The genesis of the Hashitu porphyry molybdenum

Miner Deposita DOI 10.1007/s00126-017-0745-5

ARTICLE

The genesis of the Hashitu porphyry molybdenum deposit, Inner Mongolia, NE China: constraints from mineralogical, fluid inclusion, and multiple isotope (H, O, S, Mo, Pb) studies Degao Zhai 1 & Jiajun Liu 1 & Stylianos Tombros 2 & Anthony E. Williams-Jones 3

Received: 27 October 2016 / Accepted: 14 May 2017 # Springer-Verlag Berlin Heidelberg 2017

Abstract The Hashitu porphyry molybdenum deposit is located in the Great Hinggan Range Cu-Mo-Pb-Zn-Ag polymetallic metallogenic province of NE China, in which the Mo-bearing quartz veins are hosted in approximately coeval granites and porphyries. The deposit contains more than 100 Mt of ore with an average grade of 0.13 wt.% Mo. This well-preserved magmatic-hydrothermal system provides an excellent opportunity to determine the source of the molybdenum, the evolution of the hydrothermal fluids and the controls on molybdenite precipitation in a potentially important but poorly understood metallogenic province. Studies of fluid inclusions hosted in quartz veins demonstrate that the Hashitu hydrothermal system evolved to progressively lower pressure and temperature. Mineralogical and fluid inclusion analyses and physicochemical calculations suggest that molybdenite deposition occurred at a temperature of 285 to 325 °C, a pressure from 80 to 230 bars, a pH from 3.5 to 5.6, and a Δlog fO2 (HM) of −3.0, respectively. Results of multiple isotope (O, H, S, Mo, and Pb) analyses are consistent in indicating a genetic relationship between the ore-forming fluids, metals, and the Editorial handling: S.-Y. Jiang Electronic supplementary material The online version of this article (doi:10.1007/s00126-017-0745-5) contains supplementary material, which is available to authorized users. * Degao Zhai [email protected]

1

State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences, Beijing 100083, China

2

Department of Geology, University of Patras, Rion, 26500 Patras, Greece

3

Department of Earth and Planetary Sciences, McGill University, Quebec 3450, Canada

Mesozoic granitic magmatism (i.e., δ18OH2O from +1.9 to +9.7‰, δDH2O from −106 to −87‰, δ34SH2S from +0.3 to +3.9‰, δ98/95Mo from 0 to +0.37‰, 206Pb/ 204Pb from 18.2579 to 18.8958, 207Pb/204Pb from 15.5384 to 15.5783, and 208Pb/204Pb from 38.0984 to 42.9744). Molybdenite deposition is interpreted to have occurred from a low-density magmatic-hydrothermal fluid in response to decreases in temperature, pressure, and fO2. Keywords Fluid inclusions . Mo isotopes . Porphyry Mo deposits . Hashitu . Northeast China

Introduction Economic Climax-type porphyry molybdenum deposits are relatively uncommon and are largely restricted to the Climax, Henderson, Mount Emmons, Silver Cliff, Pine Grove, Questa, and Mount Hope deposits in the USA and the Nordli deposit in Norway (White et al. 1981). They are invariably emplaced in post-subduction extensional settings, and the associated intrusions are commonly A-type granites. Usually, the molybdenum mineralization occurs in the form of quartz-molybdenite stockworks located within or proximal to F-, Rb-, Nb-, and Ta-enriched A-type granitic cupolas (Ludington and Plumlee 2009). Economic Climax-type deposits typically contain between 100 and 1000 Mt of ore grading between 0.1 and 0.3 wt.% Mo; F > 2000, Rb > 250, Nb > 20, Ta > 2, Sr < 100, and Zr < 120 ppm (White et al. 1981). They are enriched in Be, Cs, Li, Sn, Th, and W, but depleted in Cu. However, numerous aspects of their genesis, including the source of metals, the processes responsible for molybdenum ore formation, and the tectonic settings for their emplacement, are still debated (Richards 2003; Seedorff and Einaudi 2004; Audétat 2015).

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The Hashitu Mo deposit, which is located in the southern segment of the Great Hinggan Range and easternmost part of the Central Asian Orogenic Belt in NE China, is the first potential BClimax-type^ deposit discovered in NE China. It is hosted by the Late Jurassic Hashitu granitic pluton and contains more than 100 Mt of Mo ore, with an average grade of 0.13 wt.% Mo. The ore bodies are mainly in the form of quartz-molybdenite veins and stockworks that are concentrated primarily in the apical parts of the pluton. There have been few previous studies of the deposit, and these have focused mainly on the ages of the mineralization and host granite, and the nature and tectonic setting of the granite. The molybdenite has been dated, using the Re-Os method, at 147 ± 1 to 148 ± 1 Ma, which is very similar to the age of the granite (zircon U-Pb ages, 143 ± 2 to 147 ± 1 Ma; Zhang et al. 2012; Zhai et al. 2014a). Results of whole-rock major and trace element analyses and Sr-Nd-Pb isotopic analyses indicate that the intrusion comprises A-type granites, which are inferred to have been emplaced in an intra-continental extensional setting that experienced significant lithosphere thinning (Ding et al. 2016). A study of the deposit published in Chinese reported preliminary sulfur isotope data for molybdenite, pyrite, sphalerite, and pyrrhotite (δ34S values from +0.4 to +3.8‰, n = 6), and results of microthermometric analyses of a small number of melt and fluid inclusions (Zhai et al. 2012). In the current contribution, we report results of a comprehensive study of the Hashitu deposit that makes use of large numbers of fluid/melt inclusion and isotopic analyses (O, H, S, Mo, and Pb) and thermodynamic modeling of phase equilibria to determine fluid and metal sources, and the physicochemical conditions of molybdenite deposition. Using this information, we develop a genetic model for ore formation that calls upon the exsolution of a low-density fluid from a very shallowly emplaced A-type granite and molybdenite mineralization in response to decreases in temperature, pressure, and fO2. The study also examines the similarities and differences between the Hashitu Mo deposit and the wellknown Climax-type Mo deposits, and makes a case for Hashitu being a variant of the latter deposit type. The Great Hinggan Range Metallogenic Belt The Hashitu Mo deposit is located in the Great Hinggan Range (GHR) Metallogenic Belt, which lies in the easternmost part of the Central Asian Orogenic Belt (CAOB; Fig. 1a). The CAOB is rimmed by the Siberian, Tarim, and North China Cratons (SC, TC, and NCC; Fig. 1a). This region is marked by the widespread occurrence of Mesozoic volcanic and intrusive rocks (Fig. 1b, c), including I- and A-type plutons, which were emplaced in four successive stages of geotectonic evolution (Xiao et al. 2004; Wu et al. 2004, 2005b). An early phase of crustal accretion took place in the Neoproterozoic to Paleozoic, and was related to the

subduction of the Paleo-Asian oceanic plate (Li 2006; Wilde 2015). This was followed by uplift in response to collision in the early Mesozoic. During the Late Jurassic, collision was succeeded by subduction of the Paleo-Pacific plate beneath the Eurasian continental plate, in response to thickening of the crust (Li 2006; Wang et al. 2006; Wu et al. 2011a; Wilde and Zhou 2015; Sun et al. 2015). Finally, in the Early Cretaceous, there was large-scale crustal delamination and lithospheric thinning (Wu et al. 2011a; Wilde and Zhou 2015). The I- and A-type granites in the Central Asian Orogenic Belt constitute one of the largest plutonic provinces in the world and one of the most important sites of juvenile crust formation during the Phanerozoic (Wu et al. 2011a). Based on their ages, these granites have been divided in two groups. An early group, with ages of 210 to 275 Ma, is composed mainly of calc-alkaline I- and S-type plutons; the latter were products of post-orogenic extension. Available Sr-Nd isotope data suggest that these magmas were derived from asthenospheric mantle and recycled ancient crust (Chen and Jahn 2001). The second group, with ages of 130 to 160 Ma, comprises anorogenic A-type plutons associated with lithospheric thinning. These A-type granites were emplaced within NNE to NE trending extensional fault zones and formed from melts derived from the lower crust (Wei et al. 2008; Ding et al. 2016). The Great Hinggan Range (GHR) hosts a number of porphyry, skarn, epithermal, and hydrothermal vein-type ore deposits (Wu et al. 2011b, c; Zeng et al. 2011; Li et al. 2012a; Zhai and Liu 2014; Zhai et al. 2013, 2014b, c, 2015) and constitutes an important CuMo-Au-Ag-Pb-Zn-Fe metallogenic belt in NE China. Numerous ore deposits, including porphyry Cu-Mo, skarn Fe-Sn, and polymetallic veins, occur in the southern segment of the GHR Metallogenic Belt (Fig. 2a; Zhai et al. 2014b). Recent exploration has revealed that porphyry Cu-Mo and Mo deposits are particularly common in this area, and has led to the discovery of the Aolunhua, Xiaodonggou, Jiguanshan, Chehugou, and Hashitu deposits (Wu et al. 2011b, c; Zeng et al. 2011; Zhai et al. 2014a; Shuetal.2016).Ithasbeenproposedthatthedepositsinthisregion formed during two different metallogenic events (Li et al. 2012b), anearlyeventintheLatePermian(i.e.,256±7to272±3Ma)anda later event in the Cretaceous and Jurassic (i.e., 129 ± 3 to 167 ± 2 Ma). Most of the porphyry Cu-Mo and Mo deposits formed during the second metallogenic event, which was genetically related to the widespread intrusion of A-type granites (Mao et al. 2005; Liu et al. 2004; Wu et al. 2005b, 2011a). Ore deposit geology The Hashitu Mo deposit is located in the center of the southern segment of the GHR Metallogenic Belt in NE China (Fig. 2a). The main lithotype exposed within the ore district is a composite granitoid, which comprises A-type granite, granite porphyry, and NW-trending diorite dikes and sills (Fig. 2b). Field relationships reveal that the granite porphyry intruded the A-type granite as

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Fig. 1 a Tectonic scheme for the Central Asian Orogenic Belt (modified from Safonova 2009). b Simplified map of tectonic units of eastern China showing the distribution of the Mesozoic volcanic rocks (based on Wu et al. 2005a). c Geological map showing Mesozoic granites and tectonic

divisions in the Great Hinggan Range (modified from Qi et al. 2005). CAOB Central Asian Orogenic Belt, NCC North China Craton, SC Siberian Craton, CB Cathaysia Block, TC Tarim Craton, YC Yangtze Craton

younger apophyses (Fig. 2c; Zhai et al. 2014a). The Hashitu granite and porphyry are medium- to coarse-grained and are composed of equal proportions of quartz and K-feldspar as phenocrysts in a groundmass of quartz, plagioclase, K-feldspar, biotite, muscovite, and pyrope, with accessory zircon, fluorite, apatite, titanite, ilmenite, and magnetite. Both units are classified as A2-type granites (Zhai et al. 2014a). They are characterized by high alkali (K2O/Na2O > 1), fluorine (e.g., occurrence of topaz and fluorite) and Al2O3 contents, a variable FeOT content, and a low content of CaO.

(grading between 0.09 and 7.2 wt.% Mo) but there is also disseminated and stockwork mineralization (grading between 0.08 and 0.36 wt.% Mo; Fig. 3a) (Zhai et al. 2014a). The disseminations and stockworks not only commonly occur near the main veins in altered granites but are also observed distal from the vein mineralization. The South sector is the main exploration site and contains more than 60 mineralized veins, 17 of which constitute the Hashitu vein system, contain 90% of the entire Mo resource of the deposit, and occur along the contacts between the granites and the porphyries (Fig. 2b, c; Zhai et al. 2014a). The veins range in thickness from 0.5 to 9 m and have an average thickness of 2 m. They vary in length from 50 to 530 m and their average length is 80 m. The disseminated mineralization occurs mainly in the North sector, particularly in the No 4 orebody; the disseminated orebodies are sheet-like, range in thickness from 0.8 to 8.6 m, and have an average thickness of 3.5 m.

Molybdenum mineralization Five different mineralized sectors make up the Hashitu deposit, namely the South, North, East, and Out I and II sectors. These sectors contain more than 100 quartz-molybdenite veins (Fig. 2c) (Lu et al. 2009). The molybdenite ores are mainly vein-type

Miner Deposita Fig. 2 a Geological map of the southern Great Hinggan Range showing the locations of major ore deposits (modified from Zhai et al. 2014b). b Simplified geological map and c cross section of the Hashitu Mo deposit (based on Lu et al. 2009)

Three distinct sets of quartz veins have been recognized on the basis of crosscutting relationships (Fig. 3a, b). The earliest is a set of pinkish quartz veins, which was followed, in turn, by sets of smoky quartz (Fig. 3a) and greenish quartz veins (Fig. 3b). The ore minerals are either disseminated or occur as irregular masses (up to 30 vol.%) in quartz veins (Fig. 3a). Pinkish quartz veins (with widths up to 5 m) contain macroscopically euhedral molybdenite (locally up to 25 vol.%), coexisting with subordinate chalcopyrite, galena, sphalerite, arsenopyrite, and pyrite (up to 5 vol.%). Fine- to mediumgrained pinkish quartz, K-feldspar, and albite are the main gangue minerals. These veins commonly contain angular granitic fragments. Smoky quartz veins (with widths up to 1 m) contain molybdenite (up to 8 vol.%), pyrite, pyrrhotite,

sphalerite, and chalcopyrite; the gangue minerals are smoky quartz, muscovite, K-feldspar, albite, epidote, chlorite, and fluorite (Fig. 3b–e). The greenish quartz veins, which represent tension-gashes, have widths of up to 1.5 m and display well-developed open-space filling textures. Calcite, muscovite, chlorite, and epidote with subordinate dickite and paragonite, together with minor chalcopyrite and pyrite, generally accompany the quartz (Fig. 3g). Ore mineralogy and paragenesis Over thirty ore and gangue minerals have been identified in three primary paragenetic stages (I to III) and a subsequent supergene stage (IV) (Fig. 4). The early mineralization (stages

Miner Deposita Fig. 3 Alteration and mineralization from the Hashitu Mo deposit. a Stage I quartzmolybdenite vein intruded by stage II quartz-molybdenite vein. b Stage II quartz-molybdenite vein replaced by a stage III quartz vein and a chlorite vein. c Quartz coexisting with large molybdenite crystals in greisen. d Muscovite coexisting with molybdenite and quartz (BSE). e Muscovite coexisting with molybdenite (crossed polarized light). f Muscovite coexisting with pyrite and sphalerite in quartz (BSE). g Sericite and epidote coexisting with pyrite and quartz; these minerals were cut by late quartz veins (plane polarized light). Chl chlorite, Ep epidote, Fl fluorite, Mol molybdenite, Ms muscovite, Py pyrite, Qz quartz, Ser sericite, Sp sphalerite

I and II) is restricted to veins (pink and smoky quartz) and minor stockworks, whereas the late mineralization (stage III) is disseminated and/or occurs as stockworks. Molybdenite comprises more than 90% of the total ore mineral volume, with the remainder comprising minor proportions of pyrite, chalcopyrite, sphalerite, galena, pyrrhotite, and arsenopyrite (Fig. 5a). Stage I is dominated by molybdenite and is accompanied by chalcopyrite, galena, pyrite, arsenopyrite, sphalerite, and stannite (Fig. 5b). The molybdenite forms small to large isolated platy crystals within the quartz matrix and, locally, intergrowths with K-feldspar and albite. In the second paragenetic stage (stage II), which is also the main Mo mineralization stage (approximately 70% of the Mo resource), the ore minerals comprise molybdenite, pyrite, pyrrhotite, sphalerite, chalcopyrite, marcasite, and gustavite (AgPbBi3S6) (Fig. 5c, d). The accompanying gangue minerals are quartz,

muscovite (Fig. 3c–f), K-feldspar, albite, epidote, chlorite, fluorite, and locally rutile (Fig. 5e). The stage II molybdenite normally occurs as massive aggregates (Fig. 3c), isolated plates or finer disseminations. In the third and final hypogene stage (stage III), minor pyrite and chalcopyrite are accompanied by quartz, calcite, sericite (muscovite), dickite, paragonite, chlorite, and epidote (Fig. 3g). The supergene minerals (stage IV) include molybdite (MoO3) and litharge (PbO) (Fig. 5f). Hydrothermal alteration Widespread silicification affected the Hashitu granites and porphyries. Greisen developed as selvages (with widths of 30 to 50 cm) around joints in which plagioclase (andesine), Kfeldspar (orthoclase), biotite, and hornblende were replaced by

Miner Deposita Fig. 4 Paragenetic sequence of ore and alteration minerals in the Hashitu Mo deposit

fine-grained quartz, muscovite and/or paragonite, epidote, chlorite, topaz, and fluorite. In part, the greisens are spatially related to the disseminated molybdenite mineralization (Fig. 3d, e). Potassium feldspar and albite form alteration envelopes around stage I vein-type molybdenite mineralization, which was cut by veins containing stage II mineralization. Intense to moderately intense greisen alteration zones (with widths up to 30 cm) envelop and are superimposed on the potassic alteration, and grade outward into widespread argillic alteration. The greisen alteration assemblage consists mainly of smoky quartz, muscovite, and minor K-feldspar and albite. This type of alteration is pervasive, although primary textures are preserved locally. The outer argillic zone (with widths up to 10 cm) is characterized by the assemblage greenish quartz, epidote, fine-grained muscovite, dickite, paragonite, chlorite, calcite, and minor fluorite (Fig. 3g). It occurs predominately along late fractures in the Hashitu granitoids. Sampling and analytical methods Over sixty samples were collected from the Hashitu Mo deposit. These include pinkish, smoky, and greenish vein quartz and their associated sulfide minerals (e.g., molybdenite, chalcopyrite, and pyrite). Hydrogen isotopic analyses were conducted on fluid inclusionshostedinquartzandoxygenisotopicanalysesongrainsof quartz. The analyses were carried out using a MAT-253 stable

isotope ratio mass spectrometer at the Beijing Research Institute of Uranium Geology, China National Nuclear Corporation (CNNC). Oxygen was released from the quartz using the BrF5 extraction technique and hydrogen was released from fluid inclusionsbythermaldecrepitation.Amassofapproximately2gquartz was heated to between 100° and 200 °C to decrepitate low temperature secondary fluid inclusions and thereby largely eliminate their contribution to the isotope composition. Following the method of Friedman (1953), the samples were then heated above 500 °C and reacted with zinc powder at 410 °C to generate hydrogen. The analytical precision was better than ±0.2 ‰ for δ18Ο and ±2 ‰ for δD. The isotopic ratios are reported in standard δ notation (‰) relative to SMOW for oxygen and hydrogen (Table 1). In a previous study results, we reported sulfur isotope analyses for six samples obtained by hand-picking sulfide grains from crushed hand-specimens (Zhai et al. 2012). By contrast, for the present study, we used a micro-drill equipped with a 0.3-mm carbide bit, in conjunction with an optical microscope, to obtain small powdered samples (molybdenite, pyrite, chalcopyrite, pyrrhotite, and galena) from thick and thin sections. Prior to micro-drilling, all the minerals in the samples were examined with a microscope in reflected light in order to confirm that the grains of the different minerals are in apparent textural equilibrium (Fig. 6b, e). The samples were analyzed using a Finnigan MAT-252 stable isotope ratio mass spectrometer at the Department of Geological Sciences, Indiana

Miner Deposita Fig. 5 Ore mineral assemblages from the Hashitu Mo deposit. a Galena replaced by arsenopyrite and chalcopyrite, which were replaced by pyrite and sphalerite (reflected light). b Stannite coexisting with sphalerite and pyrite (BSE). c Pyrite as subhedral grains in chalcopyrite; sphalerite coexisting with chalcopyrite and gustavite occurring as disseminated grains in chalcopyrite (reflected light). d Coexisting pyrite, chalcopyrite, and sphalerite cut by late marcasite and chalcopyrite veins (reflected light). e Molybdenite as independent crystals in quartz veins and rutile coexisting with magnetite (BSE). f Gustavite coexisting with sphalerite and cutting early chalcopyrite, which coexists with pyrite; all of the minerals are cut by fine veins of molybdite and litharge (BSE). Apy arsenopyrite, Ccp chalcopyrite, Gn galena, Gus gustavite, Lit litharge, Mrc marcasite, Mob molybdite, Mol molybdenite, Mag magnetite, Py pyrite, Rt rutile, Sp sphalerite, Stn stannite

University, Bloomington, USA. The standard V2O5-SO2 method was utilized for the analyses following the procedures outlined by Lefticariu et al. (2006). The sulfur isotopic ratios are reported in standard δ notation ‰ relative to V-CDT. Table 1 Hydrogen and oxygen isotopic compositions of fluid inclusions and quartz in the Hashitu Mo deposit Sample

Stage T (°C) δDH2O (‰) δ18OSMOW (‰) δ18OH2O (‰)

HST-11 HST-04 HST-10 HST-19 HST-20 HST-22

I II II III III III

385 325 325 285 285 290

−98 −91 −87 −99 −106 −103

13.5 15.7 11.8 10.1 9.3 11.8

9.1 9.7 5.8 2.7 1.9 4.5

Temperatures are based on microthermometric results of fluid inclusions hosted in quartz; δ18 OH2O values are calculated based on quartz-water oxygen isotope fractionation equations of Clayton et al. (1972).

Analytical accuracy was better than ±0.05 ‰ and the reproducibility was within ±0.2‰ (±2σ). The sulfur isotopic ratios are listed in Table 2. The molybdenum isotopic composition of molybdenite was determined using a MC-ICP-MS, Neptune®, Thermo Finnigan at BRGM, Orléans, France, following the methods of Arnold et al. (2004) and Pietruszka et al. (2006). In order to maximize precision, all samples were analyzed at least three times in different sequences and the average value was used for subsequent data interpretation. Values of δ98/95Mo and δ97/ 95 Mo were evaluated using the equations of Pietruszka et al. (2006) and are reported as ‰. All the data are reported at the ±2σ level relative to NIST SRM 3134 with an external reproducibility of 0.09‰ and 0.06‰ for δ98/95Mo and δ97/95Mo, respectively. The Mo isotopic data are listed in Table 3. The Pb isotopic compositions of the sulfides were determined at the Chinese Academy of Geological Sciences (CAGS), Beijing, China. The analyses were carried out using a Nu

Miner Deposita Fig. 6 Micro-scale δ34S distribution in ore minerals (stages II and III) from the Hashitu Mo deposit. a Massive texture ore with an assemblage of pyrite, chalcopyrite, and galena. b Galena and chalcopyrite in apparent textural equilibrium (reflected light). c Pyrite and quartz vein crosscutting earlier chalcopyrite and molybdenite mineralization. d Massive texture ore with an assemblage of pyrite and chalcopyrite. e Pyrite and chalcopyrite in apparent textural equilibrium (reflected light). Ccp chalcopyrite, Gn-galena, Mol molybdenite, Py pyrite, Qz quartz

Plasma High Resolution type MC-ICP-MS and the Pb isotope NBS 981 standard. Long-term repeated measurements of lead isotopic ratios of standard NBS 981 yielded a 208Pb/206Pb value of 2.1674 ± 0.0005, a 207Pb/206Pb value of 0.91486 ± 0.00025, a 206 Pb/204Pb value of 16.9397 ± 0.0111, a 207Pb/204Pb value of 15.4974 ± 0.0089, and a 208Pb/204Pb value of 36.7147 ± 0.0262 (±2σ). The lead isotopic data are reported in Table 4. Microthermometric measurements were performed on fluid inclusions in quartz from Mo-bearing quartz veins and quartz phenocrysts from the mineralized granitoids. The microthermometric data were obtained using a LINKAM MDSG600 heating/freezing stage coupled to a ZEISS microscope at the School of Earth Sciences and Resources, China University of Geosciences Beijing (CUGB). The stage enables measurements within the range of −196° to 600 °C. A high temperature heating stage was utilized for melt inclusion measurements. This instrument has a temperature range from 25° to 1200 °C. The microthermometric results are reported in Table 5. Laser Raman spectroscopic analyses of individual inclusions were carried out at the Beijing Research Institute of Uranium Geology, China National Nuclear Corporation (CNNC), using a Renishaw RM-2000 Raman spectroscopic microscope. This instrument records peaks in the range

of 100 to 4000 cm−1 with a resolution of 1 to 2 cm−1. The spot size of the laser beam is about 1 μm. The inclusions were analyzed for CO2, CH4, N2, CO, H2S, SO2, C2H6, NH3, and H2.

Results Isotope geochemistry Hydrogen and oxygen isotope compositions Hydrogen and oxygen isotope compositions were obtained from pinkish, smoky, and greenish vein quartz in textural equilibrium with sulfides. As the hydrogen was not obtained from the OH site in hydrous minerals, the risk of misinterpreting the origin of the mineralizing fluid from the corresponding isotopic ratio was relatively high. The δDH2O values display a narrow range from −106 to −87‰; whereas the δ18Oquartz values display a relatively large range, from +9.3 to +15.7‰ (Fig. 7, Table 1). Fluid trapped in quartz of stage I and II veins is commonly characterized by high δDH2O values (−98 to −87‰) relative to fluid in stage III veins (−106 to −99‰).

Miner Deposita Table 2 Sulfur isotope compositions of sulfides from the Hashitu Mo deposit

Sample

Mineral occurrence/ore type

Mineral

δ34SV-CDT (‰)

T (°C)

δ34SH2S (‰)

HST-09-2 HST-14 HST-15-1 HST-15-2 HST-15-3 HST-15-4 HST-16-1 HST-16-2 HST-16-3 HST-16-4 HST-16-5 HST-16-6 HST-16-7 HST-16-11 HST-16-8 HST-16-9 HST-16-10 HST-16-11 HST-16-12 HST-16-13 HST-16-14 HST-17-1 HST-17-2 HST-17-3 HST-17-4 HST-17-5 HST-17-6 HST-17-7 HST-17-8 HST-18-1 HST-18-2 HST-18-3 HST-18-4 HST-18-5 HST-18-6 HST-18-7 HST-18-8 HST-19-1 HST-19-2 HST-1 HST-2 HST-3 HST-4 HST-5 HST-6

Vein/quartz vein Euhedral/massive Anhedral/massive Euhedral/massive Anhedral/massive Euhedral/massive Cubic/massive Anhedral/massive Anhedral/massive Euhedral/massive Vein/massive Anhedral/massive Anhedral/massive Cubic/massive Cubic/massive Cubic/massive Vein/massive Cubic/massive Cubic/massive Cubic/massive Cubic/massive Cubic/massive Vein/massive Cubic/massive Anhedral/massive Cubic/massive Anhedral/massive Vein/massive Colloidal/massive Anhedral/massive Cubic/massive Colloidal/massive Colloidal/massive Colloidal/massive Anhedral/massive Vein/massive Colloidal/massive Vein/quartz vein Anhedral/quartz Vein/quartz vein Vein/quartz vein Vein/quartz vein Euhedral/massive Vein/quartz vein Anhedral/massive

Mol Py Cp Py Cp Py Py Cp Cp Gn Py Cp Gn Py Py Py Py Py Py Py Py Py Py Py Cp Py Cp Py Py Cp Py Py Py Py Cp Py Py Mol Cp Mol Mol Mol Py Sp Po

1.7 1.1 2.4 2.0 2.7 3.5 2.0 2.8 2.8 2.5 2.3 3.0 2.9 2.4 2.5 2.4 2.3 2.4 2.6 2.5 2.1 2.3 2.0 2.6 2.4 2.7 1.9 2.6 2.7 3.1 2.6 2.5 2.4 2.5 2.3 3.0 2.4 2.1 1.7 2.6 2.3 3.8 0.4 0.9 2.3

325 325 325 325 325 325 325 325 325 325 325 325 325 285 285 285 285 285 285 285 285 285 285 285 285 285 285 285 285 285 285 285 285 285 285 285 285 325 325 325 325 325 325 325 325

1.8 1.0 2.4 1.8 2.7 3.4 1.9 2.8 2.8 2.7 2.2 3.0 3.1 2.3 2.4 2.3 2.1 2.3 2.4 2.4 2.0 2.2 1.9 2.5 2.4 2.5 1.9 2.4 2.6 3.1 2.5 2.4 2.3 2.4 2.3 2.8 2.3 2.1 1.7 2.7 2.4 3.9 0.3 0.9 2.3

Average temperatures were obtained from fluid inclusions for stages II (325 °C) and III (285 °C); δ34 SH2S values were calculated based on the H2S-sulfide equations of Ohmoto and Rye (1979), Ohmoto and Lasaga (1982); Abbreviations: Cp-chalcopyrite, Gn-galena, Mol-molybdenite, Po-pyrrhotite, Py-pyrite, Sp-sphalerite. Data HST1 to 6 are from Zhai et al. (2012), and others are from this study. Data HST-1 to 6 are from Zhai et al. (2012) and the others are from this study Cp chalcopyrite, Gn galena, Mol molybdenite, Po pyrrhotite, Py pyrite, Sp sphalerite

Sulfur isotope compositions The locations of sulfide grains that were analyzed for their sulfur isotopic compositions are shown in Fig. 6a, d and demonstrate clear equilibrium textural relationships (Fig. 6b–e). The 39 δ34SV-CDT values reported in Table 2 and 6 δ34SV-CDT values reported by Zhai et al. (2012) range from +0.4 to +3.8‰ and average +2.4‰ (n = 45; Fig. 8, Table 2). The δ34S values for molybdenite and

chalcopyrite are very similar; for both, the average is 2.5‰ and the ranges are from 1.7 to 3.8‰ (n = 5) and 1.7 to 3.1‰ (n = 10), respectively. These values are comparable to those for pyrite, which range from 2.0 to 3.5‰ (n = 24) and average 2.5‰; they exclude two relatively low values of 0.4 and 1.1‰. Single analyses of pyrrhotite and sphalerite yielded δ34S values of 2.3 and 0.9‰, respectively. Two analyses of galena yielded δ34S values of 2.5 and 2.9‰.

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Molybdenum isotopes

Fluid and melt inclusions

Molybdenum isotope compositions were obtained for stage I and II molybdenite (Table 3). The δ98/95Mo and δ97/95Mo values range from 0 to +0.37 and 0 to +0.26‰, respectively. All of the δ98/95Mo and δ97/95Mo ratios correlate linearly (r2 = 0.995) (Fig. 9), suggesting that isobaric interferences were largely eliminated and appropriate adjustments were applied to correct for analytically induced mass fractionation. In addition, the fact that the variations of δ97/ 95 Mo and δ98/95Mo are very similar indicates that the measured Mo isotopic ratios were almost entirely the product of natural mass-dependent fractionation. This is supported by the observation that intra-sample variations of up to 0.14 and 0.19‰ for δ97/95Mo and δ98/95Mo, respectively, were determined for individual molybdenite crystals within the same hand specimen (e.g., sample HST-20a and b; Table 3).

Fluid inclusions

Lead isotope compositions Lead isotope compositions were determined for molybdenite, pyrite, and pyrrhotite. The measured 206Pb/204Pb, 207 Pb/204Pb, and 208Pb/204Pb ratios are listed in Table 4. In general, molybdenite is characterized by 206Pb/204Pb and 208 Pb/ 204 Pb ratios from 18.3256 to 18.8958 and 38.1808 to 42.9744, respectively, which are higher than the corresponding ratios for pyrite and pyrrhotite. The 206 Pb/204Pb and 208Pb/204Pb ratios for pyrite are 18.2671 and 38.1084, respectively, which are similar to the ratios for pyrrhotite of 18.2579 and 38.0984. Moreover, there is no evident variation in the ratios of different textural types of molybdenite, pyrite, or pyrrhotite (Table 4).

Table 3 Mo isotope compositions of molybdenites from the Hashitu Mo deposit Sample

Ore type

n

δ97/95Mo

±2σ

δ98/95Mo

±2σ

HST-02 HST-03

Vein Vein

3 4

0.14 0.17

0.06 0.08

0.22 0.26

0.09 0.13

HST-07 HST-10 HST-18 HST-19 HST-20a HST-20b HST-22

Vein Massive Vein Massive Vein Vein Vein

4 4 4 4 3 3 4

0.19 0.17 0.00 0.16 0.12 0.26 0.07

0.03 0.07 0.06 0.04 0.06 0.05 0.07

0.27 0.25 0.00 0.24 0.18 0.37 0.10

0.02 0.07 0.06 0.04 0.11 0.03 0.11

All the Mo isotope compositions are obtained from molybdenite and reported as per mil (‰) relative to NIST SRM 3134 n the number of individual measurements for the same sample

Fluid inclusions were analyzed microthermometrically in quartz from the different vein stages (i.e., the pinkish, smoky, and greenish quartz veins) and quartz phenocrysts from granites and porphyries. Only fluid inclusions deemed to be primary from their occurrence in 3-D clusters or along growth zones were analyzed, and each cluster or set of inclusions along a growth zone was considered to represent a separate fluid inclusion assemblage (FIA). The inclusions in the clusters and along growth zones are elliptical, rod-shaped, rounded, irregular, or display negative crystal shapes and range in length from 10 to 30 μm (Fig. 10). Fluid inclusions, which occur along fractures or grain boundaries in clusters and linear arrays, were considered secondary and were not analyzed microthermometrically due to the possibility that they formed very late relative to the mineralization. Fluid inclusions were classified on the basis of the phase relationships observed at room temperature (Fig. 10). Four types of fluid inclusions were recognized: (i) aqueous liquid inclusions (L-V; Fig. 10a), which homogenize to liquid upon heating (they generally contain ~60 to 80 vol.% liquid); (ii) vapor-rich aqueous inclusions (V-L; Fig. 10b), which homogenize to vapor (they typically contain >60 vol.% vapor); (iii) monophase vapor or liquid aqueous inclusions (type V or L; Fig. 10c); and (iv) halite-bearing inclusions (L-V-S; Fig. 10d-g) with 20 to 65 vol.% vapor. Some L-V-S inclusions contain several solids in addition to halite (identified by its cubic shape). However, only halite shows consistent volume ratios with the other phases in the inclusions phase proportions. It is therefore considered to be a daughter mineral (Fig. 10d, i). The other solids, which almost invariably show inconsistent volume ratios with the other phases are interpreted to be accidentally trapped solids. In some cases, LVand V-L inclusions also contain some trapped solids. The following solids were identified in L-V-S, L-V, and V-L inclusions: hematite based on its red color (Fig. 10f), and calcite (Fig. 11f) and rhodochrosite (Fig. 11e) from their Raman spectra. Primary L-V, V-L, and L-V-S inclusions are all commonly observed in stage I and II quartz veins, whereas the L-V type inclusions dominate the stage III quartz veins. Laser Raman analyses of individual fluid inclusions indicate that the vapor consists mainly of H2O but also contains CO2 (Fig. 11a, b). Raman spectra for the liquid indicate that CO2(aq) and SO42− are the principal dissolved aqueous species in addition to Cl−, which was analyzed microthermometrically (Fig. 11c, d). No CO2-rich fluid inclusions were observed, although Raman analyses showed that CO2 is present in both the liquid and the vapor. All inclusion types except the monophase inclusions are commonly observed in the different quartz samples (Fig. 10d, e), though their relative proportions differ. Significantly, V-L and L-V or L-V-S inclusions commonly coexist along the same growth

Miner Deposita Table 4

Pb isotope compositions of sulfides from the Hashitu Mo deposit 206

Pb/204Pb

±2σ

207

Pb/204Pb

±2σ

208

Pb/204Pb

±2σ

Sample

Ore type

Mineral

10HST-02

Vein

Molybdenite

18.5288

0.0007

15.5532

0.0005

38.2127

0.0011

10HST-05

Massive

Molybdenite

18.3256

0.0008

15.5425

0.0005

38.1808

0.0014

10HST-07 10HST-09

Massive Massive

Molybdenite Molybdenite

18.8958 18.8299

0.0006 0.0007

15.5783 15.5686

0.0006 0.0007

38.4548 42.9744

0.0016 0.0021

10HST-15 10HST-16

Massive Massive

Pyrrhotite Pyrite

18.2579 18.2671

0.0010 0.0005

15.5384 15.5408

0.0008 0.0006

38.0984 38.1084

0.0023 0.0017

zones of quartz crystals, providing evidence that the fluid underwent phase separation (Fig. 10d–f, i). Microthermometric measurements were performed on L-V, V-L, and L-V-S inclusions. Most L-V inclusions and the liquid and vapor in L-V-S inclusions homogenized to liquid and V-L inclusions homogenized to vapor. Liquid-vapor homogenization temperatures were determined for 204 fluid inclusions, including 25 FIAs, and final ice melting temperatures were determined for a subset of 22 of these FIAs (Fig. 12, Table 5). Numerous L-V and V-L inclusions with accidentally trapped solids did not homogenize, despite being heated to 600 °C (the upper temperature limit of the heating/freezing stage), because of the failure of one or more of the trapped solids to dissolve; the homogenization temperature for these inclusions is that for the homogenization of liquid and vapor (Table 5). The halite in some L-V-S inclusions dissolved at a temperature below that of liquid-vapor homogenization, whereas in others it dissolved above the latter temperature. In individual FIAs, the ranges of both the halite dissolution temperature and the temperature of liquid-vapor homogenization are relatively large, suggesting that a significant proportion of the inclusions may be the products of the heterogeneous entrapment of phases (Table 5); inclusions that showed evidence of necking-down were excluded. Fluid inclusion assemblages of L-V-S and/or L-V and V-L types in stage I, II, and III veins have mean liquid-vapor homogenization temperatures of 385, 325, and 285 °C, respectively. The final ice melting temperature (Tm ) for different FIAs ranged from −24.0 to −0.4 °C (Table 5), corresponding to a salinity range from 0.7 to 25.0 wt% NaCl equivalent (salinity was estimated from the equations of Brown and Lamb 1989). Halite dissolution temperatures were utilized to calculate the salinity of L-V-S inclusions (calculations were based on Lecumberri-Sanchez et al. 2012). The corresponding salinities range from ~30 (greenish quartz veins) to ~66 wt% NaCl equivalent (greisen porphyry and pinkish quartz veins) (Table 5, Fig. 12). The different FIA types show distinct homogenization temperature and salinity ranges (Fig. 12a), i.e., L-V-S dominant FIAs have relatively high salinity and homogenization temperatures, which overlap part of the temperature range for the L-V dominant and V-L dominant FIAs.

Melt inclusions Melt inclusions ranging up to 20 μm in diameter were identified in quartz phenocrysts from the granites and porphyries and typically contain glass, liquid, and vapor at room temperature (Fig. 10h). Most of the melt inclusions homogenized to a silicate liquid at temperatures varying from ~640 to 890 °C (Fig. 12b). These temperatures are similar to the crystallization temperatures estimated for the Hashitu granitic magmas based on the Ti-in-zircon geothermometer (Watson et al. 2006), of ~583 to 760 °C (Electronic Table 1).

Discussion Temperature-pressure conditions The Hashitu A-type granite and associated porphyries are estimated from the temperatures of homogenization of melt inclusions reported above to have been emplaced at temperatures between ~640 and 890 °C. This estimate is consistent with evidence (also reported above) that zircon in these rocks crystallized between ~583 to 760 °C. The temperature estimates ignore the effect of pressure (Fu et al. 2008), which in the current context is likely to have been small. The temperature of Mo mineralization at Hashitu was estimated using a combination of fluid inclusion and stable isotope geothermometers. Microthermometric measurements of fluid inclusion assemblages hosted in stage I, II, and III quartz veins yielded average FIA homogenization temperatures (Th) of 385 °C, 325 °C and 285 °C, respectively. As L-V or L-V-S inclusions coexisted with V-L inclusions in the same FIAs, the fluid is interpreted to have boiled or condensed and these temperatures represent the temperatures of entrapment. This interpretation is supported by the observation of several FIAs with relatively wide homogenization temperature ranges (Table 5) providing evidence of probable heterogeneous trapping of fluid (Ramboz et al. 1982; Bodnar et al. 1985). Similar temperatures were obtained using the pyrite-chalcopyrite, pyrite-galena, chalcopyrite-galena, and molybdenitechalcopyrite sulfur isotope geothermometers, from 384 to 434 and 304 to 318 °C for stages I and II, respectively

Miner Deposita Table 5

Fluid inclusion microthermometric measurement results from the Hashitu molybdenum deposit Tmhalite (°C)

Salinity (wt% NaCl equiv)

Host mineral

Number

FIAs

HST-02

Pinkish quartz (stage I)

13

L-V and V-L

−3.5 to −0.4

230 to 410

–

0.7 to 17.3

6 22

L-V and L-V-S L-V and V-L

– −22 to −0.5

480 to 600 377 to 596

373 to 553 –

31.4 to 65.5 0.9 to 23.7

HST-03 HST-08

Tmice (°C)

Th (°C)

Sample

HST-21

46

L-V and V-L

−5.4 to −0.5

384 to 600

–

0.9 to 8.9

HST-21-2

9

L-V-S

–

375 to 600

235 to 456

33.7 to 55.2

Smoky quartz (stage II)

4 11

L-V-S and L-V L-V and L-V-S

−11 –

277 to 341 284 to 366

306 to 391 259 to 438

15.0 to 46.2 35.1 to 51.6

11

L-V and V-L

−6.2 to −0.4

286 to 359

–

0.7 to 9.5

Greenish quartz (stage III)

17 10

L-V and V-L V-L and L-V-S

−24 to −0.4 –

210 to 280 205 to 265

– 145 to 391

0.9 to 25.0 29.6 to 46.2

3 6

L-V and L-V-S L-V-S

– –

231 to 260 220 to 271

360 to 400 241 to 438

43.0 to 47.2 34 to 51.6

HST-04 HST-22-3 HST-21-3 HST-21-5 HST-21-6 HST-06 HST-22-4 HST-19

Quartz phenocrysts (granite)

21

L-V-S and L-V

−22 to −1.4

230 to 484

215 to 414

2.4 to 48.9

HST-23 HST-23-3

Quartz phenocrysts (greisen porphyry) Quartz phenocrysts (greisen porphyry)

22 3

L-V and V-L L-V-S

−4.1 –

370 to 535 390 to 513

– 160 to 551

6.6 30.1 to 63.1

HST-24 HST-25

Quartz phenocrysts (porphyry) Quartz phenocrysts (granite)

31 31

M (to L) M (to L)

– –

640 to 830 680 to 890

– –

– –

N the number of inclusions analyzed, FIAs fluid inclusion assemblages, L-Vaqueous liquid and vapor inclusion, V-L vapor-rich aqueous inclusion, L-V-S halite-bearing aqueous liquid-vapor inclusion, L liquid, M melt inclusion, B–^not measured, Th homogenization temperature, Tmice final ice melting temperature, Tmhalite final salt melting temperature

(using Ohmoto and Rye 1979). The pressure during the different mineralizing stages was calculated using the BHOKIE FLINCS_H2O-NACL^ spreadsheet for a boiling fluid system (Steele-MacInnis et al. 2012). These calculations indicate that stage I quartz veins formed at pressures of 300 to 460 bars (350 bars on average), which correspond to depths of 1.5 to 2.1 km (1.7 km on average), assuming lithostatic conditions, or 3 to 4.1 km, assuming hydrostatic conditions. The corresponding pressures for stage II and III quartz veins are from 80 to 230 bars (130 bars on average) and 50 to 90 bars (75 bars on average), respectively. The calculated pressures for the different vein generations are consistent with their spatial and geological occurrence. The majority of stage I quartz veins crosscut the deeper parts of the Hashitu porphyry system, whereas stage II and III veins occur at shallower depths. The variations of pressure from early to late stage veins corresponding to ~1 km in depth were mostly related to large-scale lithospheric extension and relatively rapid crustal uplift during the Mesozoic (Wu et al. 2002; Meng 2003; Wang et al. 2006). Based on data from fluid inclusion microthermometry and stable isotope geothermometry, we conclude that main stage molybdenite deposition (stage II) in the Hashitu deposit occurred at temperatures between 285 and 325 °C and near hydrostatic pressures of 80 to 230 bars. Overall, the data presented above indicate that vein formation in the Hashitu Mo deposit occurred at progressively lower pressure and lower temperature.

Other physicochemical conditions Physicochemical conditions, other than pressure and temperature during the formation of ore and alteration minerals were estimated from stability relationships among sulfides and silicates using the SUPCRT92 database (Johnson et al. 1992) and the data of Holland and Powell (1998) for the temperature and pressure conditions mentioned above. For the purpose of estimating stability relationships among the minerals, all were considered to be ideal solid solutions. From the alteration mineral assemblages, K-feldsparmuscovite-albite (stage II), and muscovite-dickite-paragonite (stage III), the log(αK+/αH+) is estimated to have been 3.64 and 1.76 and the log(αNa+/αH+) 4.54 and 3.43, respectively, assuming temperatures for stages II and III of 325 and 285 °C (Fig. 13). During the transition from greisen to argillic alteration, the ore fluids evolved by decreasing both log(αK+/αH+) and log(αNa+/αH+), which resulted in the destruction of Kfeldspar and the stabilization of dickite in stage III. To constrain fO2 conditions for the main ore stage (stage II), a logfO2–pH diagram was constructed utilizing aqueous and mineral equilibria in the Fe–O–S system, assuming a temperature of 325 °C and a pressure of 130 bars (Fig. 14). The alteration minerals associated with the sulfides in stage II were used to constrain the pH. As formation of the main ore stage veins involved precipitation of pyrite, pyrrhotite, muscovite, and K-feldspar (Fig. 3), we conclude, based on Fig. 14, that

Miner Deposita

Fig. 7 δ18O versus δD isotope compositions of the Hashitu mineralizing fluids. The fields of volcanic vapor, felsic magma, and magmatic water are from Hedenquist and Lowenstern (1994). The metamorphic water box and the kaolinite weathering and meteoric water lines were adopted from Giggenbach (1992) and Hoefs (2009)

the conditions of the main stage of ore deposition were relatively reducing (Δlog fO2 (HM) = −3.0) and weakly acidic (pH = 5.6). The presence of hematite in some fluid inclusions from stage II, however, suggests that redox conditions of metal transport could have been significantly higher than those of deposition, although we cannot rule out the possibility that the presence of hematite in these inclusions is an artifact of H2 diffusion (Mavrogenes and Bodnar 1994).

Source of metallic and hydrothermal components of the ore fluids The calculated δ34SH2S values of the mineralizing fluid interpreted to have been in equilibrium with the sulfide minerals range between +0.3 and +3.9‰ and average 2.3‰

Fig. 8 Histogram of δ34SV-CDT isotope compositions for the Hashitu sulfides

Fig. 9 Relationship between the measured δ98/95Mo and δ97/95Mo ratios for molybdenite from the Hashitu deposit

(Table 2), which is consistent with a magmatic origin for the ore fluid, e.g., the local Hashitu granitic magma. It is also noteworthy that sulfur isotope values varied during the precipitation of individual sulfide crystals, as shown by the variation of δ34SH2S on a millimetric scale in single crystals. For example, a pyrite crystal has δ34S values ranging from 2.1 to 2.6‰ (Fig. 6a). The measured δ98/95Mo and δ97/95Mo compositions of the Hashitu molybdenites are similar to those of typical granites worldwide (Anbar et al. 2001) and suggest that the most likely source of molybdenum was the genetically associated granitic magma. A positive correlation between the molybdenum and sulfur isotope values of the Hashitu molybdenite is evident from plots of δ 34 S V-CDT versus δ 98/95 Mo NIST and δ 97/ 95 MoNIST (Fig. 15). These plots indicate that the heavier Mo isotope is associated with the heavier S isotope (r2 = 0.92; Fig. 15) and suggest reduction of the fluid during molybdenite precipitation; high δ34S values can be explained by a reduction in fO2 (Ohmoto 1972). Such an interpretation is consistent with the observation that Mo in hydrothermal fluids is generally transported in the hexavalent state, but precipitates in the tetravalent state as molybdenite (Williams-Jones and Migdisov 2014). During this reduction, heavy Mo isotopes will have been partitioned preferentially into molybdenite (Tossell 2005; Greber et al. 2014). It should be noted that two samples have values lying off the linear trend referred to above (Fig. 15). We infer that these anomalous values likely indicate disequilibrium during the formation of the corresponding molybdenite, due perhaps to fluid phase separation (Heinrich et al. 1999). In order to further evaluate the metal source, we made use of previously published Pb isotopic data, including Pb isotope ratios from local Permian, Jurassic, and Cretaceous volcanosedimentary rocks and granitoids and Pb isotope data from nearby hydrothermal ore deposits (Zhang et al. 1995; Cai

Miner Deposita Fig. 10 Photomicrographs of fluid inclusions in mineralized quartz veins and melt inclusions in quartz phenocrysts of granites from the Hahistu deposit (taken at room temperature). a Type L-V fluid inclusions containing liquid and vapor as the main phases, and dominated by liquid. b Coexistence of type L-V, and V-L or L-V-S fluid inclusions; V-L inclusions are dominated by vapor and L-V-S inclusions contain halite. c Vapor fluid inclusions. d Primary FIA comprising V-L and L-V-S inclusions in a single quartz growth zone; secondary fluid inclusions occur along healed fractures. e–f L-V, V-L, LV-S, pure gaseous, and liquid fluid inclusions; note the presence of a trapped hematite crystal in a V-L fluid inclusion. g L-V-S fluid inclusions in a quartz-molybdenite vein (Mol). h Melt inclusions in a quartz phenocryst. i L-V-S inclusions containing multiple solids

et al. 2004; Guo et al. 2010; Shao et al. 2010; Zhai et al. 2014b) (Fig. 16). A considerable body of lead isotope data also has been collected for the Linxi A-type pluton, which is located near the Hashitu granite (~20 km from it) and is similar in age and composition to the latter. Based on the distribution of 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb ratios in binary plots, the most likely source of Pb in the Hashitu sulfides is the upper crust. There may also have been a minor mantle contribution (Fig. 16). The Pb isotopic ratios of the Mesozoic granitic rocks are linearly distributed between lower crust and mantle values, implying a mixed source (Fig. 16), which is consistent with previous Sr-Nd isotope studies (Ding et al. 2016). Significantly, the Pb isotopic ratios of the ore minerals (i.e., molybdenite, pyrite, and pyrrhotite) are intermediate between those of the Mesozoic granites, with which the Mo deposit is closely associated, and the Jurassic volcanic

rocks, which underlie both the Mo deposit and the intrusion. This suggests that there were contributions of Pb and possibly other metals from both sources. Two molybdenite samples have relatively high Pb isotope ratios, which displaces them from the likely sources and potential mixing line (Fig. 16). This may reflect higher proportions of U and Th, decay of which would have produced higher proportions of radiogenic Pb. Indeed, the relatively high concentrations of U (9.8~27.2 ppm) and Th (15.5~17.3 ppm) in the two molybdenite samples compared to those in other molybdenite samples (i.e., U with 1.1~5.7 ppm and Th with 4.2~9.7 ppm) supports this interpretation (Electronic Table 2). Information on the source of the fluids is provided by the oxygen and hydrogen isotopic data. The values of δ18OH2O calculated from quartz and the interpreted trapping temperatures of the fluid using Clayton et al. (1972) range

Miner Deposita

Fig. 11 Laser Raman spectra for fluid inclusions from the Hashitu Mo deposit. a–b CO2-bearing vapor phase. c CO2-bearing liquid. d Sulfate-bearing liquid. e Trapped rhodocrosite. f Trapped calcite

from +1.9 to +9.7‰. By contrast, the measured δDH2O values have a relatively narrow range, from −106 to −87‰ (Table 1). The δ18OH2O values for stage I and II veins (+5.8 to +9.7‰) are similar to those of magmatic water (+7 to +9‰; Hedenquist and Lowenstern 1994), whereas the values for stage III are significantly lower (+1.9 to +4.5‰). The δDH2O values for fluids from all three stages plot below the field for typical magmatic water (Fig. 7). We speculate that this is because the bulk fluid inclusion decrepitation method employed to provide the water for δDH2O analyses inevitably led to the sampling of a mixture of primary and secondary fluid inclusions presumably of magmatic and meteoric origin, respectively. Accordingly, we conclude that the stage I and stage II

fluids were dominantly of magmatic origin, whereas the stage III fluids contained a significant proportion of meteoric water. Fluid evolution The Hashitu Mo deposit is genetically related to an A-type granite, which exsolved a magmatic-hydrothermal fluid that was responsible for Mo mineralization and hydrothermally altered the rock. The temperature during the early porphyry stage was between ~640 and 890 °C, based on the obtained melt inclusion temperature data and the Ti-in-zircon geothermometer of Watson et al. (2006). Pressure at this stage is interpreted to have been in the range of 300 to 460 bars

Miner Deposita Fig. 12 Microthermometric data for fluid inclusion assemblages (a) and melt inclusions (b) from the Hashitu Mo deposit. The shaded regions reflect the entire range of the measured temperature and salinity data for individual FIAs. Note the gap and overlap between different FIAs. The halite saturation and critical curves are after Driesner and Heinrich (2007)

based on the homogenization temperatures of fluid inclusion assemblages that show evidence liquid-vapor separation. These pressures correspond to depths of 1.5 to 2.1 km. At these depths and pressures, the exsolved fluids were initially of low to intermediate salinity but evolved to an assemblage of high salinity L-V-S and/or L-V fluids and low salinity V-L fluids through liquid-vapor separation induced by a transition from lithostatic to hydrostatic pressure (Fig. 12a). Temperatures for the magmatic-hydrothermal fluids at stage I and II were ~385 and ~325 °C, respectively, but decreased to ~285 °C during cooling of the system.

Although there has been some debate over the capacity of low-density fluids to transport metals (Williams-Jones and Heinrich 2005), experimental studies of Migdisov and Williams-Jones (2013), Migdisov et al. (2014), and Hurtig and Williams-Jones (2014a, b) have shown that Ag, Cu, Au, and Mo can all be transported in appreciable concentrations by vapor. In the case of Mo, an oxidizing vapor-like fluid (low concentrations of reduced sulfur species) of the density

Metal transport and deposition The fluid inclusion and isotopic data presented earlier demonstrate that the fluid responsible for Mo mineralization at Hashitu was exsolved from a magma, which was also the likely source of the metals. Given the relatively shallow level of emplacement of the hydrothermal system (between 1.5 and 2.1 km), the exsolved fluid was likely of low to intermediate density (Williams-Jones and Migdisov 2014). In order to determine its probable density prior to phase separation, we assumed that the fluid had a salinity of ~10 wt% NaCl equivalent (Burnham 1979), a temperature of ~750 °C (the likely emplacement temperature of the magma) and a pressure of ~500 bars (lithostatic pressure). Based on these data, the fluid had a density of ~0.17 g/cm−3, indicating that it exsolved as a vapor-like supercritical fluid and, on cooling, condensed a high salinity liquid. This conclusion is consistent with conclusions of other studies that some porphyry systems exsolve low-density supercritical fluids (Henley and McNabb 1978; Williams-Jones and Heinrich 2005; Landtwing et al. 2010; Hurtig and Williams-Jones 2015).

Fig. 13 Log(αNa+/αH+) versus log(αK+/αH+) phase diagrams showing stability relationships among alteration minerals in the Na2O–K2OAl 2 O 3 -SiO 2 –H 2 O system. The squares in a and b represent approximate stability fields for stages II and III, respectively. The temperature and pressure for which the diagrams were constructed were constrained by fluid inclusion microthermometry. The diagrams were prepared for conditions of quartz saturation and the thermodynamic data taken from the SUPCRT92 database (Johnson et al. 1992) and Holland and Powell (1998)

Miner Deposita

Fig. 14 LogfO2–pH diagram showing stability relationships in the system Fe–Cu–S–O–H system for stage II at T = 325 °C and P = 130 bars. The dickite–muscovite and K–feldspar–muscovite stability boundaries are also shown. Dotted lines (red) separate fields of predominance of aqueous sulfur species and heavy solid lines separate the stability fields of the iron minerals (black) and copper minerals (blue). Gray dotted lines separate the stability fields of the alteration minerals. The shaded area identifies the stability field for the assemblage chalcopyrite-pyrite, and the red square defines the approximate conditions for stage II molybdenite deposition. The diagram was constructed assuming m ∑ S = 0.01 and m K + = 0.01. All the thermodynamic data were obtained from the SUPCRT92 database (Johnson et al. 1992). Hem hematite, Mag magnetite, Ccp chalcopyrite, Dck dickite, Py pyrite, Po pyrrhotite, Bn bornite, Kfs K–feldspar, Ms muscovite

inferred above would be capable of transporting 100 s of ppm Mo at near magmatic conditions as the Mo(VI) species (MoO3(H2O)y), i.e., more than sufficient to form an ore deposit (Hurtig and Williams-Jones 2015). Furthermore, the modeling conducted in the latter paper and in Hurtig and Williams-Jones (2014b) showed convincingly that molybdenum solubility decreases with decreasing temperature and particularly decreasing oxygen fugacity, which ultimately reduces Mo(VI) to Mo(IV) and sulfate to H2S, thereby causing deposition of molybdenite (Williams-Jones and Migdisov 2014). This behavior of molybdenum is consistent with the fluid evolution of the Hashitu system described above, in which temperature decreased from the early to late stage and fO2 was relatively high prior to molybdenite deposition. It is, however, important to recall that the Mo mineralization coincided with a decrease in pressure, which likely caused condensation of liquid, a process that would have sharply reduced ligand activity, hydration, and temperature, thereby promoting deposition of molybdenite (Seward et al. 2014; WilliamsJones and Migdisov 2014). Accordingly, we propose that shallow emplacement of an A-type granite, exsolution of a low-density magmatic ore fluid and a combination of

Fig. 15 Plots of δ34SV-CDT versus δ98/95MoNIST (a) and δ97/95MoNIST (b) for molybdenite from the Hashitu Mo deposit. Note that two data points for the samples with red square symbols deviate from the linear trend of the data from the other samples

decreasing temperature, fO2, and pressure all contributed to ore formation at Hashitu.

Comparisons to Climax-type Mo deposits The Hashitu Mo deposit displays numerous features that are common to those of typical Climax-type Mo deposits. It is genetically associated with A-type granites enriched in fluorine, rubidium (>500 ppm), niobium (>50 ppm), and tantalum. Furthermore, the isotope data presented here suggest that the metals and sulfur were derived largely from the host granite, which is also the conclusion reached by numerous isotope studies of Climax-type deposits (Stein and Hannah 1985; Stein 1988). Finally, the genetic model accepted for Climaxtype Mo deposits, namely that they are the products of fluids exsolved from silicate melts in the upper crust (White et al. 1981; Cline and Bodnar 1994; Klemm et al. 2008; Ludington and Plumlee 2009), is broadly the same as proposed in this paper for the Hashitu deposit.

Miner Deposita Fig. 16 Plots of 206Pb/204Pb versus 207Pb/204Pb (a) and 206 Pb/204Pb versus 208Pb/204Pb (b) for the Hashitu sulfides and local granites and andesites. The lead isotope data for sulfides from the deposit, the granites, and volcanic rocks were taken from Zhai et al. (2014b) and references therein. The Pb isotope curves for the mantle, orogene, and crust are from Zartman and Doe (1981)

There are several notable differences between the Hashitu deposit and Climax-type deposits. One of these is the mineralization style. Climax deposits are characterized by ore shells of quartz-molybdenite stockworks and/or veinlets (centimeter scale) that lie above and surround the greisen-altered apices of highly evolved granites. Although there are some larger veins, and locally breccias, most of the ore occurs as stockworks and veinlets. In contrast, the Mo mineralization of the Hashitu deposit is dominated by relatively large quartz-molybdenite veins, usually several meters thick, that occur along the contacts between the coeval granites and porphyries. This may reflect the very high level of emplacement of the Hashitu intrusions compared to those of Climax-type deposits. Another difference is the nature of the alteration. In Climax Mo deposits, the quartz-molybdenite veins are associated with potassic alteration (Ludington and Plumlee 2009), whereas in the Hashitu Mo deposit, they are intimately associated with greisen; in Climax-type deposits, as noted above, the Mo mineralization lies above or surrounds the greisen alteration. Finally, all except one of the important Climax-type porphyry molybdenum deposits occur in post-subduction and extensional tectonic settings in western North America (Ludington and Plumlee 2009), whereas the newly discovered Hashitu Mo deposit in NE China formed in a distal back arc, syn-subduction setting (Zhai et al. 2014a). In view of these differences, we propose that the Mo deposits in NE China exemplified by Hashitu represent a variant of the Climax-type deposits.

Conclusions Geological, mineralogical, fluid inclusion, and multiple isotope investigations of the newly discovered Hashitu Mo deposit provide compelling evidence that both ore-forming fluids and metals were genetically related to a felsic magma, which was emplaced at shallow crustal levels in response to oceanic subduction in NE China. Hydrothermal fluids

exsolved from the evolved magma were initially supercritical, weakly saline, acid, oxidizing and of low density, and condensed a high salinity L-V-S or L-V fluid. A model is proposed in which low-density hydrothermal fluids released from an oxidizing A-type magma transported Mo as hydrated MoO3 species and deposited large masses of molybdenite in response to a combination of decreasing temperature, decreasing oxygen fugacity, and condensation (which decreased ligand activity, hydration and temperature). The Hashitu Mo deposit shows important similarities to and differences from typical Climax Mo deposits, which may be helpful in understanding the genesis of porphyry Mo deposits in various tectonic settings. Acknowledgements We thank Mineralium Deposita referees, Thomas Ulrich and Huaying Liang, Associate Editor Shao-Yong Jiang, and Editor-in-Chief Georges Beaudoin for their constructive reviews and comments, which significantly improved this paper. Ed Ripley and Ben Underwood helped with the sulfur isotope analyses, Li Su and Hongyu Zhang with the zircon and sulfide trace element analyses, and Noémie Breillat with the molybdenum isotope analyses. Xingwang Liu and Gongwen Wang helped with the field work and Wenbing Zhu with the preparation of samples for H-O isotope analyses. Discussions with and helpful suggestions from Chusi Li and Sotirios Kokkalas helped us to clarify some of the ideas presented in the manuscript. This research was supported financially by the National Natural Science Foundation of China (Grants 41503042, 41272110), the Fundamental Research Funds for the Central Universities (Grant 2652015045), the Open Research Funds for GPMR (Grant GPMR201513), and the Chinese B111^ project (Grant B07011). An initial draft of the manuscript was prepared during the visit of DZ to Indiana University in 2013–2014, which was funded by the China Scholarship Council.

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