . . . . . . . . . . . . . 3.1.2. Porphyry Cu–Mo deposits . . . . . . . . . . . . . . . . . . . 3.1.3. Porphyry Mo deposits . . . . . . . . . . . . . . . . . . . . . 3.2. Linkage between oxidized magmas and porphyry deposits . . . . . . . . . 3.2.1. Sulfur oxidation, sulfide under saturation and residual sulfide . . . 3.2.2. Sulfate reduction . . . . . . . . . . . . . . . . . . . . . . . 3.2.3. Hematite–magnetite intergrowth . . . . . . . . . . . . . . . . 3.3. Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4. The association of porphyry deposits with reduced magmas . . . . . . . . . . . . 4.1. Reduced magmas . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1.1. Evidence for reduced magmas of Catface porphyry deposit . . . . 4.1.2. Evidence for reduced magma in the Baogutu porphyry deposit . . . 4.2. Source of copper . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.3. Formation of reduced porphyry deposits . . . . . . . . . . . . . . . . . 4.3.1. The formation of the Catface porphyry deposit . . . . . . . . . . 4.3.2. The formation of the Baogutu porphyry deposit . . . . . . . . . . 4.4. Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5. Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.1. The oxygen fugacities at convergent margins . . . . . . . . . . . . . . . 5.2. The difference between porphyry and epithermal in terms of oxygen fugacity 5.2.1. Magnetite crisis . . . . . . . . . . . . . . . . . . . . . . . . 5.2.2. Oxygen fugacity and open systems . . . . . . . . . . . . . . . 5.3. Adakite, slab melting, ridge subduction and porphyry Cu deposits . . . . . . 5.3.1. Adakite and porphyry Cu deposits . . . . . . . . . . . . . . . . 5.3.2. Ridge subduction and porphyry Cu deposits . . . . . . . . . . . 5.4. Alterations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6. Conclusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Acknowledgments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

1. Introduction Porphyry deposits are hosts to one of the most important economic mineral associations (Cooke et al., 2005; Halter et al., 2005; Heinrich et al., 2004; Mutschler et al., 2010; Sillitoe, 2010), accounting for ~80% Cu and ~95% Mo of the world's total reserves. It is also an important resource of Au, Ag, Zn, Sn and W. Most porphyry deposits are found above active subduction zones (Fig. 1) (e.g., Chiaradia, 2014; Chiaradia et al., 2012; Gonzalez-Partida et al., 2003; Hedenquist et al., 1998; Kesler, 1997; Lee, 2014; Richards, 1999, 2013; Sillitoe, 2010; Sun et al., 2011; Wilkinson, 2013), with a few occurrences at post-collisional or other tectonic settings (Sillitoe, 2010), e.g., porphyry Mo deposits in the eastern Qinling orogenic belt (Chen, 2013; Li et al., 2012a; N. Li et al., 2013) and, arguably porphyry Cu–Mo deposits in Gangdese belt on the south Tibetan Plateau (Hou et al., 2009; Qu et al., 2004; Xiao et al., 2012) and some porphyry Cu deposits in Iran (Calagari, 2003; Castillo, 2006; Haschke et al., 2010; Shafiei et al., 2009).

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The consensus is that most of the porphyry Cu ± Mo ± Au systems are initiated by injection of oxidized adakitic magma saturated with aqueous fluids that are S- and metal-rich, i.e., the parental magmas must be water rich and oxidized (e.g., Ballard et al., 2002; Burnham and Ohmoto, 1980; Garrido et al., 2002; Imai, 2002; Liang et al., 2006; Mungall, 2002; Sillitoe, 2010; Stern et al., 2007; Sun et al., 2013b). It is, however, still controversial as regards to: why high oxygen fugacity is favorable for the mineralization of porphyry deposits, how oxidized the magma could be, whether adakitic magma is essential for porphyry mineralization or whether the porphyry deposits can be associated with normal arc rocks (Fig. 2), and why the pure porphyry Mo deposits are also closely associated with highly oxidized magmas. Copper, Au and Mo are chalcophile elements, which are strongly controlled by the behavior and speciation of sulfur. Therefore, the less the quantity of residual sulfide, the higher the initial Cu contents in primary magmas (Fig. 3) (Lee et al., 2012; Sun et al., 2004a, 2013b). Experiments show that sulfate is much more soluble than sulfide in magmas

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Fig. 1. Worldwide distribution of porphyry Cu deposits. Note, most of the porphyry deposits are distributed along convergent margins. Porphyry Mo deposits are not shown. Modified after Sun et al. (2013b). Data sources: Mutschler et al. (2010).

(Beermann et al., 2011; Jugo, 2009). Therefore, considerably more sulfur is removed in the form of sulfate under higher oxygen fugacity (Lee et al., 2012; Liang et al., 2009; Sun et al., 2013b). Meanwhile, sulfide is kept undersaturated during the evolution of oxidized magmas, such that no sulfide segregation is expected (Sun et al., 2012a; Sun et al., 2013b). Therefore, Cu, Mo and Au act as moderately incompatible elements (McDonough and Sun, 1995; Sun et al., 2003a,b,c), which become enriched during the early stage of magma evolution, and suddenly go into magmatic fluids during magnetite crystallization (Sun et al., 2004a, 2013a). It has been proposed that oxygen fugacity of log fO2 N FMQ + 2 (i.e., ΔFMQ + 2, where FMQ is the fayalite–magnetite– quartz oxygen buffer) is the magic number for porphyry mineralization (e.g., Mungall, 2002; Sun et al., 2013b). Sulfate is the dominant species at log fO2 N FMQ + 2, which is much more soluble than sulfide in magmas (Jugo, 2009). Therefore, residual sulfides are more efficiently destroyed at oxygen fugacities higher than FMQ + 2, releasing their chalcophile elements (Sun et al., 2013b). Others have argued that SO2 is the main sulfur species dissolved in porphyry magmas (Richards, 2014; Smith et al., 2012), not sulfate nor sulfide. It was further argued that variations in oxidation state over typical ranges for arc magmas (ΔFMQ = 0 to +2) have no major effects on the potential of magmas to form porphyry Cu–(Mo) deposits during plate subduction (Richards, 2011b). The implication is that ΔFMQ + 2 is not of importance for porphyry Cu mineralization. Instead, water is the most important factor that controls porphyry mineralization (Richards, 2011a). The question then becomes that most arc plutons are water saturated and highly oxidized, why do only a very small fraction of special magmas (mostly adakitic characterized by high Sr/Y and high Sr) in arc settings form porphyry deposits. For example, there are essentially no porphyry deposits in Japan (Fig. 1). Note that Japan arc is associated with older and presumably wetter subducting plate, such that arc rocks are presumably wetter than subduction related rocks along the eastern margin of the Pacific Ocean (Sun et al., 2012a,b, 2013a,b). A recent study has suggested that the source region for arc magmas is probably neither unusually oxidized nor enriched in economic elements of interest, such as Cu, as has been shown by studies of primitive arc magmas (Lee et al., 2010, 2012). This is seemingly consistent with the moderate incompatibility of Cu during mantle magmatism (Sun

et al., 2003a,b), which leads to Cu depletion in the mantle wedge. Also, the moderate mobility of Cu during plate subduction (J.L. Li et al., 2013) can compensate for the depletion caused by previous melting. The implication thus is that normal arc rocks, i.e., peridotitic melts, are not favorable for porphyry Cu mineralization. Still other authors have argued that, instead of high oxidation, sulfide saturation of the magma and consequent pre-enrichment through sulfide accumulation are the most important step for porphyry ore deposits (Chiaradia, 2014; Lee, 2014; Wilkinson, 2013). This is probably true along the western margin of the North American continent and other places where reduced sediments are well developed (see more in Section 4). However, this hypothesis raises several other questions. Why are porphyry deposits usually associated with oxidized magmas, which is a situation so different from Ni–Cu sulfide deposits that clearly experienced sulfide saturation? How does porphyry remain oxidized after collecting sulfide from the pre-enriched sulfide accumulates? Why are the grades of porphyry deposits much lower than sulfide saturated Ni–Cu sulfide deposits? Why is there no Ni in porphyry deposits? Moreover, some porphyry deposits are apparently associated with reduced magmas at ΔFMQ − 0.5 to − 3 (Cao et al., 2014; Rowins, 2000; Smith et al., 2012). Although reduced porphyry deposits are rare, with tonnages much smaller than oxidizing porphyry deposits (Cao et al., 2014), the mechanism needs to be clarified. This contribution focuses on the controlling factors of porphyry mineralization. Major controversies to be discussed here are: (1) Why do most of the porphyry deposits associated with oxidized magmas occur at convergent margins? (2) What is the most favorable oxygen fugacity range for porphyry mineralization? (3) What is the genetic connection between high oxygen fugacity magmas and porphyry deposits? (4) Why are some porphyry deposits associated with reducing magmas? (5) What are the connections between porphyry deposits and reduced magmas, if there is any? 2. Brief introduction of major oxygen buffers Oxygen fugacity (fO2) is an important geological parameter affecting the stability of minerals, the evolution of magmas, and ore-forming processes. It is an equivalent of the partial pressure of oxygen in a particular

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Fig. 2. Two different models for porphyry Cu ± Au ± Mo deposits. A. Porphyry deposits are formed in normal arc rocks (after Richards, 2011a). According to this model, even the formation of giant porphyry deposits is nothing special but optimization of normal ore-forming processes, controlled by distinct tectonic configurations, reactive host rocks, or focused fluid flow that have helped to enhance the overall process (Richards, 2013). B. Porphyry deposits are associated with slab melts (modified after Wilkinson, 2013), which have high initial Cu contents (Sun et al., 2011).

environment (atmosphere, magmas, rocks, etc.) corrected for the nonideal character of the gas. Oxygen fugacity is very important to the behaviors of elements, but it is not a very precise term because in some cases, oxidation–reduction reactions do not involve any oxygen and the oxygen fugacity changes with pressure and temperature (Sun et al., 2014a,b). “Redox state” is a better term to describe the relative proportions of an element among its different oxidation states. Nevertheless, oxygen fugacity is popularly used in Earth sciences. Oxygen fugacity is usually notated as variations relative to a certain oxygen buffer. Oxygen buffer refers to an assemblage of minerals or compounds that constrains oxygen fugacity as a function of temperature and

pressure (Fig. 4), i.e., the oxygen fugacity of equilibration at a fixed pressure is defined by one of the curves in oxygen fugacity versus temperature diagrams. The concept of oxygen fugacity and oxygen fugacity buffer have been well developed and widely used in high pressure experiments. Oxygen fugacity is of critical importance for the behaviors of elements. Therefore, it is crucial to know the oxygen fugacity of a geochemical process and to control oxygen fugacity during experiments. There are several ways to control oxygen fugacity of experiments in the laboratory. Gas-mixing techniques have been commonly used for controlling oxygen fugacity of experiments conducted at high temperature

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Fig. 3. Copper contents during partial melting of mantle peridotite under different oxygen fugacities modeled by Lee et al. (2012). (A) Variation of aggregated melt and residual mantle composition with degree of partial melting (F) at 2GPa, 1350 °C and fO2 at FMQ + 0, assuming initial S content of 200 ppm, i.e. 0.06 wt.% of sulfide. (B) Copper contents in primary melt as a function of F and log fO2 (ΔFMQ); same P–T conditions as in (A). (C) Cu content of aggregate liquids versus fO2 for different F. Gray field refers to Cu in primitive MORB and arc magmas (Lee et al., 2012).

Fig. 4. Comparison of log fO2–T relationships for the Ni–NiO, Co–CoO, FMQ, MnO–Mn3O4, and HM oxygen buffers at 1 bar. The log fO2 of Co–CoO, FMQ and HM are taken from Chou (1978). The log fO2 of Ni–NiO and MnO–Mn3O4 are taken from Huebner and Sato (1970).

(above 1300 K) and 1 atm pressure (Huebner, 1987). The gas mixtures mostly used are CO2–CO and CO2–H2. At higher pressure, oxygen fugacity can be controlled by the oxygen buffer technique (Eugster, 1957) and the Shaw membrane (Shaw, 1963), and measured with the hydrogen fugacity sensor technique (Chou, 1978). This oxygen buffer technique was first developed by Eugster in 1957 to prevent oxidation of iron during growth of hydrous ferrous silicates and then later to determine the stability of annite (Eugster, 1957; Eugster and Wones, 1962; Wones and Eugster, 1965). The oxygen buffers commonly used in high pressure and high temperature experiments are fayalite– magnetite–quartz (FMQ), Co–CoO, Ni–NiO (NNO), hematite–magnetite (HM), and MnO–Mn3O4. These buffers are also useful as notations in the measurement of the oxygen fugacity of natural rocks and magmas. A comparison of the log fO2–T relationship for these buffers is shown in Fig. 4. In experiments using the conventional double-capsule method, the mixture of buffer materials and water is loaded in the outer capsule and the sample is sealed in the inner capsule. Hydrogen gas forms during the reaction between buffer materials and water in the outer capsule, then diffuses through the inner capsule wall and equilibrates with the hydrogen in the inner capsule. Capsules with high hydrogen diffusivity are usually taken as the inner capsule (e.g., Pt), whereas that with low hydrogen diffusivity is used as the outer capsule (e.g., Ag). The

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prerequisite for using the double-capsule technique is that the buffer materials themselves do not diffuse into the inner capsule. Therefore, they cannot be used for system with haplogranitic melts and H2SO4 solutions because Co and Ni diffuse into the inner capsule, resulting in formation of nickel sulfide or cobalt sulfide (Keppler, 2010). The other way to regulate the oxygen fugacity of sample assemblages is to control the hydrogen fugacity using the Shaw membrane, which introduces variable amount of hydrogen gas into the system (Shaw, 1963). The advantage of this technique is that the hydrogen fugacity can be varied independently and continuously. Moreover, the Ag–AgCl–H2O–HCl acid buffer can be used as a fH2 sensor by loading the acid-buffer assemblage in a small Pt capsule and putting the capsule in a larger capsule containing the sample assemblages (Chou, 1978). In this case, the measured concentration of HCl in the sensor indicates the fH2 of the outer sample system. 2.1. Fayalite–magnetite–quartz (FMQ) oxygen buffer The FMQ oxygen buffer is widely used in studies of natural samples, as well as for experiments. Given that the oxygen fugacity of most igneous rocks plot within several log units of the FMQ buffer, the oxygen fugacities of natural samples are often noted as log unit variations from the FMQ buffer, i.e., in ΔFMQ units. For example, ΔFMQ + 2 means oxygen fugacity 2 log units higher than the values defined by the FMQ buffer. The FMQ equilibrium has been determined experimentally by many researchers, e.g., Chou (1978), Myers and Eugster (1983) and Eugster and Wones (1962). The oxygen fugacity is controlled through reaction (1): 3Fe2 SiO4 þ O2 ¼ 3SiO2 þ 2Fe3 O4 :

ð1Þ

As shown in Fig. 5, the pressure dependence of log fO2 for this buffer is not significant, but log fH2 changes greatly from 1 bar to 10 kbar. 2.2. Hematite–magnetite (HM) oxygen buffer The HM (hematite + magnetite) equilibrium has been determined by several researchers (e.g., Chou, 1978; Eugster and Wones, 1962; Hemingway, 1990; Myers and Eugster, 1983). Oxygen fugacity is controlled through reaction (2) or (3): 4Fe3 O4 þ O2 ¼ 6Fe2 O3

ð2Þ

2Fe3 O4 þ H2 O ¼ 3Fe2 O3 þ H2 :

ð3Þ

Eq. (3) better describes the system with water present. As shown in Fig. 6, the pressure dependence of log fO2 is quite small, while log fH2 increases rapidly at low pressures but less at higher pressures. For high pressure experiments, the HM buffer is more oxidized than the intrinsic oxygen fugacity of the hydrothermal vessel. Thus, hematite was consumed quite fast, e.g., in some experiments around 3.73 mg Fe2O3 was consumed in the buffer assemblage every hour under steady-state conditions, and 100 mg of Fe2O3 lasted only 26.8 h (Chou, 1986). Therefore, the HM buffer is not suitable for experiments conducted in hydrothermal vessels with water as the pressure medium. In natural systems, the oxygen fugacity of the HM buffer is usually much higher than that of magmas. Most of the porphyry deposits, however, reach the HM buffer value during mineralization.

Fig. 5. Pressure dependence of log fO2 and log fH2 of FMQ oxygen buffer. The log fO2 of Ni– NiO buffer is taken from Chou (1978). Method for calibrating log fH2 of FMQ buffer is the same as that of the HM buffer.

Ni þ H2 O ¼ NiO þ H2 :

ð5Þ

As shown in Fig. 7, the pressure dependence of log fO2 is not significant, but log fH2 increases rapidly at low pressures but less at higher pressures. The Ni–NiO buffer is commonly used in experiments, bringing with it distinct advantages. First, the oxygen fugacity of Ni–NiO is quite close to the intrinsic oxygen fugacity of the hydrothermal vessel (NiNiO + 0.5). Thus, even quite small amounts of Ni and NiO powder could last for a long time during experiments, e.g., around 200 mg of Ni and NiO is enough for around a 10-day experiment. Moreover, oxygen fugacity controlled by the Ni–NiO buffer is quite important for the oxidation states of many elements, e.g., sulfur in the fluids is mostly as H2S at the oxygen fugacity of NNO, but SO2 and SO3 appear when the oxygen fugacity increases by 0.5 log unit (Binder and Keppler, 2011). 2.4. Pyrite + pyrrhotite + magnetite (PPM) oxygen buffer The PPM (pyrite + pyrrhotite + magnetite) equilibrium controls f S2 and fO2 through reactions (6)–(8):

2.3. Ni–NiO(NNO) oxygen buffer

2Fe1−x S þ ð1−2xÞS2 ¼ 2ð1−xÞFeS2

ð6Þ

The Ni–NiO oxygen buffer has been experimentally determined by Huebner and Sato (1970) and O'Neil and Pownceby (1993). The oxygen fugacity is controlled through reactions (4) and (5):

6Fe1−x S þ 4O2 ¼ 2ð1−xÞFe3 O4 þ 3S2

ð7Þ

Ni þ 1=2O2 ¼ NiO

3FeS2 þ 2O2 ¼ Fe3 O4 þ 3S2 :

ð8Þ

ð4Þ

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Fig. 6. Pressure dependence of log fO2 and log fH2 for the HM buffer. The log fO2 of HM buffer is taken from Chou (1978). The log fH2 is calculated using log K = log f H2 + 1/2log fO2, where K is the equilibrium constant of H2O = H2 + 1/2O2 determined by Supcrt92 with the database DPRONS2003.

This buffer has been used in hydrothermal experiments, e.g., Spry and Scott (1986) and Crerar et al. (1978). The fO2 of the PPM buffer has been determined by Kishima (1989) and Shi (1992), and their results are in general agreement. Fig. 8. shows a comparison of the fO2–T relationship of the PPM buffer with those of the NNO and HM oxygen buffers. The log fO2 of the PPM oxygen buffer is located between those of NNO and HM buffers, e.g., at 700 °C, the log fO2 of PPM oxygen buffer is around 1 log10 unit lower than that of the HM buffer, but 3.3 log10 unit higher than that of the NNO buffer. As shown in Fig. 8b, log fO2 decreases slightly with increasing pressure (Kishima, 1989; Shi, 1992). At high temperature, the assemblage PPM is replaced by magnetite and pyrrhotite (MPo) because pyrite is not stable (Sun et al., 2013a,b; Tomkins, 2010). Consequently, SO2 becomes the dominant species, and fH2, fH2O, and fH2S decrease abruptly. The breakdown temperature of pyrite increases with increasing pressures, e.g., at 2 kbar, the breakdown temperature is 745 °C, and increases to around 800 °C to 10 kbar (Shi, 1992).

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Fig. 7. Pressure dependence of log fO2 and log fH2 (Ni–NiO oxygen buffer). The logfO2 of Ni–NiO buffer is taken from Huebner and Sato (1970). Method for calibrating log f H2 of Ni–NiO buffer is the same as that of HM buffer.

almost always attributable to the injection of an oxidized adakitic magma that is saturated with S- and metal-rich aqueous fluids, i.e., the parental magmas are water rich and oxidized (Ballard et al., 2002; Mungall, 2002; Sillitoe, 2010). Several small “reducing porphyry deposits” constitute the rare exceptions (Cao et al., 2014; Rowins, 2000; Sillitoe, 1999) (see detailed discussion in Section 4 below). Intergrowths of magnetite and hematite indicate that oxygen fugacities of porphyry deposits often reach the values defined by the HM buffer (Fig. 10). In this section, we review the oxygen fugacities of a number of famous deposits and then discuss the genetic connections between oxidized magmas and porphyry deposits. The main questions discussed in this section include that: whether oxygen fugacity is a controlling factor that dictates the distribution of porphyry deposits. What is the genetic link between high oxygen fugacity magmas and porphyry deposits? What is the most favorable oxygen fugacity range for porphyry mineralization? And how do oxygen fugacities change during porphyry mineralization? 3.1. Large porphyry deposits

3. The association of porphyry deposits with oxidized magmas It has long been proposed that most porphyry deposits are closely associated with oxidized magmas (Burnham and Ohmoto, 1980; Candela, 1992; Hedenquist and Lowenstern, 1994; Liang et al., 2006; Liang et al., 2009; Mungall, 2002; Sillitoe, 2010; Sun et al., 2012a, 2013b), also known as the magnetite-series magmas (Ishihara and Terashima, 1989) (Fig. 9). These porphyry Cu–(Mo)–(Au) deposits are

The close association between highly oxidized magmas and porphyry deposits may best be illustrated by a FeO versus fO2 diagram (Fig. 9). The highly oxidized nature of porphyry magmas have been reported in essentially all types of porphyries (Mungall, 2002; Sillitoe, 2010; Sun et al., 2013b; Vila et al., 1991), including porphyry Cu and Cu–Au deposits (Sillitoe, 2010), porphyry Au deposits (Vila and Sillitoe, 1991; Vila et al., 1991), to porphyry Cu–Mo deposits (Cuadra and Rojas, 2001; Lynch and Ortega, 1997; Stern et al., 2007) and pure porphyry

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Ohno, 2005; Imai et al., 1993; Sun et al., 2013b), and from active subduction zones to post-collisional zones (Hou et al., 2004, 2007b, 2009; Li et al., 2012a; Xiao et al., 2012). One common indicator of the high oxygen fugacity of porphyry deposits is sulfate, i.e., both magmatic and hydrothermal sulfate, e.g., hydrothermal anhydrite and magmatic anhydrite and gypsum are abundant in essentially all large porphyry deposits (Cooke et al., 2011; Zhang, Ling et al., 2013; Halter et al., 2005; Imai et al., 2007; Kavalieris et al., 2011; Li et al., 2008; Liang et al., 2009; Stern et al., 2007; Vila et al., 1991). In addition to sulfate, hypogene hematite and specularite have also been reported in many porphyry deposits (Baker et al., 1997; Hedenquist et al., 1998; Imai, 2001; Imai et al., 2007; Li et al., 2008; Seedorff and Einaudi, 2004b; Sillitoe, 2010; Spry et al., 1996; Vila and Sillitoe, 1991; Vila et al., 1991). The hematite–magnetite (HM) oxygen buffer has been taken as the upper limit of oxygen fugacities that are favorable for porphyry mineralization (Sun et al., 2013b). Previous authors also have proposed that the oxygen fugacities of different porphyry deposits are slightly different, in the order of: porphyry Cu and porphyry Au deposits N porphyry Cu–Mo deposits N porphyry Mo deposits (Fig. 9) (Thompson et al., 1999).

Fig. 8. Comparison of the log fO2 of PPM (pyrite–pyrrhotite–magnetite) buffer with those of Ni–NiO and hematite–magnetite buffer (A) and pressure dependence of the logfO2 of PPM buffer. The log fO2 of PPM buffer is calibrated from Shi (1992).

Mo deposits (Seedorff and Einaudi, 2004a,b). They are found in different tectonic settings, ranging from continental arc (Sillitoe, 2010; Stern et al., 2007, 2011) to island arc (Hedenquist et al., 1998; Imai and

Fig. 9. Schematic plot of Fe content in magmas versus oxidation state (fO2) for calc-alkaline to alkaline magmas associated with porphyry Cu, Cu–Mo and Mo deposits and W, Sn deposits. The approximate boundary between magnetite- and ilmenite-series magmas (Ishihara, 1977) is also shown. Nevertheless, the HM buffer is the upper limit of oxygen fugacity favorable for porphyry deposits. Modified after Thompson et al. (1999).

3.1.1. Porphyry Cu and Au deposits Porphyry Cu and Au deposits almost do have systematically higher oxygen fugacities and FeO contents compared to other porphyry deposits (i.e., Cu–Mo and Mo). The best examples are the Cenozoic porphyry Cu and Au deposits occurring in the southwestern Pacific islands, and to a less extent, the Paleozoic porphyry Cu–Au deposits in the Central Asian Orogenic Belt. From this one could surmise that porphyry Cu–Au deposits are associated with island arcs, whereas porphyry Cu–Mo deposits are associated with continental arcs. For example, Cenozoic porphyry deposits located in island arcs in the southwestern Pacific are all Cu–Au deposits, e.g., Grasberg and Batu Hijau in Indonesia; Panguna and Ok Tedi in Papua New Guinea; Lepanto–Far South East, Tampakan, Atlas and Sipilay in the Philippines (Cooke et al., 2005). Some researchers have even argued that Cu comes from the mantle, whereas Mo comes from the continental crust (Mao et al., 2011). This is not the case, however, although porphyry Cu (±Au and ±Mo) deposits indeed do have higher εNd isotopic values (Hou et al., 2007a), implying less contribution from the continental crust. This observation, however, does not necessarily support a mantle origin for Cu, either, because both the depleted mantle and the oceanic crust have high εNd values, whereas enriched mantle and the continental crust have low εNd values. As argued below (Section 3.2.1), a strong case can be made that most of the Cu comes from subducted oceanic crust. The distribution of porphyry deposits does not support the premise that the Cu comes from the mantle. Moreover, porphyry Cu–Au deposits are not always formed in island arc environments, either. For example, there are also many porphyry Cu–Au deposits (without economic levels of Mo) in the western American continents, including some of the world's top 25 largest porphyry deposits, e.g., Cerro Casale in Chile; Minas Conga in Peru; Bajo de la Alumbrera in Argentina; Pebble Copper in the USA; and Prosperity in Canada. Here we give brief introductions to three famous porphyry Cu–Au deposits. 3.1.1.1. Grasberg, Indonesia. Grasberg in Irian Jaya, Indonesia is one of the largest high-grade hypogene porphyry Cu–Au deposit (38.32 Mt Cu @ 1.12% and 3662 t Au @ 1.07 g/t) in the world, ranking no. 1 among porphyry Au deposits and among the top 10 porphyry Cu deposits (Cooke et al., 2005). It was formed ~3 Ma ago. The presence of anhydrite and hematite indicates very high oxygen fugacities (Cooke et al., 2005; Mathur et al., 2000), near the HM buffer. This deposit is related to the closure of a small backarc basin between Indonesia and the South China Sea. As is characteristic of Cenozoic porphyry deposits in southwestern Pacific, it contains no Mo.

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Fig. 10. Images of magnetite–hematite intergrowths from major porphyry deposits in China, indicating that the oxygen fugacity reached the HM buffer (Zhang, Ling et al., 2013). A. Dexing (Zhang, Ling et al., 2013), B. Xiongcun, C. Duobuza, D. Yulong (Sun et al., 2013a,b), E. Qulong, F. Zijinshan (Sun et al., 2013a,b).

3.1.1.2 . Far Southeast–Lepanto, Philippines. The Far Southeast–Lepanto porphyry and epithermal Cu–Au deposits, Philippines, are parts of a large epithermal–porphyry Cu–Au body with a total of 5.48 Mt Cu @ 0.8 wt.% and 973 t Au @ 1.42 g/t (Table 1; Cooke et al., 2005; Hedenquist et al., 1998). Sulfates, including anhydrite and alunite (likely related to the epithermal mineralization) are abundant in the mineralization system, and hematite formed during chlorite alteration, indicating high oxygen fugacity up to the HM buffer (Hedenquist et al., 1998). 3.1.1.3. Santo Tomas II (Philex), Philippines. The Santo Tomas II (Philex) porphyry deposit, Philippines, has total reserves of 1.2 Mt Cu @ 0.33 wt.% and 233 t Au @ 0.64 g/t (Table 1). It is also associated with a highly oxidized magma, as indicated by anhydrite veinlets and a

magnetite–titanohematite assemblage, which indicates oxygen fugacities near the HM buffer at nearly magmatic temperature (Imai, 2001). In summary, essentially, most if not all, porphyry Cu–Au deposits are highly oxidized (Figs. 9, 10). In addition to those mentioned above, hematite flakes in quartz veinlets have been reported in the Waisoi porphyry Cu deposit (Namosi district), Viti Levu, Fiji (Imai et al., 2007). The Tongshankou porphyry skarn Cu–Au deposit, in the Lower Yangtze River belt, central eastern China, also has primary hematite next to sulfides (Li et al., 2008). The euhedral characteristics of the hematite imply a hydrothermal origin. Porphyry Au deposits have oxygen fugacities similar to, if not higher than, porphyry Cu deposits. For example, Marte, a large porphyry Au

Table 1 Age, grade and tonnages of major porphyry deposits discussed in the text. Deposit

Cu–Au

Cu–Mo

Grasberg Lepanto-Far South East Santo Tomas II El Teniente Chuquicamata Dexing

Qulong Yulong Sar-Cheshmeh Batu Hijau Panguna Ok Tedi Tampakan Atlas Mo Henderson Reduced Catface Reduced Baogutu

Province

Ref

Age

Tonnage Au grade

Au

Cu grade

(Ma)

(Mt)

(t)

(wt.%) (Mt)

3662 973

1.12 0.8

233 437 301 19

0.33 0.92 0.71 0.45

324 572 799 446 336 331

0.5 0.84 1.20 0.44 0.46 0.64 0.55 0.50

Irian Jaya N. Luzon

Cooke et al. (2005), Mutschler et al. (2010) Cooke et al. (2005), Mutschler et al. (2010)

4–3 1.5–1.2

Philippines Central Chile Nothern Chile China

Mutschler et al. (2010) Cooke et al. (2005), Mutschler et al. (2010) Cooke et al. (2005), Mutschler et al. (2010) Mutschler et al. (2010), Zhang, Ling et al. (2013), Zhang, Sun et al. (2013) Xiao et al. (2012) Mutschler et al. (2010), Liang et al. (2006) Cooke et al. (2005), Mutschler et al. (2010) Mutschler et al. (2010) Cooke et al. (2005) Mutschler et al. (2010) Cooke et al. (2005) Cooke et al. (2005) Mutschler et al. (2010) Mutschler et al. (2010) Cao et al. (2014)

1 7.1–4.6 33.6 170–148

China China Iran Indonesia Bougainville PNG Philippines Philippines United States Canada China

14 Paleogene 12 5 3.5 1.2–1.1 3.3–2.2 61 30–27 Middle Eocene

3409 685 364 11,845 15,052 1500

(g/t) 1.07 1.42 0.64 0.0035 0.04 0.18

850 1200 0.27 1644 0.35 1415 0.57 700 0.64 1400 0.24 1380 0.24 727 308 0.1 14

0.37 0.28 0.63

Cu

Mo grade

Mo

(wt.%) (Mt)

38.3 5.48 1.2 109 106 8.4 10.4 7.14 14.4 7.26 6.51 4.48 7.70 6.90

0.02 0.024 0.01

2.50 1.81 0.29

0.03 0.03 0.03

0.5 0.15 0.36

0.17 1.14 0.01 0.011 0.018

1.24 0.02

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deposit located in the Maricunga belt of the Andean Cordillera, northern Chile, contains several percents of hematite, magnetite, anhydrite and gypsum (Vila et al., 1991), clearly reaching the HM buffer (Figs. 10, 11). It has also been argued that iron oxides–copper–gold (known as IOCG) deposits are the low S version of porphyry Cu–Au deposits, controlled by secular changes in oceanic sulfate content and the geothermal gradients at the end of the Precambrian (Richards and Mumin, 2013). This is still elusive. Nevertheless, IOCG deposits are all highly oxidized, up to the HM buffer.

3.1.2. Porphyry Cu–Mo deposits In general, porphyry Cu–Mo deposits have FeO contents systematically lower than porphyry Cu, but higher than pure porphyry Mo deposits, with oxygen fugacities also falling between them (Fig. 8). Many porphyry Cu–Mo deposits have high sulfate contents, with much less magnetite. No hypogene hematite has yet been reported in some of the supergiant porphyry Cu–Mo deposits, partially due to severe supergene oxidation, which makes it difficult to identify primary hematite. Porphyry Cu–Mo deposits have Cu and Au grades comparable to those of porphyry Cu ± Au deposits, with low Mo grades, ranging from 0.01 to 0.03 wt.% (Table 1) (Cooke et al., 2005). Some of the supergiant deposits may have large Mo reserves of several million tons (Cooke et al., 2005). The Cu–Au and Mo mineralizations that occur in the same ore bodies usually do not happen at the same time, indicating different sources and/or different mineralization processes. Consistently, the Mo/Cu ratios of these deposits are generally more than 20 times higher than the primitive mantle ratio value (McDonough and Sun, 1995), suggesting that Mo has been added to the porphyry deposits from different sources. One of the most important geologic processes that enriches Mo is the oxidation–reduction cycle operating during chemical weathering at the Earth's surface, i.e., Mo is mobilized during oxidation and then enriched in organic-rich sediments due to reduction (Li et al., 2012a,b). The

Fig. 11. Stability domains of the trisulfur ion S− 3 , sulfate, and sulfide in an aqueous solution, as a function of oxygen fugacity (log10 fO2) and acidity (pH = −log10 mH+, in mol per kg) at 350 °C and 0.5 GPa illustrating sulfate reduction (modified after Pokrovski and Dubrovinsky, 2011; Sun et al., 2013a,b). Also shown is the oxygen fugacity of the major mineral buffers (HM, thick horizontal orange line; NNO and FMQ, horizontal dashed lines) and the neutrality point of pure water (the vertical dashed line). The orange field between HM and ΔFMQ +2 is the optimum condition for porphyry Cu, Au, Mo mineralization (Sun et al., 2013b). Lines E1, E2, E3 and E4, show trajectories for sulfur reduction (Sun et al., 2013b). Stability domains of trisulfur ion S− 3 at total dissolved sulfur concentrations of 1 wt.% and 0.1 wt.% are shown.

involvement of such reducing agents as organic matter seemingly explains the system's lower oxygen fugacity. The most famous porphyry Cu–Mo deposits are the Cenozoic ones located along the eastern Pacific margin. 3.1.2.1. El Teniente. The supergiant El Teniente deposit in Chile is the largest porphyry Cu–Mo–Au deposit in the world, with total reserves of ~ 109 Mt Cu @0.92%, 2.5 Mt Mo @ 0.02% and 437 t Au @0.035 g/t (Table 1; Cooke et al., 2005; Mutschler et al., 2010). The deposit spatially corresponds to the Juan Fernandez Ridge (Cooke et al., 2005), and thus has been attributed to slab melting during ridge subduction (Sun et al., 2010). Others argued that ridge subduction is not responsible for making the supergiant porphyry deposit based on the inferred migration history of the arc (Kay et al., 2005). El Teniente is a nested porphyry system that, according to different authors, was active for ~ 1 Ma (Baker et al., 2011; Cannell et al., 2003) up to more than 7 Ma (14.2–6.5 Ma) (Barra, 2011; Stern et al., 2011; Vry et al., 2010). High Fe2O3/FeO values of 1 to 3 (Garrido et al., 2002), quartz anhydrite veins and anhydritecemented breccias as well as gypsum clearly indicate high oxygen fugacities (Klemm et al., 2007; Vry et al., 2010). Although it is not spelled out, the high Fe2O3/FeO ratios of 1 to 3 (Garrido et al., 2002) indicate abundant hematite, reaching the HM buffer. Hydrothermal rutile also indicates high oxygen fugacity (NNO + 1.3) (Rabbia et al., 2009), which is however much lower than the HM buffer, though. Nevertheless, the oxygen fugacity undoubtedly fluctuated during porphyry mineralization (Liang et al., 2009; Sun et al., 2013b), and the hydrothermal rutile only records the oxygen fugacities under which it crystallized. Most of the Cu in El Teniente was emplaced during the late magmatic stage (Cannell et al., 2005). A small sulfur- (N 3 wt.%) and copper-rich (N0.5 wt.%) fine-grained igneous rock known as “Porphyry A” stock (b1 km3, 6.09 ± 0.18 Ma) (Stern et al., 2011) has abundant igneous anhydrite, with varied textures ranging from interstitial to poikilitic and corresponding modal abundances from 10 to 20%, respectively. These igneous anhydrite grains with planar crystal boundaries, occur along with fresh and unaltered biotite, feldspars, quartz, and Fe-oxides (Stern et al., 2007). Because the anhydrite-rich stock is isotopically similar to all the other igneous rocks of the late Miocene deposit (Baker et al., 2011; Stern et al., 2007) formed at a time of regional compressive deformation, it has been proposed that an oxidized parent magma in the large productive magma chamber was undergoing igneous fractionation at that time. During a period of recharge by mantle-derived mafic magmas into the base of the chamber and associated volatile transfer and concentration near its roof, the opportunity arose to produce the Cu- and S-rich magmas that formed the anhydrite-bearing intrusive rocks (Stern et al., 2007). 3.1.2.2. Chuquicamata. Chuquicamata is another supergiant porphyry deposit in Chile, with a total reserve of ~ 106 Mt Cu @ 0.71%, 1.81 Mt Mo @ 0.024% and 301 t Au @ 0.04 g/t (Table 1. Cooke et al., 2005; Mutschler et al., 2010). Similar to El Teniente, the formation of the Chuquicamata porphyry Cu deposits also lasted for several million years, between 36 and 31 Ma (Ballard et al., 2001; Ossandon et al., 2001). It is closely associated with highly oxidized adakite (Oyarzun et al., 2001, 2002). Because quantitative oxygen barometers based on Fe–Ti oxides are prone to resetting, and primary whole rock Fe(III)/ Fe(II) ratios and anhydrite, if originally present, are unlikely to have survived the hydrothermal alteration and surficial weathering of this deposit, the zircon Ce4+/Ce3+ ratio has been used to indicate its high oxygen fugacity (Fig. 12) (Ballard et al., 2002). 3.1.2.3. Dexing. The Dexing porphyry Cu deposit is located in southeastern China, with total reserves of 8.4 Mt Cu @ 0.45%, 0.29 Mt Mo @ 0.01% and 19 t Au @ 0.18 g/t (Table 1. Zhang, Ling et al., 2013; Zhu et al., 1983). The deposit consists of a cluster of three porphyries, with Tongchang as the largest in the center, Fujiawu the second in the southeast and Zhushahong, a small deposit in the northwest (Li et al., 2007; X.F. Li

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Eurasian continents, Qulong is taken as a typical post-collisional porphyry deposit. Nevertheless, the Indian plate is still subducting northward. It is not clear whether the portion of the subducting slab underneath Qulong was continental or oceanic. Intergrowths of hematite and magnetite (Fig. 10) indicate the oxygen fugacity of Qulong reached the HM buffer. Adakites found in the Qulong deposit are mixtures of continental and slab melts (Sun et al., 2012a).

Fig. 12. Ce4+/Ce3+ and (Eu/Eu*)N ratios of zircon grains for Dexing and Shapinggou porphyry deposits as well as ore-bearing and ore-barren porphyries from Chile. High Ce4+/Ce3+ signifies high oxygen fugacity. Data of ore-bearing and ore-barren samples in Chile from Ballard et al. (2002), Dexing from Zhang, Ling et al. (2013) and Shapinggou from Zhang, Li et al. (2013).

et al., 2013; Zhou et al., 2013). The ore minerals comprise pyrite, chalcopyrite, molybdenite, minor tetrahedrite, bornite and chalcocite. Gangue minerals include quartz, muscovite, chlorite, calcite, minor epidote and anhydrite. The porphyry is an adakite formed at ~170 Ma (Zhang, Ling et al., 2013; Wang et al., 2006a,b). Another poly-metal deposit, Yinshan, formed roughly at the same time and may be paragenetically related (Wang et al., 2013). The Dexing porphyry deposit is closely associated with highly oxidized magmas (Zhang, Ling et al., 2013; Li and Sasaki, 2007), e.g., it is famous for abundant specular hematite, which mostly formed during late stage hydrothermal alteration (Fig. 13). Intergrowths of magnetite and hematite indicate its oxygen fugacity reached the magnetite–hematite buffer (Fig. 11). Hematite is also found in fluid inclusions in Dexing (Liu et al., 2011).

3.1.2.5. Yulong. The Yulong porphyry Cu–Au deposit belt is distributed along the northwestern extension of the Red River–Ailao Shan fault system, at the eastern margin of the Tibetan Plateau (Hou et al., 2007b; Jiang et al., 2006; Liang et al., 2006), and covers an area of ~ 300 km long and ~ 20 km wide. Five major porphyries, which contain most of the Cu reserves so far discovered in the belt, are located in a narrow, elongated domain of approximately 50 km long and 10 km wide and are closely associated with Cenozoic high potassium intrusive rocks (Liang et al., 2006, 2007). The Yulong porphyry is the largest one in the Yulong copper deposit belt, with reserves of N7.14 Mt of Cu @ 0.84% and 0.15 Mt of Mo @ 0.028% (Table 1; Liang et al., 2006; Mutschler et al., 2010). All five known porphyry deposits all together contain a total of more than 8 million tons of Cu and Mo reserves. Zircon U–Pb dating shows that the formation ages of deposits within the Yulong ore belt range from 41.2 to 36.9 Ma, extending over a period of similar to 4.3 Ma, with formation ages decreasing systematically from northwest to southeast (Liang et al., 2006). Zircon grains from the Yulong ore-bearing porphyries have higher Ce4 +/Ce3 + values than those from barren porphyries in the region (Liang et al., 2006). Abundant magnetite and hematite suggest that the ore-bearing porphyries are highly oxidized, reaching the HM buffer (Figs. 10, 11) (Sun et al., 2013b).

3.1.2.4. Qulong. The Qulong porphyry Cu–Mo deposit is now the largest porphyry-type deposit in China, with reserves of 10.4 Mt Cu @ 0.5% and 0.5 Mt Mo @ 0.03% (Table 1. Xiao et al., 2012). It is located in the Gangdese orogenic belt in southern Tibet. In addition to abundant hydrothermal anhydrite of up to 10% or more, magmatic anhydrite is also reported in unaltered granodiorite porphyry. These anhydrite occurrences indicate that the Qulong magmatic–hydrothermal system was highly oxidized and sulfur-rich, with abundant sulfates (Xiao et al., 2012; Yang et al., 2009) (Fig. 14). Given that it formed at ~ 14 Ma, which post-date the initial collision between Indian and

Fig. 13. An image of specularite from the Tongchang deposit in the Dexing porphyry Cu deposits. Specularite cutting across carbonate veins, indicating that it was formed very late, likely through hydrothermal alteration.

Fig. 14. Images of magmatic anhydrite from the Qulong giant porphyry deposit in the Gangdese porphyry belt, south Tibetan Plateau. Bi = biotite; Anh = anhydrite; Kf = K-feldspar; Py = pyrite; Cpy = chalcopyrite; Ap = apatite.

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3.1.2.6. Sar-Cheshmeh. The Sar-Cheshmeh is one of the top 20 largest porphyry Cu–Au–Mo deposit in the world with reserves of 14.4 Mt Cu @ 1.2%, 0.36 Mt Mo @ 0.03%, and 324 t Au @ 0.27 g/t (Cooke et al., 2005). It is located in southwestern Iran and is associated with several intrusive pulses of Miocene stocks (~12.2 Ma) (Mutschler et al., 2010), ranging in composition from diorite through granodiorite to quartz– monzonite (Hezarkhani, 2006). Molybdenum enrichment and deposition took place before Cu. Anhydrite is a popular mineral identified in most veins (Hezarkhani, 2006). Hematite has been reported, but it occurs as a secondary supergene ore forming mineral (Shahabpour, 1991). The Sar-Cheshmeh and several other giant Miocene porphyry deposits of Iran and Pakistan are located in a region undergoing continent–continent collision. The geodynamic and architectural controls on porphyry formation in this complex tectonic zone are unclear (Cooke et al., 2005). The ore-forming magmas in this section of the Tethys margin have high Sr contents (N 500 ppm) and Sr/Y (N 50) (Haschke et al., 2010), which are clearly adakitic (Defant and Drummond, 1990). Previous studies suggested that these magmas are a product of continental arc-style magmatism (Cooke et al., 2005). They were attributed to either the attempted subduction of the Arabian plate beneath the Eurasian plate, or the change from subduction of oceanic to continental crust (Cooke et al., 2005), or partial melting of a fertile copper- and sulfur-enriched arc crustal keel (Haschke et al., 2010). Further studies are needed on the formation of Sar-Cheshmeh and other Cenozoic porphyry deposits in the belt. 3.1.3. Porphyry Mo deposits Porphyry Mo deposits generally have higher SiO2 and lower FeO contents. It also has been proposed that porphyry Mo deposits have systematically lower oxygen fugacities than porphyry Cu (Au) and Cu–Mo deposits (Fig. 9). There are three major porphyry Mo deposit belts in the world, the Henderson–Climax (Klemm et al., 2008; Pettke et al., 2010; Seedorff and Einaudi, 2004a,b; Singer, 2008), Qinling–Dabie (Chen, 2013; Chen et al., 2000; Li et al., 2012a; Mao et al., 2008; Stein et al., 1997; Yang et al., 2013; Zeng et al., 2013; Zhang, Li et al., 2013) and Xing'an–Mongolia belts (Wu et al., 2011; Y. Zhang et al., 2013; Zeng et al., 2013). The oxygen fugacity can also reach the HM buffer during sulfate reduction and mineralization (Fig. 10), likely because of the lower FeO contents. 3.1.3.1. Henderson. Henderson porphyry Mo deposit is located in the Climax–Henderson Mo belt, Colorado. It consists of 12 Oligocene rhyolitic stocks in three centers, Henderson (oldest), Seriate, and Vasquez (deepest and youngest), at depths of ~ 1 km below the surface at the Red Mountain (Seedorff and Einaudi, 2004a), with a reserve of 1.24 Mt Mo @ 0.17 wt.% (Table 1). Previous studies proposed that, in the Henderson porphyry Mo deposit, Fe was leached at high to moderately high temperatures and then fixed in the rock at lower temperatures, first mainly as magnetite and then as pyrite and minor pyrrhotite and specular hematite (Seedorff and Einaudi, 2004a). Other porphyry deposits in the Climax–Henderson belt also contain hematite, indicating these deposits reached the HM buffer. 3.1.3.2. Shapinggou. The Shapinggou porphyry Mo deposit is the largest Climax-type Mo deposit in the world, with total proven Mo reserves of over 2.2 million metric tons @ 0.17 wt.% (Zhang, Li et al., 2013). It is located in the western Dabie Mountains, along the east extension of the East Qinling Mo mineralization belt (Chen, 2013; Chen et al., 2013; Gao et al., 2010; H.Y. Li et al., 2012; Han et al., 2013; N. Li et al., 2013; Stein et al., 1997; Yang et al., 2013; Zeng et al., 2013; Zhu et al., 2010), to the north of the Triassic suture between the north and south China blocks (Sun et al., 2002). Both Re–Os isochron and U–Pb zircon dating give the same age of ~111 Ma (Zhang, Li et al., 2013), which assigns it to the third mineralization pulse of the East Qinling Mo belt (Li et al., 2012a; Mao et al., 2008, 2011). High zircon Ce4+/Ce3+ ratios indicate

very high oxygen fugacity, comparable to those of Chuquicamata (Fig. 12) (Zhang, Li et al., 2013). 3.2. Linkage between oxidized magmas and porphyry deposits Copper, Au and Mo are all chalcophile elements, the behaviors of which are mainly controlled by reduced sulfur, e.g., sulfide, hydrosulfide − complexes (Sun et al., 2004a), or polysulfide (e.g., S2− 2 , S3 ) complexes (Sun et al., 2013b). Although most geologists agree that porphyry deposits are closely associated with oxidized magmas (Sillitoe, 2010; Sun et al., 2013b), it is still hotly debated as regarding to why oxidized magmas favor porphyry mineralization and how oxidized the oreforming magmas should be. Major disagreements include: (1) What is the most favorable oxygen fugacity for porphyry mineralization? Some workers proposed that ΔFMQ + 2 is a magic number for porphyry mineralization (Mungall, 2002), whereas ΔFMQ + 2 to +4 is the most favorable range of oxygen fugacity for porphyry deposits (Sun et al., 2013b). Yet others claimed that variations in oxidation state over typical ranges for arc magmas formed during plate subduction (ΔFMQ = 0 to +2) have no major effects on the potential to form porphyry Cu ± Au ± Mo deposits during plate subduction (Richards, 2011b). It is even argued that some porphyry deposits formed in reduced magmas (Cao et al., 2014; Rowins, 2000; Smith et al., 2012). (2) What is the main sulfur species in porphyry? Some workers concluded that sulfate is the predominant sulfur species in ore-forming porphyries (Cooke et al., 2011; Field et al., 2005; Liang et al., 2009; Sotnikov et al., 2004; Sun et al., 2013a,b), whereas others argue that the main sulfur species is SO2 dissolved in porphyry magmas (Richards, 2014; Smith et al., 2012), i.e., neither sulfate nor sulfide. (3) Is sulfide saturation during magma evolution important for porphyry mineralization? Some workers proposed that sulfide undersaturation is important for porphyry mineralization, i.e., no residual sulfide remains in the source (Li et al., 2012b; Liang et al., 2009; Sun et al., 2013b), whereas others argued that sulfide saturation and accumulation are of critical importance for porphyry mineralization (Wilkinson, 2013). In general, the porphyry mineralization process consists of 3 major phases: (1) source, i.e., extraction of chalcophile elements from the source; (2) transportation and concentration of ore-forming elements; and (3) deposition and fixation of the ore forming elements into orebodies. The relationship between oxidized magmas and porphyry deposits during these three steps are discussed below. 3.2.1. Sulfur oxidation, sulfide under saturation and residual sulfide Copper, Au and Mo are all chalcophile elements, with high partition coefficients between sulfide and melts, e.g., DCu = 1334 ± 210 (Patten et al., 2013), DAu = 4500 –11,200 (Mungall and Brenan, 2014), and DMo = 0.15–5.15 (Li and Audetat, 2013), respectively. Some experiments suggest that DMo increases with decreasing oxygen fugacities (Li and Audetat, 2013). Therefore, the behaviors of Cu and Au are strongly controlled by sulfide, whereas Mo is much less sensitive to sulfide, especially at high oxygen fugacities. During mantle melting, Cu as well as Au and Mo are all moderately incompatible elements, with partition coefficients ranging between those of Yb and Ce (McDonough and Sun, 1995; Sun et al., 2003a,b, 2004a). The estimated Cu abundance in the primitive mantle is 30 ppm (McDonough and Sun, 1995), whereas the Cu concentrations in MORB so far published range from 70 to 150 ppm (Hofmann, 1988; Sun et al., 2003b). When there is no residual sulfide in the mantle source, Cu and also other chalcophile elements become highly incompatible (Lee et al., 2012; Liu et al., 2014). Partial melting of mantle peridotite may form melts with high Cu contents up to 350 ppm at high oxygen fugacities, at very low degree of partial melting (Fig. 3) (Lee et al., 2012). Nevertheless, most arc magmas formed at much N 10% partial melting, resulting in much lower Cu contents (e.g., 100 ppm), which is not favorable than reducing magmas for porphyry mineralization. In contrast to mantle peridotite, MORB has sulfur abundances of about 1000 ppm, the effects of oxygen fugacity on

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residual sulfides, and, consequently, the Cu content in the melts, becomes much more obvious when considering partial melting of subducted oceanic slabs. The S abundance in the primitive mantle is estimated to be about 200–250 ppm (McDonough and Sun, 1995), and about 150 ppm in depleted mantle (Lorand, 1990; Mavrogenes and O'Neill, 1999). Mainly, the sulfur speciation in magmas is controlled by oxygen fugacity. Sulfide is the predominant sulfur species at oxygen fugacities lower than the FMQ buffer. Sulfate proportions start to climb up above the FMQ buffer. Most of the sulfur in the magmas is present as sulfate at ΔFMQ + 2 (Fig. 11) (Jugo, 2009; Jugo et al., 2005, 2010). Experiments show that sulfate is much more soluble than sulfide in magmas. The sulfur content at sulfide saturation (SCSS) increases solely as a function of increasing fO2 (Jugo, 2009): h i 2– SCSS ¼ S ð1 þ expð2:23ΔFMQ –2:89ÞÞ

ð9Þ

and  h i h i 6þ 2– ΔFMQ C ¼ 1:29 þ 0:45 ln S – ln S

ð10Þ

where ΔFMQC refers to the critical fO2 for simultaneous saturation of sulfide and sulfate (Jugo, 2009). The predicted SCSS content at sulfide saturation in basalts ranges from 1300 ppm at ΔFMQ − 1 and 1500 ppm at ΔFMQ + 0.5, to 7500 ppm at ΔFMQ + 2 and 1.4 wt.% at ΔFMQ + 2.3 (Fig. 15) (Jugo, 2009). At oxygen fugacities of ΔFMQ 0 to + 2.5, the more oxidizing the system, the higher the S contents in the magmas, and thus high oxygen fugacity is the most efficient way to eliminate residual sulfides (Lee et al., 2012; Sun et al., 2013b). The solubility of sulfide in magmas is independent of oxygen fugacity below the FMQ buffer (Mavrogenes and O'Neill, 1999), implying that most of the sulfur is removed in the form of sulfate under high oxygen fugacity (Jugo, 2009), and with its removal releasing much more of the chalcophile elements (Lee et al., 2012; Liang et al., 2009; Sun et al., 2004b). The average S concentration in MORB is ~ 1000 ppm (O'Neill and Mavrogenes, 2002), while laboratory experimental measurements have determined the S content at sulfide saturation in basalts to be 1300 ppm at oxygen fugacities of ΔFMQ 0 (Jugo, 2009). The small difference between the average MORB and experimental values is likely due to slight differences in oxygen fugacities and pressure effects, i.e., sulfide is less soluble under higher pressures (Mavrogenes and O'Neill, 1999). Given that most MORB forms through ~ 10% partial melting of the depleted MORB mantle, it extracts about 100 ppm of sulfur from the mantle source, leaving behind ~ 150 ppm of sulfur as residual sulfide. In case all the sulfur is present as sulfide, ~ 25% partial melting is needed to eliminate all the residual sulfides from the mantle source at ΔFMQ 0 (Fig. 3) (Lee et al., 2012). Such large degrees of partial melting would dilute the released Cu and produce picrite or even komatiite, not basalt. For the mantle wedge above subducting slabs, the estimated S content at sulfide saturation ranges from 1500 ppm (at ΔFMQ + 0.4) to 4500 ppm (at ΔFMQ + 1.7) (Jugo, 2009). As shown in Fig. 16, the highest oxygen fugacity in arc magmas may be NΔFMQ + 3, which allows up to 1.4 wt.% of S at sulfide saturation (Jugo, 2009). At oxygen fugacities below ΔFMQ + 1, 10% or less partial melting of the mantle wedge allows preservation of residual sulfides, assuming that the S abundances in the mantle wedge is also 250 ppm. At higher oxygen fugacity, no residual sulfides can be expected unless there is additional sulfur added from the subducting slab, which in actuality is probably the case. During metamorphism of the subducting slab, S is mobile from metamorphic rocks ranging in grade from blueschist to amphibolite (Sun et al., 2013a; Tomkins, 2010), such that they can release S and chalcophile elements to the mantle wedge (J.L. Li et al., 2013).

Fig. 15. Sulfur speciation curve vs. fO2 for basaltic glasses based on the S6+/ΣS estimates of XANES (in bold), EPMA (dashed line) (Jugo, 2009; Jugo et al., 2005, 2010). Also shown are fields of different tectonic settings, (A) Japan arc and Mexican arc; (B) MORB, OIB and mantle wedge, and arc. MORB = mid-ocean ridge basalt; IAB = Island arc basalt; BABB = Back arc basin basalt; OIB = Oceanic island basalt.

Nevertheless, normal arc rocks are not favorable materials for porphyry Cu mineralization. Arc magmas are probably neither unusually oxidized nor enriched in economic elements of interest, such as Cu (Lee et al., 2010; Lee et al., 2012). Instead, most porphyry Cu deposits are closely associated with adakite (Oyarzun et al., 2001; Sajona and Maury, 1998; Sun et al., 2010, 2012a; Thieblemont et al., 1997), or so called high Sr/Y porphyries (Chiaradia et al., 2012; Richards, 2011a). Most of the ore forming adakites were formed through slab melting (Sun et al., 2011, 2012a), deriving from the subducting oceanic crust itself rather than the overlying mantle wedge. MORB has higher Cu contents (70–150 ppm) (Hofmann, 1988; Sun et al., 2003b) than other portions of the subducted oceanic slab, and thus is likely the main contributor to mineralization. Therefore, partial melting of subducted oceanic slabs forms magmas with high Cu contents (Sun et al., 2011, 2012b), which plausibly explains the close associations between adakites, especially formed during ridge subductions, and porphyry Cu deposits. These papers, however, did not consider the effects of oxygen fugacity and dramatically underestimated the contributions of slab melting.

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Fig. 16. Oxygen fugacity of different tectonic settings. Note the range of oxygen fugacity for continental peridotite is much less than that of Bryant et al. (2007), after removing peridotite samples from arc settings. The oxygen fugacities of convergent margin magmas are systematically higher than those of intraplate settings. Magmas from intraplate settings without influence from plate subduction are too reduced for mineralization (Sun et al., 2013b). Modified after Bryant et al. (2007).

Slab melts formed at oxygen fugacities higher than ΔFMQ + 2 are favorable for porphyry Cu mineralization. MORB has an average S content of ~1000 ppm (O'Neill and Mavrogenes, 2002), from which it is easy to get the 22% and 10% partial melting needed to eliminate residual sulfides at ΔFMQ + 1.7 and ΔFMQ + 2, respectively. Correspondingly, the Cu contents of such melts would be 450 and 1000 ppm, respectively. Slab melts formed at ΔFMQ + 2 would have Cu contents more than twice as high as those formed at ΔFMQ + 1.7, and thus are considerably more favorable for Cu porphyry mineralization than mantle wedge melts. Only four times greater enrichments of Cu are needed to reach economic porphyry deposit proportions of ~ 4000 ppm. This ought to be easily achieved through magma evolution plus hydrothermal processes. At oxygen fugacities of ΔFMQ + 2.3 or higher, only 5% partial melting is needed to remove all the residual sulfides, forming melts with initial Cu contents of ~ 2000 ppm. In contrast, unrealistic N50% partial melting is needed to eliminate all residual sulfides at oxygen fugacities lower than ΔFMQ + 1, corresponding to a Cu content of less than 200 ppm in the melt. Residual sulfide would exist until the entire slab melted, in which case sulfide would remain the predominant sulfur species in both melt and slab (under reducing conditions of b ΔFMQ 0). In order to form supergiant porphyry deposits, e.g., El Teniente, up to 600 km3 of magma with a Cu content of 100 ppm would be needed to get the necessary amount of metal (Stern et al., 2011). Note that most

porphyry deposits have very small surface exposures, ranging in area from less than 1 km2 (mostly) to ~ 5 km2 (in rare cases). Thus, the magma chamber would need to be unrealistically thick, e.g., thicker than the continental crust. Alternatively, the magma must scavenge laterally for distances over 10 km and concentrates the Cu at one relatively small spot. None of these scenarios are practical. For adakitic melts with initial Cu contents of ~ 1000 ppm, the amount of magma needed measures only 60 km3 for the supergiant El Teniente. A magma chamber extending to a depth of several kilometers is sufficient. Slab melts need to pass through the mantle lithosphere before they eventually are emplaced in the upper crust, and thus may lose or collect chalcophile elements by interactions with the mantle. Sulfide saturated melts would lose some chalcophile elements. For slab melts formed at melting degrees higher than that require for SCSS, they would be undersaturated in sulfide and thus could assimilate more sulfides and chalcophile elements from the surrounding mantle while ascending, depending on the partitioning relationships between melt and residual sulfide in the mantle. For example, 10% partial melting of MORB with a S content of 1000 ppm at ΔFMQ N 2.3, would form a melt with a S content of ~1 wt.%, which is only half of its SCSS. This melt then would have the capacity to assimilate large amounts of sulfide. This plausibly explains the common association between high-Mg adakites and porphyry Cu deposits. Meanwhile, sulfide is kept undersaturated during the evolution of the oxidized magma, such that no sulfide segregation occurs, and Cu, Mo, and Au act as moderately incompatible elements. They would become enriched at an early stage of magma evolution, and dissolve into fluids following reduction accompanying magnetite crystallization (Sun et al., 2004a, 2013b), or other reduction processes. It has long been proposed that oxygen fugacities higher than ΔFMQ + 2 have reached the magic number for porphyry mineralization (Mungall, 2002; Sun et al., 2013b). As discussed above, this is exactly the point at which residual sulfide is eliminated by ~10% partial melting of a subducted slab, forming melts with initial Cu contents up to 1000 ppm. In contrast, at oxygen fugacity even only slightly lower, e.g., ΔFMQ + 1.7, the highest Cu contents that can be reached through slab melting is ~450 ppm. The systematically low oxygen fugacities in Japan arc volcanic rocks (bΔFMQ + 2) compared to those from the western American continents (up to ΔFMQ + 3) (Fig. 16), is an important factor controlling the distribution of Cu porphyry deposits, i.e., abundant porphyry Cu deposits are located in the western American continents, whereas none occur in Japan (Fig. 1) (Sun et al., 2012b, 2013b). 3.2.2. Sulfate reduction At oxygen fugacities higher than ΔFMQ + 2, most of the sulfur in magmas is present as sulfate, and such magmas usually have higher contents of total S and chalcophile elements. Copper however, is a chalcophile element, and its final deposition during mineralization should be controlled mainly by the behavior of reduced sulfur (Liang et al., 2009; Sun et al., 2004a, 2013b). Therefore, the final stage of miner2− alization inevitably requires the reduction of sulfate (S6+: HSO− 4 /SO4 ) 2− − 2− in the oxidized source magmas to sulfides (S : H2S/HS /S ) or − polysulfides (e.g., S2− 2 , S3 ) (Sun et al., 2013b). In principle, there are several ways to lower the oxygen fugacities during magma evolution. Among them are degassing of oxidized volatile species (i.e. CO2, SO3), assimilation of reducing country rocks (Ishihara and Matsuhisa, 1999; Smith et al., 2012), and reduction of sulfate by other elements, e.g., fractional crystallization of magnetite and even hematite (Liang et al., 2009; Sun et al., 2004a, 2013b). It has also been proposed that degassing of SO2 may lower the oxygen fugacity (Kelley and Cottrell, 2012). As illustrated in Eq. (11), degassing of SO2 indeed consumes oxygen, such that it may lower the oxygen fugacity of the system. However, degassing of SO2 cannot reduce

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sulfate to sulfide in the magmas. Instead, it drives the reaction (Eq. (11)) to the right, which leads to the oxidation of sulfide. 2H2 S þ 3O2 ¼ 2H2 O þ 2SO2

ð11Þ

Degassing of CO2 may reduce sulfate to sulfide, through a process similar to that described by Eq. (12). 2C þ SO3 þ H2 O ¼ 2CO2 þ H2 S

ð12Þ

As shown in Fig. 17, however, the C–CO2 buffer is lower than the SSO buffer (Mungall, 2002), therefore there is little C in the system at SSO buffer. Moreover, degassing of any type is not likely to be a key process during porphyry mineralization, which is usually occurring at depths of ~2–4 km (Sillitoe, 2010). Assimilation of reducing sediments may lead to sulfate reduction as proposed by previous authors (Shen and Pan, 2013; Smith et al., 2012) (also see Section 4). Nevertheless, assimilation occurring during the emplacement of porphyries would result in heterogeneous reduction, because reduction is mainly focused along the porphyry–country rock interface. This kind of reduction process has been reported for Baogutu (Shen and Pan, 2013), which was classified as a reduced porphyry Cu deposit associated with ilmenite-series magmas (Cao et al., 2014) (see detailed discussion in Section 4.3). Assimilation during magma chamber development may be better mixed and more homogenized. However, it would result in early sulfide saturation and segregation from the melt. Because sulfide is denser than silicate melt, it would tend to sink, which is more likely to occur during the mineralization of Cu–Ni deposits, rather than porphyry deposits. Elements that can exist in variable oxidation states and that are present in sufficient abundances to affect the redox state of the silicate Earth are C, H, S, and Fe (Mungall, 2002). In highly oxidized porphyry magmas, sulfate, CO2 and H2O are the dominant species, which are already oxidized and thus cannot reduce sulfate to sulfide, leaving ferrous iron as the only probable reducing agent (Sun et al., 2013b). Iron is a major element, with about 20 to 30% of the total Fe in a typical porphyry occurring as ferrous iron, corresponding to an oxygen fugacity of less than ΔFMQ + 2 (Fig. 18). Magnetite contains 66.7% ferric iron, whereas hematite contains 100% ferric iron. Therefore, the crystallization of magnetite and hematite tends to lower the proportions of ferric iron in the magma, which apparently would lower the oxygen fugacity. In the case of the Manus deposit in volcanic rocks, the Fe3 +/(total Fe) did not change much during magnetite crystallization (Fig. 18). This was attributed to sulfate buffering (Sun et al., 2004a), i.e., the oxidation of

Fig. 18. Ferric/ferrous Fe ratios and estimated oxygen fugacities of Manus glasses (Sun et al., 2004a). Oxygen fugacity of Manus glasses was calculated using the following equa  2 O3 tion: ln XFe ¼ a lnƒO2 þ bT þ c þ ΣdiXi, where, a = 0.196, b = 1.1492 × 104, c = XFeO −6.675, dAl2O3 = −2.243, dFeO = − 1.828, dCaO = 3.201, dNa2O = 5.854, dK2O = 6.215 (Kress and Carmichael, 1991).

ferrous iron is coupled with the reduction of sulfate to sulfide (Eqs. (13), (14)). Magnetite crystallization at Manus and many other arc volcanic rock series is coupled with dramatic decreases in Cu and Au (Sun et al., 2004a), which is an important factor in understanding Au and Cu mineralization (Sun et al., 2004a). The reduction in the amount of Cu and Au during magma evolution is well-known in arc magmas (Moss et al., 2001; Sun et al., 2011; Togashi and Terashima, 1997), and has been referred to as the magnetite crisis (Jenner et al., 2010). Indeed, magnetite crystallization/alteration is often taken as the controlling process for porphyry Cu and Au mineralization by causing sulfate reduction and presumably also oxygen fugacity fluctuations (Liang et al., 2009). There are several ways to describe the oxidation of ferrous Fe. In magma system, this is better described as oxidation of FeO by sulfate or another oxidant (Eq. (13)) (Sun et al., 2004a): 2−

2−

SO4 þ 8FeO ¼ 4Fe2 O3 þ S

:

ð13Þ

Eq. (13) is simplified, Fe2O3 represents the ferric Fe in Fe3O4. This reaction can be more precisely described by Eq. (14): 2−

SO4

2−

þ 12FeO ¼ 4Fe3 O4 þ S

:

ð14Þ

These reactions have no effects on the pH value of the system. This is one reason why the oxygen fugacity of the Manus magma did not change much (Sun et al., 2004a), or even decrease slightly during the crystallization of magnetite and sulfate reduction (Jenner et al., 2010) (see also Section 5).

Fig. 17. Sulfate reduction and oxygen buffers. CCO = carbon dioxide–carbon oxide buffer; SSO = sulfide–sulfur oxide buffer. After Mungall (2002) and Sun et al. (2014a,b).

3.2.3. Hematite–magnetite intergrowth Interestingly, in addition to abundant magnetite in porphyry deposits (Astudillo et al., 2010; Audetat et al., 2004; Baker et al., 1997;

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Dilles et al., 2011; Imai, 2005; Liu et al., 2011; Yang et al., 2002), hematite and specularite (a hydrothermal variety of hematite) are common in porphyry copper deposits (Figs. 10, 13) (Zhang, Ling et al., 2013; Sillitoe, 2010; Sun et al., 2013b). As mentioned previously, primitive hematite–magnetite intergrowths have been reported in many porphyry Cu deposits all over the world (Fig. 10), including nearly all major porphyry deposits in China (Sun et al., 2013b), as well as some deposits in South America (Ballard et al., 2002; Patricio and Gonzalo, 2001; Vila and Sillitoe, 1991) and the Southwest Pacific islands (Hedenquist et al., 1998; Imai et al., 2007), and Mongolia (Khashgerel et al., 2008). Most of the hematite and/or specularite formed during the late stages of porphyry mineralization (Sillitoe, 2010; Sun et al., 2013b). The occurrence of primitive hematite in closely association with magnetite strongly suggests a very high oxygen fugacity (Fig. 11), such that reaching the HM oxygen fugacity buffer (equivalent to ~ΔFMQ + 4) during porphyry Cu mineralization. This suggests that, in contrast to hydrothermal systems in the Manus backarc basin (Jenner et al., 2010; Sun et al., 2004a), the oxygen fugacity of porphyries increases during magnetite crystallization (Sun et al., 2013b). The further oxidation of magnetite is controlled by the pH values of the system (Sun et al., 2013b). In contrast to the magma system, ferrous Fe is mainly present as Fe2+ in aqueous fluids. In this case, the oxidation of ferrous Fe is better described by Eq. (15): 2−

2þ

SO4 þ 12Fe

2−

þ 12H2 O ¼ 4Fe3 O4 þ S

þ

þ 24H :

ð15Þ

As shown in Eqs. (2) and (1), HM and FMQ oxygen fugacity buffers do not change with pH. The oxidation potential of sulfate, however, depends strongly on pH (Figs. 11). The oxidation of ferrous iron by sulfate leads to a decreases in pH in ore-forming fluids within the porphyry system (Eq. (15)), which drives up the oxidation potential of sulfate. Consequently, the fO2 increases during the reduction of sulfate (Fig. 11) (Pokrovski and Dubrovinsky, 2011; Sun et al., 2013b). It is estimated that the amount of H+, released during the reduction of sulfate and oxidation of ferrous iron, may lower the pH down to +2, depending on temperatures, and if there are no pH buffers in the system to significantly elevate the oxygen fugacity (Sun et al., 2014a,b) and oxidize magnetite or ferrous Fe in hydrothermal fluids to hematite and/or specularite (Eq. (16)). 2−

SO4

þ 8Fe

2þ

2−

þ 8H2 O ¼ 4Fe2 O3 þ S

þ 16H

þ

2−

þ 8Fe3 O4 ¼ 12Fe2 O3 þ S

2−

:

þ

2þ

2þ

Fe

þ

þ 2H2 S ¼ CuFeS2 þ 3H þ §H2 þ

þ 2H2 S ¼ FeS2 þ 2H þ H2

−

2S3 þ 20H2 O þ 15Fe

2þ

2þ

−

¼ 2S3 þ 19Fe3 O4 þ 104H

¼ 6S

2−

þ

þ

þ 5Fe3 O4 þ 40H :

ð22Þ ð23Þ

Both of these reactions (Eqs. (22), (23)) also would lower the pH of the fluids, which in turn would elevate the oxidation potential of sulfate. Such conditions will promote the formation of hematite and would be favorable for porphyry mineralization. The oxidation potential of the − SO2− 4 –S3 reaction also rises with decreasing pH (Fig. 11), so that magnetite may be further oxidized to hematite by SO24 −, releasing OH− (Eq. (24)). This in turn increases the pH and promotes magnetite formation (Sun et al., 2013b), which may explain magnetite and hematite intergrowths (Fig. 10). In addition to direct oxidation of magnetite, hematite and specularite may form directly from fluids (Eq. (25)), releasing H+, which lowers pH and promotes further oxidation of ferrous Fe. Pure hydrothermal hematite and specularite can often be determined to have formed during the late stages of porphyry mineralization (Sillitoe, 2010), e.g., the specularite occurring in late stage veins at Dexing (Fig. 13). 2−

−

38Fe3 O4 þ 6SO4 þ 5H2 O ¼ 57Fe2 O3 þ 2S3 þ 10OH 2þ

2−

−

þ

þ 6SO4 þ 33H2 O ¼ 19Fe2 O3 þ 2S3 þ 66H

−

ð24Þ ð25Þ

38Fe

ð17Þ

In magmas, in contrast to fluids, FeO is more abundant than Fe2+. If S− 3 is indeed stable as proposed (Pokrovski and Dubrovinsky, 2011), then the reduction of sulfate in magmas first leads to an increase in pH (Eq. (26)), which is compensated by further reduction of S− 3 (Eq. (27)). Therefore, S− 3 may indeed cause more fluctuations in oxygen fugacities through a different reaction pathway, but it has no major influences on the final results (Sun et al., 2014a,b).

The formation of hydrothermal hematite will further lower the pH, whereas further oxidation of magnetite will not change the pH. All the sulfate reduction reactions provide S2 −, which promotes mineralization, forming chalcopyrite (Eq. (18)) and more pyrite (Eq. (19)) (Heinrich, 1990) and releasing H2. Hydrogen may further react with CO2 and/or O2 (Eqs. (20), (21)) or escape from the porphyry system during degassing or diffusive loss (Sun et al., 2014a,b). Cu þ Fe

2−

6SO4 þ 52H2 O þ 57Fe

ð16Þ

Or further oxidation of magnetite (Eq. (17)): SO4

The porphyry mineralization process requires a continuous reduction of sulfate to sulfide, coupled with consumption of oxygen and even reduction of CO2 to CH4. Given that porphyry magmas are highly oxidized, reaction (21) is not likely to occur. No significant amounts of methane have been detected in most cases, except in reduced porphyries (see details in Section 4). Based on high pressure experiments, it was proposed that the trisulfur ion S− 3 is an important sulfur species (Fig. 11) at pressures and temperatures (Pokrovski and Dubrovinsky, 2011) that are common for porphyry systems (Hedenquist and Lowenstern, 1994; Seo et al., 2009; Sillitoe, 2010; Sun et al., 2014a,b). It has been further proposed that in general geological processes where the tri-sulfur ion may have been involved need to be reconsidered (Pokrovski and Dubrovinsky, 2011), and this would require the investigation of different reaction pathways (Sun et al., 2013b). In fluids, the reactions can be described by Eqs. (22) and (23)

ð18Þ

ð19Þ

2H2 þ O2 ¼ 2H2 O

ð20Þ

4H2 þ CO2 ¼ CH4 þ 2H2 O

ð21Þ

2−

−

6SO4 þ 5H2 O þ 57FeO ¼ 2S3 þ 19Fe3 O4 þ 10OH −

2−

2S3 þ 5H2 O þ 15FeO ¼ 6S

þ

þ 5Fe3 O4 þ 10H

−

ð26Þ ð27Þ

For large porphyry deposits, the reduction of sulfate inevitably results in the oxidation of magnetite to hematite as the pH is lowered (Figs. 10, 11). Therefore, the hematite –magnetite intergrowths may be taken as a possible ore indicator in future prospecting for copper deposits.

3.3. Summary Sulfur is one of the most important geosolvents that controls the behavior of copper and other chalcophile elements, therefore it is essential

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to the understanding of mineralization processes of copper and a variety of other metal resources. Sulfate is the dominant sulfur mineral species in porphyries associated with large copper deposits. This is because the oxidation of sulfide to sulfate during partial melting is essential for the efficient extraction of chalcophile elements out of the source region, especially from subducted oceanic slabs, which have more than 5 times more sulfur than the mantle. In contrast, metals of porphyry Cu deposits are hosted in sulfides, which require low oxygen fugacity for their stabilization during the final stage of mineralization. Therefore, the key process of porphyry mineralization is oxidation and reduction of sulfur, controlled by ferrous/ferric Fe and pH values. Sulfate reduction in hydrothermal fluids lowers the pH and consequently elevates the oxygen fugacity of the system up to the HM buffer. The acid released during sulfate reduction causes alterations. In contrast, sulfate reduction in magmas does not change the pH, such that the oxygen fugacity is slightly reduced. All types of porphyry deposits, ranging from Cu + Au, Cu + Mo, and Mo deposits, may be oxidized sufficiently to reach the redox states of the HM buffer. 4. The association of porphyry deposits with reduced magmas Although most of the porphyry deposits are closely associated with oxidized magmas, a small group of porphyry deposits are reported as apparently related to reduced magmas with oxygen fugacities ranging from ΔFMQ − 0.5 to − 3 (Fig. 19) (Cao et al., 2014; Rowins, 2000; Smith et al., 2012). Reduced porphyry deposits so far reported are: 17 Mile Hill, Western Australia; the Minãs de San Anton, Mexico (Rowins, 2000); the Baogutu, northwestern China (Cao et al., 2014; Shen and Pan, 2013; Shen et al., 2010a, 2012); North Fork Deposit, west Central Cascades, western North America (Smithson and Rowins, 2005); Catface; Vancouver Island, British Columbia (Smith et al., 2012); and a few other small porphyry deposits (Rowins, 2000). Three criteria have been proposed for the classification of reduced porphyry Cu deposits: (1) lack of primary hematite and sulfate minerals, (2) rich in hypogene pyrrhotite and CH4, and (3) oxides belongs to the ilmenite-series, reduced, I-type granitoids (Rowins, 2000). Instead of primary hematite, magnetite, and sulfate (e.g., anhydrite, gypsum) that are frequently seen in oxidized porphyries, these “reduced” porphyry Cu–Au deposits contain abundant hypogene pyrrhotite, and commonly have carbonic-rich ore fluids with substantial proportions of CH4, clearly indicating a reducing condition (Cao et al., 2014; Rowins, 2000; Smith et al., 2012). More importantly, they are associated with ilmenite-bearing, reduced I-type granitoids (Cao et al., 2014; Rowins, 2000), in direct contrast to oxidized porphyries that are related to a magnetite-series mineralogy (Ishihara and Sasaki, 1991; Thompson et al., 1999). For example, ilmenite is an end-member of the hematite–ilmenite solid solutions with a very low hematite component, varying from XHem (0.01–0.1), for the Catface phase and XHem (0.04–0.09) for the Hecate Bay phase (Smith et al., 2012). Magnetite is nearly pure Fe3O4 with less than 2% Fe2TiO4 component for Catface (Smith et al., 2012) and Baogutu (An and Zhu, 2010; Cao et al., 2014; Shen and Pan, 2013; Shen et al., 2010a,b). Ilmenite is more abundant than magnetite in the reduced porphyries (Cao et al., 2014; Smith et al., 2012), e.g., all intrusive phases at Catface have accessory FeTi oxides with ilmenite/magnetite ratios of ~9:1 (Smith et al., 2012), indicating a relatively reduced oxidation state, i.e., belonging to the ilmenite-series (Ishihara, 2004). Experiments with apatite have shown that SO3 contents reflect the oxygen fugacity of the magmas (Peng et al., 1997). The uniform and low SO3 contents in apatite from reduced magma deposits were used to argue that the low oxygen fugacity of these porphyries is primary (Fig. 20) (Cao et al., 2014; Smith et al., 2012). The reduced porphyry deposits range in age from the Late Archean to the Oligocene (Rowins, 2000). Although the tonnages of these reduced porphyry deposits are generally much smaller than oxidized porphyry deposits, the mechanism that controls these mineralization

Fig. 19. Diagram of log fO2 versus temperature and 1000/T illustrating the oxygen fugacities of (A) oxidized porphyry deposits and (B) reduced deposits, Catface and Baogutu. The systematic difference in oxygen fugacities is likely not primary. Baogutu is hosted in ilmenite-series diorites, containing native antimony (An and Zhu, 2010), hypogenepyrrhotite, and methane-rich fluid inclusions. Previous studies suggested that assimilation occurring during the emplacement of porphyries resulted in reduction, which is mainly focused along the porphyry–country rock interface (Shen and Pan, 2013). Modified after Oyarzun et al. (2001) and Cao et al. (2014) and Smith et al. (2012), respectively.

process needs to be clarified. Some major questions to be answered are: How reduced magmas formed in an arc environment, where oxidized magmas are common, e.g., in western North America? In contrast, why are there no porphyry deposits in Japan, where the ilmenite-series is well-developed (Ishihara and Murakami, 2006)? How do Cu and Au get enriched in reduced magmas? What's the relation between oxidized and reduced porphyries? 4.1. Reduced magmas Arc magmas usually have high oxygen fugacities (Fig. 16) (Arculus, 1994; Ballhaus, 1993; Carmichael, 1991; Kelley and Cottrell, 2009; Sun et al., 2012a, 2013a,b), a condition that has been attributed to a variety

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Fig. 20. Diagram of Cl versus SO3 in apatite indicating the oxygen fugacities of Catface and Baogutu. The SO3 contents of Catface apatite are all lower than 0.15 wt.%, corresponding to oxygen fugacities below the NNO buffer. In contrast, the SO3 contents of Baogutu apatite range from less than 0.1 wt.% to much higher than 0.2 wt.%, corresponding to varied oxygen fugacities, ranging from below the NNO buffer to close to the NNO +1. HDP = hornblende diorite porphyry; GDP = Granodiorite porphyry; D = diorite. Modified after Cao et al. (2014) and Smith et al. (2012).

of mechanisms. The addition of subduction released oxidizing components to the mantle wedge or directly to arc magmas is proposed as the most straightforward way, e.g., water (Kelley and Cottrell, 2009), oxidizing fluids and/or melts (Brandon and Draper, 1996), fluids with oxidizing components, e.g., hematite, sulfate (X.M. Sun et al., 2007), slab melts with high Fe3+/Fe2+ratios (Mungall, 2002), as well as other oxidized components, e.g., carbonates, or other oxidized sediments. Arc magmas may also get oxidized during their evolution and ascent due to the following mechanisms (Ballhaus, 1993; Lee et al., 2005, 2010): (1) Changing oxygen buffers from graphite CO2 (CCO) equilibria in the mantle source to Fe3 +–Fe2+ equilibria after graphite has been eliminated by partial melting (Ballhaus, 1993); (2) Fractional crystallization of olivine and other minerals that favors ferrous iron over ferric iron (Carmichael, 1991); (3) Assimilation of oxidizing country rocks (Lee et al., 2005); (4) Degassing of reduced volatile species (e.g., H2, H2S and CH4) (Ballhaus, 1993; Lee et al., 2005); (5) Changing pH (e.g., lowering pH in a sulfate–sulfide system) (Sun et al., 2013a,b); (6) Magma recharging which preferentially enriches ferric Fe (Lee et al., in press). There are also several mechanisms that, in principle, may lower the oxygen fugacity of magmas: (1) Addition of reducing sediments from plate subduction (Takagi, 2004); (2) Assimilation of reducing country rocks (Ishihara and Matsuhisa, 1999; Smith et al., 2012); (3) Degassing of oxidized volatile species (i.e. CO2, SO3); (4) Addition of subduction released reducing fluids (Song et al., 2009); and (5) Fractional crystallization of magnetite and even hematite (Liang et al., 2009; Sun et al., 2004a, 2013a,b). As discussed previously (Section 3.2.2), the magnetite crystallization reduces the oxygen fugacity only in magmatic systems (reactions (13) and (14)), and may only lower the oxygen fugacity to the FMQ oxygen buffer. Here we discuss two reduced porphyry deposits in detail. 4.1.1. Evidence for reduced magmas of Catface porphyry deposit Catface is the largest reduced porphyry Cu deposit so far reported, emplaced at 40.4–41 Ma, with an indicated reserve of 56.9 Mt @ 0.4% Cu and an additional inferred resource of 262.4 million tons @ 0.38% Cu. Three potential mechanisms for the formation of Catface reducing magmas had been discussed before degassing was assigned as the causative process: (1) upwelling of the reducing asthenospheric mantle induced by the opening of a slab window during ridge subduction,

(2) evolvement of subducted reducing sediments, and (3) assimilation of reduced sediments during magma emplacement (Smith et al., 2012). Previous studies suggested that a slab window introduces hot, upwelling asthenospheric mantle in the subduction zone environment, forming non-arc-like alkalic and adakitic magmatism in the volcanic arc (Abratis and Worner, 2001; Groome and Thorkelson, 2009; H. Li et al., 2011, 2012; Kinoshita, 1997). The Kula–Farallon ridge in the northeastern Pacific began descending under the North American continent in Alaska in the Late Cretaceous (Madsen et al., 2006; Scharman et al., 2012), and then between 62 and 11 Ma migrated southward (Cole and Stewart, 2009; Cole et al., 2006; Liu et al., 2008), forming a semi-continuous forearc magmatic belt from Alaska to Oregon (Madsen et al., 2006). Ridge subduction forms oxidized adakites by slab melting (Defant and Drummond, 1990; Ling et al., 2009, 2013), followed by reduced A-type granites (Abratis and Worner, 2001; H. Li et al., 2012; Thorkelson and Breitsprecher, 2005; Yogodzinski et al., 1994). Interestingly, the Catface porphyry deposit is closely associated with Mt Washington, which is an adakite (Smith et al., 2012) and seemingly would support the slab window model. Nevertheless, adakites formed by slab melting have much higher initial Cu contents than mantle-derived melts, and thus are favorable for porphyry mineralization (Sun et al., 2010, 2011), suggesting that the reduced Catface porphyry could be a country rock hosting the ore deposit, rather than the causative porphyry. Subduction of reducing sediments also reduces the oxygen fugacity of arc magmas, as seen in the Japan arc (Takagi, 2004). Carbon rich sediments in the descending plate react with H2O to produce CH4 and CO2 (Ballhaus, 1993; Takagi, 2004), or directly releases CH4 by devolatilization (Song et al., 2009). This process, however, was excluded from further consideration based on the low 87Sr/86Sr of Catface (Smith et al., 2012). Methane, however, may decouple from Sr, such that Sr isotopes cannot give a conclusive answer. Nevertheless, the oceanic plate subducting underneath Japan is Cretaceous in age, which has experienced eight Ocean Anoxic Events (Jenkyns, 2010), and thus contains abundant organic rich sediments. In contrast, the oceanic plates subducting underneath the western North American continent are younger and have experienced only one Ocean Anoxic Event. Moreover, ocean ridges are young and generally sediment-poor. Therefore, subduction of reducing sediments is not a plausible explanation for the low fO2 shown by the Catface porphyry. The third mechanism proposed for the reduced condition of the Catface porphyry was by assimilation of reduced sediments. The Catface porphyry may have ascended through and interacted with graphite-rich reducing country rock known to exist in the region (Smith et al., 2012). The reduction of ferric Fe through reaction with reducing materials, e.g., graphite, in sediments was proposed as a potential factor responsible for the low fO2 of the parental Catface magma. This mechanism was excluded based on Sr isotopes (see details below) (Smith et al., 2012). The authors then proposed that the reduced Catface porphyry was due to degassing of SO2 from the magma. This argument is based on the low S content in the rock's apatite and a recent work by Kelley and Cottrell (2012), which proposed that fractional crystallization coupled with degassing of S in a sub volcanic arc magma chamber can result in a reduction of N 2 log fO2 units in the magma. This calculated reduction was based on the rock's ferric iron content and S in melt inclusions, and would hold true for magmas varying in composition from basaltic andesite to dacite (Kelley and Cottrell, 2012). Degassing of SO2 indeed has a significant influence on the oxygen fugacity in volcanic systems. It is, however, not likely to be a major process in plutons, which should experience much less degassing. Moreover, apatite from the Catface has a homogenous S content throughout the crystal (Smith et al., 2012), which argues against major degassing induced S loss during magma evolution. Experiments on felsic liquids showed that the SO3 content in apatite increase with increasing fO2 from 0.04 wt.% at the FMQ buffer, up to N1.0 wt.% at the HM buffer (Peng et al., 1997). As pointed out by Smith et al. (2012), apatite

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crystallizes below 950 °C in silicic calc-alkaline magmas (Green and Watson, 1982). If the magma at Catface was reduced during degassing of SO3, there should be compositional zonation in the apatite, which is not the case (Fig. 20) (Smith et al., 2012).

4.1.2. Evidence for reduced magma in the Baogutu porphyry deposit Baogutu is located in West Junggar, Central Asian Orogenic Belt, and is the second largest reduced porphyry Cu deposit so far reported, with controlled reserves of 0.63 Mt Cu @ 0.28 wt.%, 14 t Au @ 0.1 ppm, 1.8 × 104 t Mo @ 0.011 wt.% and 390 t Ag @ 1.8 ppm (Cao et al., 2014). It contains native antimony (An and Zhu, 2010), abundant hypogene pyrrhotite, and methane-rich fluid inclusions. The low apatite SO3 content, whole rock Fe2O3/FeO, and fluid compositions indicate a low fO2 of ~ NNO for the magma and NNO–NNO − 2 for associated hydrothermal fluid (Cao et al., 2014; Shen et al., 2010b). Based on H–O and sulfide S–Pb isotope data, it has been argued that the methane-rich ore-forming fluids were derived from a deep mantle source with little contamination from penetrated sediments (Cao et al., 2014). The assimilation model, however, does not contradict with a deep mantle source for the S, Pb, nor Cu. In fact, most Cu should have come from slab melts (Sun et al., 2011, 2012b), as also should have come the S and Pb, which are present in abundances similar to mantle derived magmas. The H isotopes of fluid inclusions from Baogutu have very low δD values (Shen et al., 2012; Zhang et al., 2010), which are much lower than for either magmatic fluids or metamorphic fluids (Fig. 21). This was interpreted as the main evidence for degassing (Cao et al., 2014). Degassing, however, should result in heavier hydrogen and oxygen isotopes in the residual magmas. The authors used published heavy H values from circum-Pacific volcanic gases (Giggenbach, 1992b) to argue that degassing should leave behind fluids with even lighter H than contained in the magmatic and metamorphic fluids (Cao et al., 2014). The heavy H of circum-Pacific volcanic gases, however, was originally explained as having arisen from the addition of seawater to the gases (Giggenbach, 1992a). Basic principles of isotope geochemistry dictate instead that in degassing, i.e., evaporating, the remaining liquid becomes isotopically heavier in H and O. In addition,

Fig. 21. δD versus δ18O diagram and calculated isotopic composition of waters in hydrothermal fluids derived from measured isotopic composition of quartz and its fluid inclusion for the Baogutu deposit from Cao et al. (2014). Reference lines and boxes are as follow: meteoric water line (Craig, 1961), felsic magmatic water (Taylor, 1992), residual water in intrusion after degassing and crystallization (Taylor, 1974), low-salinity vapor discharges from high-temperature volcanic fumaroles (Giggenbach, 1992b), primary magmatic water area, metamorphic water area and multiple formation water area (Hoefs, 2004). H–O isotopic data are from (Zhang et al., 2010) and (Shen et al., 2012). Low-salinity vapor shows strong seawater signals, which cannot represent “degassing” fluids.

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degassing cannot explain the decoupling between δD and O isotope compositions. Moreover, the degassing of porphyries should be much less pronounced than the degassing of volcanic rocks. The systematically lighter H coupled with a slightly lighter O can best be explained by the addition of meteoritic water. Baogutu is located in the center of the Eurasian continent. Meteoritic water there should be very light in H and O isotope composition. Addition of such water should dramatically reduce the δD of the magma with a much lesser effect on O because magmas have low hydrogen contents and abundant oxygen contents. 4.2. Source of copper For oxidized magmas, as discussed in Section 3, excess sulfur is precipitated in the form of sulfate, leaving behind less residual sulfide, such that Cu, Au and other chalcophile elements become enriched in magmas (Lee et al., 2012; Sun et al., 2012a, 2013a,b). In contrast, reduced magmas usually retain residual sulfides, such that they have low primary Cu contents (Fig. 3) (Lee et al., 2012). Then, how do reduced magmas get enriched in Cu? As pointed out by previous authors, porphyry deposits associated with reduced magmas are generally small (Cao et al., 2014). This was attributed to the originally low Cu and Au contents of the magmas, less magmatic fluids released due to deep emplacement, or tectonic settings that are not favorable for porphyry mineralization (Cao et al., 2014). They also argued that significantly lower sulfur solubility in reduced melt (Jugo, 2009; Jugo et al., 2005, 2010) keeps S2− as the dominant sulfur species, which potentially isolates sulfides from the magma during its migration to the site of final emplacement, thus producing only relatively small chalcophile endowments (Cao et al., 2014). None of these arguments are convincing to us. First, lower Cu and Au contents would first result in lower grade, not necessarily smaller tonnages. The grades of porphyry Cu and Au deposits associated with reduced magmas are comparable to, if not higher than, those associated with oxidized magmas (Cooke et al., 2005; Rowins, 2000). Second, there is no systematic difference in terms of emplacement depths and tectonic settings between reduced and oxidized porphyry deposits. More importantly, compared to oxidized magmas, reduced magmas indeed have lower sulfur contents with S2− occurring as the dominant sulfur species (Jugo et al., 2010), but this does not necessarily mean higher S2− in reduced magmas as claimed by Cao et al. (2014). Given that the solubility of sulfide is independent of oxygen fugacity, but increases with decreasing pressure under reducing conditions (Mavrogenes and O'Neill, 1999), the high proportions of S2− cannot “isolate sulfides from the magma during migration” as proposed. Based on a synthesis of theoretical, experimental, and field data, it has been proposed that Cu and Au can be transported via the vapor phase to distal sites as far as several kilometers away from the causative porphyry due to fluid boiling or immiscible phase separation. Consequently, the source porphyry becomes a low-grade sub-economic Cu– Au core or failed porphyry Cu system (Rowins, 2000), which is more or less similar to an epithermal ore system. Experiments find that the main transporting agents of Cu at the porphyry level are brines and that models based on transporting copper in the vapor phase are incorrect (Lerchbaumer and Audetat, 2012). It is further demonstrated, using experimental studies, that brine–vapor separation in porphyry deposits does not cause selective Cu transfer to the vapor, but is more likely to destabilize Cu complexes and promote copper ore deposition during decompression and unmixing of the two fluid phases. In contrast, Au may be selectively transferred into the vapor phase, allowing its transport from the deeper porphyry copper deposits to form shallower epithermal gold deposits (Seo and Heinrich, 2013). This explains the Au-rich capping feature of many reduced porphyry deposits. Meanwhile, Cu may be transported to distal sites through normal fluids. Such a reduced porphyry Cu–Au mineralization model does not contradict the current understanding of porphyry Cu–Au formation. In fact,

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the recognition of reduced porphyry Cu–Au systems encourages a search for distal sites that are favorable for focusing and precipitating Au and Cu-rich vapors (Rowins, 2000). This implies that such porphyry deposits themselves are not genetically related to reduced magmas, i.e., they are not the causative magma, but the host rock. 4.3. Formation of reduced porphyry deposits Porphyry deposits related to reduced magmas have two other distinct characteristics. They are associated with carbonic-rich ore forming fluids (Cao et al., 2014; Rowins, 2000; Smith et al., 2012), and multiple intrusive events (Cao et al., 2014; Rowins, 2000; Smith et al., 2012). In addition to the transportation of Cu and Au from the causal porphyry to distal sites of deposition (Rowins, 2000), we propose two other ways that may form reduced porphyry deposits: (1) by reduction of oxidized magmas during their ascent in the crust, and (2) by using reduced magmatic rocks that are present only as country rocks to host ore deposits originating from underlying oxidized magmas. 4.3.1. The formation of the Catface porphyry deposit As discussed above (Section 3.2.2), assimilation of reduced sediments may have reduced the oxygen fugacity of the Catface porphyry, a possibility that was excluded based on Sr isotopes. The 87Sr/86Sr ratio of the Catface intrusions is 0.704, whereas those of the graphiterich metasediments of the Pacific Rim terrane range from 0.706 to 0.708 (Smith et al., 2012). We find such small differences in Sr isotopes to be negligible, and thus do not argue strongly against the assimilation model. Graphite and especially methane need not be considered coupled with silicate minerals, such that the reducing action of the graphite and methane is not proportionally related to the assimilation of silicates nor sulfides, thereby decoupling them from the Sr, Pb, and S isotopes. Relationship between Sr and C aside, the amount of C needed is very small. The total iron (expressed as Fe2O3) content ranges from 2.25 wt.% to 5.53 wt.% for the Catface porphyry. One carbon atom reduces 4 ferric Fe atoms (Eq. (28)), assuming all the Fe in the parental Catface magmas was ferric Fe, and all the ferric Fe was reduced by graphite, then only 0.04 to 0.1 wt.% of graphite is required, which may only have had a limited influence on Sr isotopes. More interestingly, the Catface intrusions have fO2 values nearly identical to the C–CO– CO2 buffer at similar pressure and temperature conditions (Smith et al., 2012), which would support the assimilation of graphite-rich sediments. C þ 2Fe2 O3 ¼ 4FeO þ CO2

ð28Þ

Reduced sediments usually also contain methane (Cao et al., 2014), which is a more efficient reductant, i.e., 1 methane molecule reduces 8 ferric Fe atoms (Eq. (29)), such that only ~0.02 to 0.05 wt.% of methane is needed. Therefore, the parental Catface magmas may have acquired a low fO2 during their ascent through the crust. The homogenous SO3 in apatite may simply be due to an early reduction of the magmas, before apatite crystallization. CH4 þ 4Fe2 O3 ¼ 8FeO þ CO2 þ 2H2 O

ð29Þ

Considering that the Catface porphyry is not adakitic, it may well be simply the country rock that hosts the deposit. The Catface porphyry deposit is closely associated with a nearby pluton on Mt Washington (35–41 Ma), which is an adakite (Smith et al., 2012). The mineralization (40.9 Ma) and the emplacement age of the Catface porphyry (40.4– 41 Ma) (Smith et al., 2012) are both within the age range of the Mt Washington adakite. Adakites along the eastern Pacific rim are mostly formed by slab melting (Liu et al., 2010; Sun et al., 2012a,b), which are likely to have high initial Cu contents, and are thus favorable for porphyry mineralization (Sun et al., 2010, 2011). In contrast, the Cu contents of the asthenosphere (~ 30 ppm) (McDonough and Sun, 1995) and the continental crust (~27 ppm) (Rudnick and Gao, 2003) are much lower

than MORB (~ 100 ppm) (Sun et al., 2003b), such that melt derived from them should have much lower Cu contents and be less favorable for porphyry Cu mineralization. Therefore, the reduced Catface porphyry is likely to be a country rock that only hosts the deposit. Based on the above discussion, we propose the following model for the formation of the Catface porphyry. Intrusions in the Vancouver Island range in age from 51 to 35.5 Ma (Madsen et al., 2006), spanning the time during which the Kula–Farallon ridge collided with the continent. Accompanying the oblique subduction of the ridge (Madsen et al., 2006), its two limbs separated, with a slab window opened in between. The first limb formed early through partial melting of the hot subducting plate. The resulting adakites may then have been reduced through reaction with the thick (maximum thickness of 4 km) carbon-rich Cretaceous Nanaimo Group sediments (Madsen et al., 2006) as they migrated upward. Consequently, sulfate is reduced to sulfide, leaving behind Cu-rich sulfide accumulations in the lower crust. This is followed by emplacement of the Catface porphyry at 41– 40.4 Ma (Smith et al., 2012), which took place at a time that the slab window was open. The mantle derived parent magmas of the Catface porphyry are expected to have been more reduced (near the FMQ oxygen fugacity buffer) and drier than the adakite, and to have a lower initial Cu contents. Nevertheless, it became wetter and even more reduced (ΔFMQ − 0.3 to − 3) with a higher Cu content after assimilating the Cu-rich sulfide accumulations in the lower crust. Meanwhile, ridge subduction induced compression and consequent uplift and erosion occurred, which favored the exposure of the porphyry deposit. Oxidized Mt Washington adakite was emplaced at 41–35.3 Ma during the subduction of the west limb of the ridge, bringing more ore forming fluids into the Catface porphyry. We infer that large proportions of these adakites are still buried due to the dramatically lessened compression, uplifting and erosion following ridge subduction.

4.3.2. The formation of the Baogutu porphyry deposit Based on previously obtained fluid inclusion H–O isotope data and sulfide S–Pb isotope data, it was proposed that the methane-rich ore forming fluids in the Baogutu porphyry deposit were derived from a deep mantle source with little contamination from sedimentary components (Cao et al., 2014). As discussed above (Section 4.1.2, Fig. 21), the authors' understanding about H–O isotopes is in error, whereas only small amount of C is enough to lower the oxygen fugacity of the Baogutu low Fe magmas. Therefore, S–Pb isotopes cannot place any decisive constraints on the role of assimilation in forming the deposit. As pointed out by the authors, detailed studies are needed to clarify the origin of the CH4. More importantly, there are three intrusive phases at Baogutu, which from old to young are: (1) the main diorite phase, (2) dikes of diorite porphyry and granodiorite porphyry intruding the early diorite stock, and (3) dikes of hornblende diorite porphyry intruding all the three phases. All the samples with the exception of two were collected from the diorite, which are not porphyry at all. The reduced diorite magma may have no relation with the porphyry mineralization, except to act as a host rock. We propose that the reduced features of the Baogutu porphyry deposit are secondary and occurred during emplacement, thus having no major influence on the mineralization process. A similar model has been proposed based on geochemical and mineralogical studies by Shen and Pan (2013), in which mineral composition data suggest that the primary magma of the Baogutu porphyry deposit is oxidized. The heterogeneous and reduced characteristics of the deposit are attributed to significant country-rock contamination after emplacement (Shen and Pan, 2013) in an arc setting (Shen et al., 2013a,b). Similar to Catface, there are also adakites in Baogutu. The Baogutu adakites have been attributed to the mixing of ~ 95% slab melt with ~ 5% sediment-derived melt in the Late Carboniferous close to a subducting spreading ridge (Tang et al., 2010). Therefore, more work

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on porphyry dykes are needed to clarify the origin of the Baogutu porphyry deposit. 4.4. Summary Reduced magmas are not favorable for porphyry mineralization. The reduced porphyry deposits so far reported are either simply distal host rocks located as far as several kilometers away from or above the causative porphyry buried underneath, or originally oxidized magmas that were reduced through assimilation of reducing components. Degassing of SO2 is not likely to be a main process for forming porphyry deposits because it cannot explain the features of reduced porphyries. Careful studies are needed to identify the causative porphyries. Oxidized adakitic (slab melts) rocks are overwhelmingly the most favorable candidates for porphyry mineralizations. 5. Discussion 5.1. The oxygen fugacities at convergent margins The oxygen fugacity of the mantle and volcanic arc has been under study for a long time (Ballhaus, 1993; Carmichael, 1991; Cottrell and Kelley, 2013; Kelley and Cottrell, 2009; Lee et al., 2005, 2010; Parkinson and Arculus, 1999). The consensus is that the oxygen fugacity of arc magmas is systematically higher than that of MORB (Fig. 16) (Carmichael, 1991; Kelley and Cottrell, 2009; Parkinson and Arculus, 1999; Sun et al., 2012a, 2013a,b). Nevertheless, it is still hotly debated as regards to how arc magmas get oxidized. As mentioned in Section 3, a variety of mechanisms have been proposed. The most straightforward way to elevate the oxygen fugacity of arc magmas is by the addition of oxidizing components to the mantle wedge or directly to the magmas. Water is the most abundant volatile component released during plate subduction. It has been proposed that H2O reacts with FeO, forming Fe2O3 with the release of H2 (Eq. (30)) (Brandon and Draper, 1996; Kelley and Cottrell, 2009). Most of the ferric Fe will be transferred into the melt because ferric Fe is more incompatible than ferrous Fe (Lee et al., in press). This reaction, however, is not controlled by water. Instead, it is controlled by ferrous Fe, which is more stable under high pressure. Recent experiments show that water and H2 can coexist as two immiscible phases. This immiscibility implies that water is stable in the mantle under high pressure (Bali et al., 2013). More accurately, the ferrous Fe in fluids reacts with water, releasing H2 and H+ (Eq. (31)). This is supported by the abundance of oxidized components, e.g. magnetite–hematite and sulfates, found in the subduction released fluids of ultrahigh pressure quartz veins, which may elevate the oxygen fugacity of the mantle wedge (X.M. Sun et al., 2007). It has also been proposed that slab melts are the most efficient way to transfer high Fe3+/Fe2+ ratios responsible for high oxygen fugacity in adakites (Mungall, 2002) and arc magmas (Brandon and Draper, 1996). H2 O þ 2FeO ¼ Fe2 O3 þ H2 2þ

3H2 O þ 2Fe

¼ Fe2 O3 þ H2 þ 4H

ð30Þ þ

ð31Þ

The addition of other oxidized materials, e.g., Fe3+, C4+, and S6 +, from subducted sediments and oceanic crust to the mantle wedge may elevate the redox states of the mantle as well (Fig. 18) (Evans et al., 2012). Arc magmas may also get oxidized during their evolution and ascent by processes accompanying melting, crystallization, assimilation, degassing, etc (Ballhaus, 1993; Lee et al., 2005, 2010). For example, graphite and diamond are stable in the deep mantle. During mantle melting, CO2 is incompatible whereas C is compatible, and graphite may be consumed through oxidation melting (Stagno et al., 2013). Most of the C in magma occurs as CO2, therefore the oxygen buffer changes

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from graphite–CO2 (CCO) equilibria in the mantle source to Fe3 +– Fe2+ equilibria as the magma ascends (Ballhaus, 1993). Consistently, it becomes more oxidizing with decreasing depths and pressures (Stagno and Frost, 2010; Stagno et al., 2013). Moreover, most mantle minerals, such as olivine, favor ferrous Fe over ferric Fe, and fractional crystallization of these minerals elevates the Fe3+/Fe2+ (Carmichael, 1991). Given that ferric Fe is highly incompatible, magma recharging will further enrich the ferric Fe (Lee et al., in press), resulting in higher oxygen fugacities. In addition, assimilation of oxidizing country rocks (Lee et al., 2005), degassing of reduced volatile species (e.g., H2, H2S and CH4) (Ballhaus, 1993; Lee et al., 2005) and lowering of pH values in a sulfate–sulfide dominated system (Sun et al., 2013a,b) may also elevate the oxygen fugacity of the magmas. Based on similar Zn/FeT ratios of mantle peridotite and primitive arc magmas, it has been argued that primary arc magmas are not necessarily oxidized (Lee et al., 2010). Increasing Zn/FeT ratios with decreasing MgO content argues that the high oxygen fugacity of arc magmas is acquired through magma evolution (Lee et al., 2010), indicating that continuing chemical evolution during magma transport and emplacement may have major effects on the oxidized characteristics of arc magmas. This may equally explain the diversity in oxygen fugacities at convergent margins. Nevertheless, high oxygen fugacities developed during magma evolution may provide little or no contribution to eliminating residual sulfides, and thus little contribution to porphyry mineralization. This decoupling may partially explain why most arc magmas are highly oxidized, whereas only a small portion of them form porphyry Cu deposits. 5.2. The difference between porphyry and epithermal in terms of oxygen fugacity Most Cu porphyry deposits form at depths of 2–4 km, and are usually associated with epithermal deposits at shallow depths if not removed by erosion (Figs. 22d, 23a, b) (Cooke et al., 2011; Hedenquist et al., 1998; Heinrich, 2005; Heinrich et al., 2004; Hollings et al., 2005; Sillitoe, 2010). This seemingly implies that porphyry and epithermal deposits are closely related. Many epithermal deposits, however, are not linked to porphyry deposits. For example, epithermal deposits formed in the mid-ocean ridge and backarc basins usually do not have any associated porphyry deposit. This is probably mainly because the crust in backarc basins is too thin. Regardless, the oxygen fugacities of epithermal deposits have a much larger ranges than porphyry deposits. 5.2.1. Magnetite crisis The magnetite crisis (Jenner et al., 2010) refers to the dramatic decreases in Cu and Au during magnetite crystallization in arc volcanic rocks (Sun et al., 2004a), which is a common phenomenon (Moss et al., 2001; Sun et al., 2011; Togashi and Terashima, 1997). This magnetite crisis is taken as a main process that leads to the formation of oreforming fluids, responsible for the Au and Cu mineralization of both epithermal and porphyry deposits (Liang et al., 2009; Sun et al., 2004a, 2013a,b) (see also Section 3): Sulfate is reduced to sulfide during magnetite crystallization, forming hydrosulfide complexes, which scavenge chalcophile elements into fluids and subsequently transport the metals to favorable places for forming deposits (Sun et al., 2004a). This mineralization model has been challenged by some later studies. Although using the same set of samples (Jenner et al., 2012), and essentially repeating the results of the previous study (Sun et al., 2004a), it was proposed that magnetite fractionation triggers sulfide saturation (Jenner et al., 2010). This conclusion was drawn based mainly on the behavior of Se, which was assumed as a proxy that follows S closely during magmatic evolution except that it is not lost during low-pressure (near sea-floor) degassing (Jenner et al., 2010). Fluid extraction, however, is different from degassing, and thus Se provides no constraints on the process. In fact, as shown in the supplementary information, sulfide is undersaturated in Manus glasses (Fig. 24) (Sun et al., 2004a).

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Fig. 22. Different models for the mineralization of porphyry deposits. After (A) Lee (2014): Mantle-derived magmas (red) intrude into the cold upper plate (black) of subduction zones, generating crystallizing magma chambers. Thick continental arcs are more favorable for porphyry mineralization because of accumulation of copper-rich sulfide cumulates; (B) after Wilkinson (2013), which also proposes that sulfide saturation and accumulation are very important for the formation of porphyry Cu deposits; (C) after Richards (2011b), illustrating porphyry and epithermal mineralizations in arc and postcollisional settings; (D) after Richards (2013), highlighting features or processes that may result in supercharging of these systems to form giant deposits.

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Fig. 22 (continued).

There are indeed sulfide inclusions in mineral phenocryst from Manus glasses (Fig. 25) (Sun et al., 2004a). The sulfide phase contains high Au and Ag contents in addition to Cu in sulfides, but no other chalcophile elements, such as Ni, Re, and Pt (Jenner et al., 2010). This was used to argue that the sulfide phase is crystalline rather than an immiscible sulfide melt (Jenner et al., 2010). Crystallization of sulfide cannot explain this phenomenon, because Ni and Pt are even more chalcophile than Cu, with partition coefficients between sulfide and silicate melts of several hundreds and more than 20 thousand, respectively (Table 2). We propose that these sulfides formed directly from magmatic fluids. As shown in Fig. 25, sulfides in phenocrysts are associated with fluid inclusions (Sun et al., 2004a). It is very likely that the preferential enrichments of Cu and Au over

Ni and Pt are controlled by sulfide complexes in fluid. This may explain why there is no Ni and Pt in most porphyry and epithermal deposits. The lack of Re in these sulfides can be plausibly interpreted by the high oxygen fugacity, under which most of the Re is Re6+, and thus behaves as lithophile rather than chalcophile elements. As shown in Fig. 26, in contrast to Cu and Au (Sun et al., 2004a) and also V, Co, Pt, and Se (Jenner et al., 2010), Re concentrations in Manus glasses keep increasing when magnetite starts to crystallize, and then decrease gradually (Sun et al., 2003a). This can best be explained by the reduction of Re6 + to Re4 + during magnetite crystallization. Correspondingly, Re changes from incompatible (Sun et al., 2003a,b,c), to compatible (Mallmann and O'Neill, 2007).

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Fig. 23. Cartoons illustrate the relationship between some epithermal deposits with porphyry deposits underneath. Modified after: (A) Hedenquist and Lowenstern (1994) and (B) Heinrich (2005). Note, not all epithermal deposits are associated with porphyry deposits.

5.2.2. Oxygen fugacity and open systems In contrast to the dramatically elevated oxygen fugacities during magnetite crisis, the oxygen fugacity of the Manus magmas did not change much, or even lessened slightly (Jenner et al., 2010) during the crystallization of magnetite and sulfate reduction (Sun et al., 2004a). This can be interpreted as a magmatic process, i.e., the magnetite crystallization and sulfate reduction recorded in glasses occurred during magma evolution. As mentioned in Section 3.2.2, magnetite crystallization in magmas reduces sulfate without changing the pH values. The solubility of sulfides in fluids is very high, with partition coefficients

of ~ 500 (Keppler, 2010). Sulfides formed through sulfate reduction are released into magmatic fluids, where they scavenge chalcophile elements (Jenner et al., 2010; Sun et al., 2004a). Meanwhile, the reactions of Eqs. (16) and (17) are driven to the right, promoting sulfate reduction. Hydrothermal magnetite and even hematite may also form in fluids released from arc volcanic magmas, e.g., Manus glasses, releasing H+. Nevertheless, volcanic magma systems are much more open than porphyry systems, so that the formation of hydrothermal iron oxides does not necessarily affect the oxygen fugacity of the arc magmas.

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sulfides. This explains why epithermal deposits are distributed in a variety of tectonic settings, ranging from mid-ocean ridges and backarc basins to arcs. 5.3. Adakite, slab melting, ridge subduction and porphyry Cu deposits

Fig. 24. Sulfide contents in Manus submarine volcanic glasses, showing that sulfide is far below saturation (after Sun et al., 2004a). Therefore, sulfide segregation induced by sulfide saturation as proposed by previous authors (Jenner et al., 2010) is not a feasible way to explain the magnetite crisis. Filled and open circles represent volcanic glasses analyzed by different authors, filled triangles represent melt inclusions.

Given that volcanic systems are far more open, the initial Cu, Au contents are not of critical importance to porphyry mineralization. Metals can be leached out as far as there is enough water circulation. Therefore, it does not require very high oxygen fugacities to eliminate residual

Fig. 25. Images of sulfide inclusion in an olivine phenocryst of Manus glass under (A) reflected light and (B) and transparent light. Only one big sulfide globule is clearly identified under reflected light, which seemingly indicates sulfide saturation. This grain is actually associated with fluid inclusions as shown under transparent light. In addition, there are several sulfide grains, which are all associated with fluid inclusions (Sun et al., 2004a). The unique composition of sulfides in Manus glasses (low Ni, Pt, etc.) may be plausibly interpreted by sulfides crystallized from magmatic fluids.

Adakite was initially named for rocks formed through partial melting of subducted young oceanic crust (b25 Ma, represented by midocean ridge basalt, MORB) (Defant and Drummond, 1990; Kay, 1978). In contrast to most other rock types, adakite is defined by geochemical compositions (e.g., SiO2 ≥ 56 wt.%, Al2O3 ≥ 15 wt.%, Y ≤ 18 ppm, Yb ≤ 1.9 ppm and Sr ≥ 400 ppm) without detailed petrographic constraints. Therefore, (1) both eruptive and intrusive rocks can be classified as adakites; and (2) adakites may be produced simply by partial melting of mafic rocks in the presence of garnet and absence of plagioclase. Different mechanisms have been proposed to produce adakites, e.g., partial melting of the lower continental crust (Chung et al., 2003; Gao et al., 2004; Guo et al., 2006; Xu et al., 2002, 2006; Zhang et al., 2001b) or underplated new crust (Hou et al., 2009; Martin, 1999), or by fractional crystallization of normal arc magmas (Castillo, 2006; Macpherson et al., 2006; Richards and Kerrich, 2007). Given that the oceanic crust is very different from continental crust, slab melts can be distinguished from lower continental crust melts using geochemical criteria (Ling et al., 2011; Liu et al., 2010; Sun et al., 2012a). 5.3.1. Adakite and porphyry Cu deposits Most porphyry Cu deposits are associated with adakites (Oyarzun et al., 2001; Sajona and Maury, 1998; Sun et al., 2011, 2012a,b, 2013a, b; Thieblemont et al., 1997), but the association is not true vice versa. Many adakites, e.g., those from the Dabie Mountains, are not mineralized at all (Huang et al., 2008; Ling et al., 2013; Liu et al., 2012; Wang et al., 2007a). Partial melting of thickened eclogitic lower continental crust (Wang et al., 2006a,b, 2007a,b; Zhang et al., 2001a) and fractional crystallization of garnet (Macpherson et al., 2006) or amphibole (Richards and Kerrich, 2007) may also form high Sr/Y magmas. The lower continental crust melts have much lower Cu abundance and lower oxygen fugacity than subducting slabs, such that they are not favorable for forming porphyry Cu deposits. Fractional crystallization of garnet and/or amphibole plays no positive role in Cu mineralization, either (Sun et al., 2011, 2012a). Adakite formation through slab melting, on the other hand, does favor porphyry Cu mineralization (Mungall, 2002; Sajona and Maury, 1998; Sun et al., 2011, 2012a; Thieblemont et al., 1997). Several different explanations for the fertility of adakitic slab melts have been proposed: Oxidized. It has been argued that slab melts might be unusually oxidized and rich in sulfur (Oyarzun et al., 2001), due to high Fe3+ content from oxidative sea-floor alteration. As a consequence, this elevated fO2 state causes oxidation of chalcophile metal-bearing sulfide phases in the mantle wedge, releasing metals to the silicate melt phase (Mungall, 2002). However, although oxidation is crucial to porphyry mineralization, adakites are not systematically more oxidized than normal arc magmas (Ballhaus, 1993; X.M. Sun et al., 2007; Sun et al., 2013a,b). Normal arc melts are not systematically more enriched in Cu than MORB, either (Lee et al., 2012). Water: It has been proposed that slab melts might be unusually water rich (Sajona and Maury, 1998). High water contents in magmas may suppress the crystallization of plagioclase, and promote the formation of amphibole, resulting in high Sr/Y signatures (Richards, 2011a, 2012). The problem with this model is that no studies have ever demonstrated that adakites are more hydrous than normal arc rocks. Moreover, the crystallization of amphibole plays no role in Cu mineralization. Copper is incompatible in most major silicate minerals, but could be compatible in amphibole, depending on the composition of the magmas and amphibole (Sun

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Table 2 Summary of sulfide/silicate partition coefficients. Cu

Au

Ag

Ni

Pt

Re

Mo

References

212–278 1383 250–970 330–1070 1040–1624 1050–1850

– 1.5E+4–1.9E+4 – 7E+3–1.3E+4 – 4.1E+03–1.12E+04

– – – 300–1680 853–1528 –

240–308 500–900 540–4400 210–1270 646–965 –

– – – – – 2.41E+05–4.19E+05

– – – – – 361–960

– – – 0.20–14.48 – –

Rajamani and Naldrett (1978) Peach et al. (1990) Gaetani and Grove (1997) Li and Audetat (2012, 2013) Patten et al. (2013) Mungall and Brenan (2014)

et al., 2012b, 2014a). Therefore, none of the arguments concerning water abundance can be plausibly interpreted to explain an association between adakites and porphyry deposits. Because of these uncertainties, it has been asserted that there is no single, obvious reason why slab melts have an unusually high potential to form porphyry deposits (Richards, 2013). Felsic: Adakite is formed by partial melting of basaltic rocks, such that they should logically be more felsic than peridotite melts (Sajona and Maury, 1998). It has been further proposed that adakites might more readily crystallize as intrusive plutons because of their viscous felsic nature, leading to the generation of a more efficient crustal magmatic–hydrothermal systems (Sajona and Maury, 1998). In fact, adakites associated with porphyry Cu deposits are mostly intermediate in composition, not felsic. Compression. Adakites are often generated by flat subduction of young oceanic crust with associated compressional stress in the upper plate. Such an environment should be favorable for trapping

Fig. 26. Diagrams of SiO2 versus Cu and Re. Note, in contrast to Cu, Re keeps increasing when Cu drops suddenly at the point when magnetite starts to crystallize. Rhenium is much less chalcophile than Cu under high oxygen fugacities. It is present mainly in the form of Re6+ and acts as an incompatible element before magnetite crystallization and was gradually reduced to Re4+ during magnetite crystallization. Modified after Sun et al. (2003a, 2004a).

magma in a non-erupting, closed-system pluton where sulfur might precipitate as hydrothermal sulfides and sulfates instead of being degassed as SO2 (Oyarzun et al., 2001, 2002). In addition, compression also results in uplifting and erosions, which are later favorable for the exposure of porphyry deposits. The question is again why normal arc rocks in compressed environments do not form porphyry deposits. High Cu contents. Oceanic crust has a much higher Cu abundance (~ 100 ppm) (Sun et al., 2003a,b) than the mantle (30 ppm) (McDonough and Sun, 1995) or the continental crust (~ 27 ppm) (Rudnick and Gao, 2003). It has been proposed that partial melting of the subducted oceanic crust forms adakites with systematically higher Cu initial contents—favorable for porphyry Cu mineralization (Sun et al., 2011, 2012a,b). Previous studies, however, did not quantitatively model the influence of oxygen fugacity, although its positive effect has been emphasized (Sun et al., 2011, 2013a,b). The effects of oxygen fugacity on normal arc magmas have been nicely modeled (Fig. 3) (Lee et al., 2012). Normal arc rocks form by partial melting of peridotite from the mantle wedge. The primitive mantle contains ~250 ppm of S (McDonough and Sun, 1995), whereas depleted mantle peridotite contains ~ 150 ppm of S (O'Neill and Mavrogenes, 2002). Partial melting of mantle peridotite can easily eliminate residual sulfide even under reducing conditions, e.g., by ~20% partial melting at ΔFMQ 0 (Fig. 3) (Lee et al., 2012). In contrast, oceanic crust has sulfur abundances of over 1000 ppm (O'Neill and Mavrogenes, 2002), which is about 4 times greater than the primitive mantle value (McDonough and Sun, 1995). As discussed in Section 3, oxygen fugacities higher than ΔFMQ + 2 are of critical importance to eliminate residual sulfides during slab melting. However, at oxygen fugacities higher than ΔFMQ + 2, ~5–10% of partial melting is enough to eliminate residual sulfide from the subducted oceanic crust, forming adakitic melts with high Cu and S contents up to 2000 ppm and percent levels, respectively. Such elevated contents in the magma, of course, are consistent with the high Cu, S contents present in ore bearing porphyries. 5.3.2. Ridge subduction and porphyry Cu deposits More than half of the world porphyry Cu deposits are located along the western coasts of the North and South American continents comprising the eastern Pacific margin. The total resources there are estimated to be N1.8 billion tons, accounting for about 60% of the world's total Cu resource estimation (Mutschler et al., 2010). Twenty out of the world's top 25 giant porphyry Cu deposits are located there. In contrast, there are essentially no porphyry Cu deposits located along the northwestern Pacific margin, e.g., Japan. Many large porphyry Cu–Au deposits are connected to subduction of spreading and aseismic ridges (Fig. 27) (Cooke et al., 2005; Sun et al., 2010). As discussed above, slab melting is usually involved in creating magmas with high Cu contents and adequately high oxygen fugacity, which are the two main controlling factors for porphyry Cu mineralization (Section 3). Subduction of young ridges, both spreading and aseismic, in particular, produces adakites with high oxygen fugacity, making it the best geological process for porphyry Cu deposits (Sun et al., 2010, 2013a,b). There are several subducting ridges along the east Pacific margin, e.g., in Chile and Peru in the South America. These

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Fig. 27. Many large porphyry Cu deposits are closely associated with ridge subduction, because subduction of young ridges is the most favorable geologic process for slab melting in the Phenozoic, forming highly oxidized melt with high initial Cu contents. Modified after Sun et al. (2010).

ridges were mostly younger than 25 Ma when they began to subduct, and are closely associated with large porphyry Cu–Au deposits (Fig. 27). There are also several ridges (most of which are aseismic, i.e., island chains) on the western Pacific plate (Fig. 28). Most of these ridges, however, are older than 100 Ma (W.D. Sun et al., 2007), and are not likely to form adakites. In addition, the oxygen fugacities in Japan, Izu–Bonin– Mariana, and other arcs along the northwestern Pacific margins, are systematically lower than ΔFMQ + 2 (Fig. 16a). For these reasons, we find it not surprising that no economically viable porphyry deposits associated with volcanic arcs are known in the northwestern Pacific region. There are porphyry ore deposits located at the southwestern Pacific margin, which, however, are much less developed in terms of tonnage and the number of deposits. It is noteworthy that these deposits are associated with the closure of backarc basins younger than 25 Ma (Fig. 29). Considering the geotherm's concave downward shape of subducting slabs, ridge subduction is the most favorable tectonic setting for slab

melting in the Phanerozoic. Geochemical signatures of ridge subduction are important exploration targets for large porphyry Cu–Au deposits. 5.4. Alterations Porphyry deposits have very well developed alteration zones that typically affect several cubic kilometers of rock (Lowell and Guilbert, 1970; Sillitoe, 2010; Titley, 1981) (Fig. 30), which is of critical importance to understanding porphyry mineralization processes and improving their exploration. Alteration is mainly controlled by pH values of the ore-forming fluids (Sillitoe, 2010). The amount of H+ released during mineralization (e.g., Eqs. (15), (16), (18), (19), (22), (23), (25)) together with alkali contents in the porphyry together control advanced argillic lithocap formation and alteration (Sillitoe, 2010). For example, sericitic and advanced argillic alteration are much less well developed in porphyry Cu deposits associated with alkaline than with calc-alkaline intrusions

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Fig. 28. Distributions of aseismic ridges (island chains) in the Western Pacific. These aseismic ridges are all much older than 25 Ma, such that do not get melted during subduction. Modified after W.D. Sun et al. (2007).

Fig. 29. Ages of backarc basins in the southwestern Pacific. Subduction of young backarc basin crust forms adakite, which is favorable for porphyry mineralization.

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(Lang et al., 1995; Sillitoe and Burrows, 2002), reflecting control of the K+/H+ ratio by magma chemistry (Sillitoe, 2010). As summarized recently (Sun et al., 2014b), the main alteration reactions are Eqs. (32)–(37):

125

þ

þ

3KAlSi3 O8 þ2H ¼ KAl2 Si3 AlO10 ðOHÞ2 þ2K þ 6SiO2

Potassium feldspar

ð33Þ

sericite þ

þ

þ

3NaAlSi3 O8 þK þ 2H ¼ KAl2 ½AlSi3 O10 ðOHÞ2 þ6SiO2 þ 3Na Sodiumfeldspar

þ

2KðMg; FeÞ3 AlSi3 O10 ðOHÞ2 þ4H Biotite

2þ

¼ AlðMg; FeÞ5 AlSi3 O10 ðOHÞ8 þðMg; FeÞ chlorite

þ

þ 2K þ 3SiO2

ð32Þ

þ

þ

2KAl3 Si3 O10 ðOHÞ2 þ2H þ 3H2 O ¼ 3Al2 Si2 O5 ðOHÞ4 þ2K Sericite

ð34Þ

sericite

kaolinite

Fig. 30. Alteration patterns of porphyry deposits. After: (A) Lowell and Guilbert (1970); (B) Sillitoe (2010) and (C), Richards (2011b).

ð35Þ

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Fig. 30 (continued).

KFe3 AlSi3 O10 ðOHÞ2 þ1=2O2 ¼ Biotite

KAlSi3 O8 potassium feldspar

þFe3 O4 þ H2 O

CaAl2 Si2 O8 þ2KCl þ 4SiO2 ¼ 2KAlSi3 O8 þCaCl2 : Anorthite

ð36Þ

ð37Þ

potassium feldspar

The alteration zone in porphyry Cu deposits starts from barren, early sodic–calcic upward through potentially ore-grade potassic, chlorite– sericite, and sericitic, to advanced argillic, and finally the lithocap (Sillitoe, 2010). In general, the alteration–mineralization zones become progressively younger upward (Fig. 30), consequently the shallower alteration–mineralization zones overprint deeper ones. Sodic–calcic alteration is typically sulfide and metal poor (except for Fe as magnetite) but can host mineralization in Au-rich porphyry Cu deposits. It is commonly located in the immediate wallrocks of the porphyry intrusion, or are found as a centrally located zone of some porphyry Cu stocks (Sillitoe, 2010). Potassic alteration is located in the center and deeper portions of the porphyry. Dominant mineral changes from biotite in relatively mafic porphyry intrusions and host rocks, to K-feldspar in more felsic, granodioritic to quartz monzonitic settings (Sillitoe, 2010). Quartz-K ± Nafeldspar overprints may destroy the more typical potassic assemblages. The chalcopyrite ± bornite ore in many porphyry Cu deposits is largely confined to potassic zones. Potassic-altered wallrocks may attain N 1 km thickness. The potassic alteration generally becomes less intense from the older to younger porphyry phases (Sillitoe, 2010). This is likely due to a lowering of the pH as mineralization continues. Chlorite–sericite alteration produces pale-green rocks and is widespread in the shallower parts of some porphyry Cu deposits, overprinting preexisting potassic assemblages. The alteration is characterized by transformation of mafic minerals to chlorite, plagioclase to

sericite (fine-grained muscovite) and/or illite, and magnetite to hematite (martite and/or specularite), along with deposition of pyrite and chalcopyrite (Sillitoe, 2010). Sericitic alteration in porphyry Cu deposits normally overprints the potassic and chlorite–sericite assemblages. It may be subdivided into two different types—a less common, early greenish to greenish-gray in color alteration, and a far more common white alteration. The sericitic alteration is commonly pyrite dominated, implying effective removal of the Cu (±Au) present in the former chlorite–sericite and/or potassic assemblages. It may also constitute ore with Cu either in the form of chalcopyrite or as high sulfidation-state assemblages (Sillitoe, 2010). The lower portion of argillic lithocaps may overprint the upper parts of the porphyry Cu deposits, whereas the sericitic alteration transforms upwardly to quartz–pyrophyllite. The advanced argillic alteration preferentially affects lithologic units with low acid-buffering capacities (Sillitoe, 2010). 6. Conclusion The key processes of porphyry mineralization are oxidation and reduction of sulfur. Most of the porphyry deposits are closely associated with oxidized magmas, with sulfate as the dominant sulfur mineral species. Sulfur is one of the most important geosolvents that controls the behaviors of copper and other chalcophile elements, therefore knowledge of its geochemical behavior is essential to the understanding of mineralization processes for copper and a variety of other metal resources. Given that for most chalcophile elements, the partition coefficient between sulfide and melt is very high, elimination of residual sulfide is essential for the extraction of chalcophile elements from the source and thus the formation of porphyry deposits. The solubility of sulfur depends strongly on sulfur speciation, which in turn depends on oxygen fugacities. Sulfate is over 10 times more soluble than sulfide.

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At oxygen fugacities higher than ΔFMQ + 2, most of the sulfur in melts is present as sulfate, such that the solubility of sulfur increases from ~1300 ppm to 2 wt.%. Consequently, slab melts becomes sulfur undersaturated at N5% partial melting, with Cu contents of over 1000 ppm. Therefore, ΔFMQ + 2 is often considered the magic number for porphyry mineralization. Metals of porphyry Cu deposits are hosted in sulfides, which require reduction of sulfate to sulfide during the final stage of mineralization. In principle, sulfate can also be reduced by assimilation of reducing sediments or degassing of oxidizing gases. Porphyry deposits are usually mineralized throughout the whole pluton, whereas interactions with country rocks occur mainly at the interface, such that assimilation of reducing sediments is not likely to be the controlling process. Degassing is not the main process in deeply emplaced porphyry bodies, such that it is not likely to be a major process for sulfate reduction, either. Ferrous iron is the most important reductant that is responsible for sulfate reduction during porphyry mineralization. The highest oxygen fugacity favorable for porphyry mineralization is the HM buffer. Otherwise, there is no ferrous Fe in the system. The reduction of sulfate and oxidation of ferrous Fe lower the pH value. This, in turn, elevates the oxidation potential of sulfate, driving the oxygen fugacities up to the HM buffer. These behaviors explain the popular occurrence of hypogene magnetite and hematite (and specularite) in porphyry deposits. Low pH fluids cause formation of pervasive alteration zones in porphyry Cu deposits, starting from barren, early sodic–calcic alteration upward through potentially ore-grade potassic, chlorite–sericite, and sericitic alterations, to advanced argillic alterations, and finally forming the lithocaps. The amount of H+ released during mineralization and the alkali content in the porphyry together control advanced argillic lithocap formation and alterations. Hydrogen and methane form during the final mineralization process of porphyry deposits. Most of the hydrogen and methane should have been oxidized by ferric Fe. In special cases, some of the reduced gases may escape from the system, and even get trapped in fluid inclusions. Therefore, small amount of reduced gases in fluid inclusions cannot argue against the oxidized feature of the magmas. Reduced magmas are not favorable for porphyry mineralization. There are indeed several small porphyry deposits that appear to be related to reduced porphyries. In our opinion, the reduced porphyry deposits so far reported are either host rocks at distal sites as far as several kilometers away from or above the causative porphyry lying deep underneath, or a consequence of fluid transportation. Some of the reduced porphyries were originally associated with oxidized magmas but were reduced through assimilation of reducing components during emplacement. Degassing of SO2 is not likely to be a main process for porphyry deposit formation, and it doesn't reduce sulfate to sulfide, either. Therefore, degassing of SO2 cannot explain the features of reduced porphyries. Given that a large portion of metals in a porphyry deposit are hosted in wall rocks, attention is needed to identify the causative porphyries.

Acknowledgments We would like to thank Professor Cin-ty A. Lee for constructive discussions. This study is supported by the National Natural Science Foundation of China (nos. 41090374, 41121002, 41172080). This is contribution no. IS-1939 from GIGCAS.

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