Geology, chronology, fluid inclusions, and H–O–S isotopic compositions of the Hongyuntan magnetite deposit, Eastern Tianshan, NW China

Accepted Manuscript Geology, chronology, fluid inclusions, and H–O–S isotopic compositions of the Hongyuntan magnetite deposit, Eastern Tianshan, NW China Zhi-Yuan Sun, Ling-Li Long, Yu-Wang Wang, Zhao-Hua Luo, Qi-Tao Hu, Meng-Long Wang PII: DOI: Reference:

S1367-9120(18)30407-3 https://doi.org/10.1016/j.jseaes.2018.09.016 JAES 3653

To appear in:

Journal of Asian Earth Sciences

Received Date: Revised Date: Accepted Date:

4 September 2017 20 September 2018 20 September 2018

Please cite this article as: Sun, Z-Y., Long, L-L., Wang, Y-W., Luo, Z-H., Hu, Q-T., Wang, M-L., Geology, chronology, fluid inclusions, and H–O–S isotopic compositions of the Hongyuntan magnetite deposit, Eastern Tianshan, NW China, Journal of Asian Earth Sciences (2018), doi: https://doi.org/10.1016/j.jseaes.2018.09.016

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Geology, chronology, fluid inclusions, and H–O–S isotopic compositions of the Hongyuntan magnetite deposit, Eastern Tianshan, NW China Zhi-Yuan Suna, b, Ling-Li Longb, Yu-Wang Wangb, *, Zhao-Hua Luoa, Qi-Tao Huc, Meng-Long Wangc

a. School of Earth Sciences and Resources, China University of Geosciences, Beijing, 100083 b. Beijing Institute of Geology for Mineral Resources, Beijing, 100012 c. Shanshan Baodi Mining Limited Liability Company, Shanshan, 838204

Abstract

The

Hongyuntan

magnetite

deposit,

which

is

located

in

the

Aqishan–Yamansu belt of Eastern Tianshan, Xinjiang, is hosted within Carboniferous submarine volcanic-sedimentary sequences. The deposit is characterized by five metallogenic

stages:

Stage

I

(garnet–diopside),

Stage

II

(magnetite–tremolite–chlorite–pyrite), Stage III (magnetite–quartz–albite–actinolite), Stage IV (pyrite–quartz–epidote) and Stage V (hematite–limonite). Pyrite associated with magnetite formed during the Stage II has a Re-Os isochron age of 324 ± 31 Ma. Four types of fluid inclusions were observed in the mineral assemblages, namely, two-phase liquid and vapor inclusions, multi-phase daughter mineral-bearing inclusions, multi-phase CO2-bearing inclusions, and monophase vapor inclusions. Microthermometry data reveal that

the ore-forming

fluids evolved

from

high-temperature (average 532.5 °C) and high-salinity (average 26.9 wt.% NaCleqv) fluids to moderate–low-temperature (average 203.6 °C) and low-salinity (5.07 wt.% NaCleqv) fluids. Magnetite precipitation may have been greatly assisted by mixing and boiling in Stage II and Stage III, respectively. The hydrogen and oxygen isotopic values of Stage I garnet (δDV-SMOW = –107.9‰ to –76.7‰, δ18Ofluid = 6.4‰ to 10.7‰) and Stage III quartz (δDV-SMOW = –85.5‰ to –83.2‰, δ18Ofluid = 3.8‰ to 4.7‰) indicate that the ore-forming fluids were dominated by magmatic water, whereas fluid mixing occurred during Stage II (δDV-SMOW = –117.3‰ to –112.6‰, δ18Ofluid = 3.3‰ to 5.7‰) and Stage IV (δDV-SMOW = –115.2‰ to –58.5‰, δ18Ofluid = –1.3‰ to 4.5‰). The variable δ34S values of pyrite (–3.8‰ to 4.7‰) and δ18Omgt values of magnetite (–0.6‰ to 4.1‰) suggest that they were not sourced from a single reservoir. The Hongyuntan magnetite deposit is spatially and temporally associated with volcanism and may be genetically related to it. The fluid inclusions and isotopic compositions of this deposit indicate that its ore-forming fluid was derived from the exsolution of deep-seated magmas and that fluid mixing could have effectively triggered the deposition of ore-forming materials. Multiple periods of volcanic activity could have incorporated additional heat and ore metals into the ore-forming system. Based on our data, we conclude that the origin of the Hongyuntan magnetite deposit is related to the volcanism in the Carboniferous period. Key words Re-Os geochronology; Fluid inclusions; Fluid boiling and mixing; H–O–S isotope; Hongyuntan magnetite deposit; Eastern Tianshan; Xinjiang

1. Introduction The boiling and mixing of ore-bearing fluids effectively change their temperature and solubility, and thus, these processes constitute the most important mechanisms of mineral precipitation in magmatic–hydrothermal systems (Bodnar et al., 1985; Cox et al., 2001; Cooke and McPhail, 2001; Simmons et al., 2005). The concentrations of volatiles and metals in the residual melt increase when magma crystallizes and when fluids separate from the melt upon the occurrence of “second boiling” (Burnham and Ohmoto, 1980). The separated fluids carry considerable amounts of volatiles and metals (Cline and Bodnar, 1991; Roedder, 1992), and these metals can precipitate under appropriate physical and chemical conditions, such as decompression or cooling (Beane and Bodnar, 1995). Accordingly, it is important to investigate the boiling and mixing processes of the fluid to determine the metallogenic process and understand its mechanism (Hemley and Hunt, 1992; Ulrich et al., 2001). According to previous studies on the fluid inclusions and stable isotopes of volcanogenic iron oxide deposits, mineralization is attributable to multi-stage fluid boiling (Xiao et al., 2002; Luo et al., 2015) and/or mixing (Farber et al., 2015; Li et al., 2015). The same processes are largely responsible for the formation of iron deposits (Yu et al., 2015; Hofstra et al., 2016; Johnson et al., 2017). However, few studies have focused on fluid inclusions within submarine volcanic-hosted iron deposits and their metallogenic mechanisms (Yan et al., 2016). The Hongyuntan magnetite deposit, which is located in the western region of the Aqishan–Yamansu belt in the Eastern Tianshan, Northwest China, is representative of deposits associated with volcanic

activity. This deposit has long been researched due to its large tonnage, high grade, and accessibility by open-pit mining. However, the genesis of this deposit and its relationship with volcanism remain controversial due to the lack of precise mineralization age determinations. Moreover, the source(s) of ore-forming fluids and metals and the associated mineralization processes and precipitation mechanisms are still poorly understood. In this paper, Re-Os dating of pyrite is used to provide direct constraints on the timing of magnetite mineralization in the Hongyuntan deposit. We sample minerals from different metallogenic stages, investigate their fluid inclusions, and present evidence of fluid boiling/mixing. Furthermore, to constrain the sources and nature of the ore-forming fluids in the Hongyuntan magnetite deposit, we acquire hydrogen, oxygen, and sulfur isotopic data. These comprehensive data help constrain the origin of the deposit and reveal the relationship between fluid evolution and mineralization, and they also help elucidate the metallogenic processes and mechanisms of this type of deposit. 2. Regional geology The Eastern Tianshan metallogenic belt is located at the intersection of the Siberian, Junggar–Kazakhstan, and Tarim plates, and it is an important part of the Central Asian Orogenic Belt (Fig. 1a) (Qin et al., 2002; Windley et al., 2007; Xiao et al., 2013). From north to south, the Eastern Tianshan mountains successively occupy the Bogda–Harlik tectonic belt, the Turpan–Hami basin, the Jueluotage tectonic belt, and the Middle–Tianshan block (Xiao et al., 2013) (Fig. 1b). As indicated by the network of faults and the distribution of ore deposit types (Fig. 1b), the Jueluotage

tectonic belt comprises the copper belt along the southern margin of the Turpan–Hami basin (north belt), the Kangguer gold belt (middle belt), and the Aqishan–Yamansu iron–copper–silver belt (south belt) (Wang et al., 2006). The Aqishan–Yamansu belt is well known for the iron–copper deposits hosted in its Carboniferous submarine volcanics (Mao et al., 2005; Zhang et al., 2016). Typical and well-known Fe deposits therein include the Yamansu and Hongyuntan deposits, which have similar mineralization styles and whose orebodies are structurally developed into banded or lenticular formations. These deposits are in conformable contact with Carboniferous volcanic rocks, and they have no evident spatial association with specific igneous intrusions (Hou et al., 2014). Additionally, these deposits are closely related to submarine magmatic activity (Jiang and Wang, 2005; Li et al., 2015). 3. Deposit geology 3.1. Stratigraphy, structures, and igneous rocks The Hongyuntan magnetite deposit is located approximately 120 km to the southeast of Shanshan County in Xinjiang Province. This deposit contains estimated ore reserves of 670.4 × 104 t with a rich average iron grade of 44 % total Fe (XJBGMR, 2009). This deposit contains two ore sections: the east section, which contains orebodies III and V, and the west section, which contains orebody II. The exposed stratum in the Hongyuntan district is the Lower Carboniferous Yamansu Formation (Fig. 2), which primarily consists of mafic–intermediate volcaniclastic rocks, acidic volcanic lavas, and clastic sedimentary rocks. The stratum trends nearly NE with dip angles ranging from 30° to 60°. From its base upward, the Yamansu

Formation is characterized by deposits produced during three periods of volcanic activity: C1 y1, C1y2, and C1 y3. C1 y1, which is located in the northern part of the district, consists of stratiform tuffaceous breccia, tuffaceous glutenite, and tuffaceous siltstone. The lower C1y2 lithologies consist of andesitic tuff and tuffaceous siltstone interbedded with breccia, ignimbrites, and intermediate volcanic lavas. These rocks have been hydrothermally altered by silicification, garnetization, and chloritization. Orebodies are mainly hosted in the C1y2 strata. The uppermost portion of the Yamansu Formation, C1 y3, is mainly composed of quartz keratophyre. Multiple periods of volcanic activity have occurred in the Hongyuntan district. A previous study related these periods of volcanic activity to the multilayered orebodies, multilayered alteration-mineralization zones, and ranges of magnetite decrepitation temperatures in this district (Wang et al., 1980). Furthermore, based on analyses of the ZK502 drill hole (Fig. 3), at least two eruption cycles occurred during C1 y2. The first, which ranges from 460 to 251 m, comprises tuffaceous breccia → andesitic crystal tuff → andesitic tuff → tuffaceous sandstone. The second, which ranges from 251 to 52 m, comprises andesitic crystal tuff → andesitic tuff → sedimentary rocks. The orebodies in C1 y2 and C1 y3 occur along the axis of a NE-trending normal syncline in the Hongyuntan district (Fig. 2). Widespread fault systems are observed in this district, and the main faults (F1, F2, and F6) belong to NE-trending fault systems. The secondary nearly NS-trending faults (F7, F8, and F9) postdate the mineralization and crosscut the volcanic rocks. The predominant intermediate and acidic intrusions in this district are diorite,

quartz diorite, granodiorite, granite, and moyite. Zircon U-Pb data obtained through laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) reveal that the quartz diorite and moyite crystallized at (351.5 ± 1.2) Ma and (297.36 ± 0.51) Ma, respectively (Zheng, 2015). Moreover, the orebodies are crosscut by moyite and diabase dikes. 3.2. Orebodies and ores The orebodies are hosted in the andesitic tuff, crystal tuff, and tuffaceous siltstone of the Yamansu Formation. Structurally, these orebodies display banded or lenticular forms, and they are in conformable contact with volcanic rocks. The orebodies commonly trend nearly NE (180° to 240°), and they dip northwestward at angles of 20 ° ~ 70 °. They are 100 m ~ 300 m long and 5 m ~ 15 m wide, and their maximum depth in the profile reaches up to 240 m. Sixteen orebodies can be recognized in the Hongyuntan district. The main orebody in the west section is the No. II3 orebody, which trends nearly northeast. This orebody can be divided into four ore blocks: II3-1, II3-2, II3-3, and II3-4. The II3-2 ore block is relatively shallow and has consequently been mined. The II3-1, II3-3, and II3-4 ore blocks are distributed in parallel (Fig. 4). It is worth noting that no spatial association between the orebodies and intrusions has been observed during geological field surveys within the Hongyuntan district. Massive, disseminated, banded, and veined ores are found in the Hongyuntan ore district. The massive ores display massive structures with highly variable amounts of magnetite and subordinate pyrite (Fig. 5a). The magnetite crystals in the massive ores

are euhedral to subhedral and exhibit platy textures (Fig. 5e), and they range in length from 200 μm to 500 μm. The disseminated ores are characterized by disseminated magnetite in chlorite (Figs. 5b, f). The magnetite grains range from subhedral to xenomorphic, and they are 50 μm to 100 μm in diameter. The banded ores typically exhibit parallel or subparallel magnetite and epidote bandings that are 1 cm to 2 cm in width (Fig. 5c). However, microscopic observations reveal indistinct boundaries between the different minerals in the banded ores (Fig. 5g). The veined ores are characterized by magnetite veins or by quartz + magnetite + actinolite ± pyrite veins with variable widths (0.5 cm ~ 5 cm), and these veins crosscut massive magnetite + chlorite (Fig. 5d) or garnet (Fig. 7b). In addition, the textural relationship between the intergrown pyrite and magnetite is unclear (Fig. 5h), implying that they precipitated synchronously. 3.3. Alteration, mineral assemblages, and generations The wall rocks in the Hongyuntan district have been hydrothermally altered to different extents. The morphologies of the alteration zones from the bottom up (as determined in the ZK502 drill hole, see Fig. 3) are albitization, garnetization, chloritization, actinolitization, silicification, epidotization, and carbonatization. Based on the paragenesis and crosscutting relationships of the minerals and veins, we can identify three principal periods, or five alteration stages (Fig. 6). The early period, which is characterized by abundant magnetite, garnet, chlorite and tremolite, can be divided into two sub-stages: Stage I and Stage II. Stage I is dominated by garnet and diopside with lesser amounts of albite. The garnet, which is

brown to reddish-brown in color and ranges from 1 mm to 5 mm in size, is commonly present as massive cumulate structures that are cut by quartz + actinolite, magnetite, or quartz + epidote veins (Figs. 7a, 7b, 7d). Under a binocular microscope, the garnet is replaced along concentric zoning rims by magnetite + quartz (Fig. 7i) or enclosed by magnetite (Fig. 7j). The diopside grains are replaced by magnetite (Fig. 7k), and these are both cut by quartz + epidote vein (Fig. 7c). As the major magnetite mineralization stage, Stage II is characterized by a considerable amount of magnetite, chlorite, and tremolite, wherein the chlorite is accompanied by massive and disseminated magnetite (Figs. 5b, 5f). The middle period is typically characterized by abundant quartz, pyrite, and epidote as well as subordinate amounts of magnetite and actinolite. The quantity of ore resources in this period is lower than that in the early period. Minor amounts of albite, K-feldspar, and chalcopyrite are also present. This period can also be divided into two sub-stages: Stage III and Stage IV. The metasomatism of Stage III is mainly recorded as quartz, magnetite, pyrite, and actinolite. The magnetite is intergrown with quartz and pyrite and is characterized by veined structures (Figs. 5d, 7e). Later albite and actinolite crosscut the massive magnetite + chlorite ores (Fig. 7f). The minerals of Stage IV are dominated by quartz, epidote, and pyrite with lesser amounts of chalcopyrite. Large quantities of pyrite occur as euhedral to subhedral grains (Figs. 7g, 7h). Chalcopyrite associated with quartz and pyrite occurred as veins that replace magnetite (Fig. 7l). Finally, the late period is characterized by minor supergene minerals, including limonite, malachite, and hematite.

4. Sampling and analytical methods Five pyrite samples were collected from the massive ores (Fig. 5a) in Stage II within the mining pit of the Hongyuntan district for Re-Os isotope dating. Single pyrite samples weighing approximately 200 mg ~ 600 mg were obtained using traditional isolation methods. Re-Os isotope analyses were carried out using a negative thermal surface ionization mass spectrometer (Triton-plus, Thermo Fisher Scientific) at the National Research Center of Geoanalysis, Chinese Academy of Geological Sciences. The analytical procedures are described in detail by Du et al. (2009). Samples used for fluid inclusion and stable isotope analyses were collected from drill holes, mining pit and shallow levels of ore-bearing veins. Doubly polished wafers (200 μm thick) were examined and petrographically photographed. The fluid inclusions in garnet, tremolite, quartz, epidote, and calcite were observed using standard techniques (Roedder, 1984), and microthermometric measurements were performed on a Linkam THMS600 heating–freezing stage (from –196 °C to 600 °C) at the Beijing Institute of Geology for Mineral Resources. To calibrate the stage, we measured the melting points of distilled water (0.0 °C), pure CO2 inclusions (–56.6 °C), and potassium bichromate (398 °C). The accuracies of the measured temperatures were approximately ±0.2 °C during cooling and ±2 °C between 100 °C and 600 °C during heating. We performed measurements near the clathrate-melting temperatures and the melting temperatures of the carbonic phase at a heating rate from 0.1 °C/min to 0.2 °C/min; for the other measurements, we used a heating rate

from 0.2 °C/min to 0.5 °C/min. The δD values of the fluid inclusions and the δ18O values of their host minerals were measured at the Beijing Research Institute of Uranium Geology, China. We performed thermal decrepitation at 400 °C to release the water contained in the fluid inclusions of the host minerals and then collected, froze, and purified the released water. Then, using reductive zinc, we replaced and released the hydrogen in the water to perform mass spectrography. A detailed explanation of these analytical methods can be found in Coleman et al. (1982). The analytical precision of the δD values exceeded ±2.0‰. The oxygen isotopes were analyzed by the conventional method (Clayton and Mayeda, 1963), in which the minerals reacted with BrF5 at high temperatures (500 °C ~ 550 °C) for at least five hours. Then, the liberated oxygen was converted to carbon dioxide using a hot platinized carbon rod at high temperature (600 °C) for mass spectrometry analysis. The analytical precision of the δ18O values was within ±0.2‰. We reported the hydrogen and oxygen isotopic data per milliliter relative to the Vienna Standard Mean Ocean Water (SMOW) standard. We separated the pyrite samples from the massive, disseminated, banded, and veined ores for sulfur isotope analysis. The sulfur isotopic data of pyrite in volcanic rocks, breccia ores, and mineralized rocks have been reported in previous studies. Sulfur isotope measurements were performed at the Beijing Research Institute of Uranium Geology, China. Fourteen pyrite samples were converted to SO2 under high-temperature vacuum conditions. The δ34S values were measured on the MAT251 gas isotope mass spectrometer. We reported the sulfur isotopic data as δ34S in per

milliliter relative to the Canyon Diablo Troilite (CDT) standard, and the precision was within ±0.2‰. 5. Results 5.1. Re-Os geochronology The Re-Os data of five pyrite samples from the Hongyuntan magnetite deposit are listed in Table 1. The total Re concentrations of the pyrite samples vary greatly (7.31 to 72.87 ppb), but the samples have low common Os concentrations (0.01515 to 0.07155 ppb) and high

187

Re/188Os ratios (844.8 to 7489), which are similar to the

low-level highly radiogenic (LLHR) sulfides defined by Stein et al. (2000). Therefore, isochron plots of the

187

Re versus

187

Os are used, and the isochronal age was

calculated using the ISOPLOT software (Ludwig, 2003). The five pyrite samples yielded an isochron with an age of 324 ± 31 Ma, an initial 187Os ratio of 0.001 ± 0.017, and a mean square weighted deviation (MSWD) of 9.6 (Fig. 8). 5.2. Fluid inclusion petrography and inclusion types Four types of primary fluid inclusions were observed in the host minerals based on the phases present at room temperature (Roedder, 1984), including Type I: two-phase

liquid–vapor

inclusions

(L–V),

Type

II:

multi-phase

daughter

mineral-bearing inclusions (S–L–V), Type III: multi-phase CO2-bearing inclusions, and Type IV: monophase vapor inclusions (V). Type I inclusions are the most common type of inclusions, and they can be observed in every stage. These inclusions consist of a liquid (L) and a vapor bubble

(V). Based on the volumetric dominance of the vapor phase at room temperature and the phase transitions during heating and cooling, Type I inclusions are divided into two subtypes: Ia and Ib. Type Ia inclusions, which usually consist of a dominant liquid and a subordinate vapor (comprising 10 % ~ 40 % of the total volume), are elliptical, rod-shaped, round, and irregular, and they include some negative crystal forms (Figs. 9a, 9d, 9n, 9o). They range in size from 6 μm to 23 μm. Type Ia inclusions are distributed as clusters or as isolated or minor linear arrays, and they are mainly homogenized to the liquid phase. Type Ib inclusions are usually dominated by the vapor phase (Figs. 9b, 9c, 9g, 9i), wherein the vapor comprises 60 % ~ 90 % of the total volume. These vapor-rich inclusions appear as irregular isolates in garnet and quartz. The shapes of the inclusions are generally elliptical, rod-like, or round, ranging in size from 6 μm to 22 μm, and they are mainly homogenized to the vapor phase. Type II inclusions occur in Stage I garnet and Stages III and IV quartz. These inclusions typically contain a liquid, a vapor, and one or more solids at room temperature. Type II inclusions are typically irregularly shaped and range in diameter from 7 μm ~ 22 μm. Based on their cubic shapes, most of the daughter minerals in the Type II inclusions appear to be halite (Figs. 9a, 9d, 9j, 9k). Furthermore, some isotropic sylvite crystals (Fig. 9l), transparent crystals (Figs. 9e, 9i), and opaque crystals (Figs. 9b, 9d, 9e, 9l) are also observed in Type II inclusions. Minor hematite crystals (Fig. 9m) are observed only in Stage III quartz. Based on the temperature correlation between the final halite dissolution and the final disappearance of vapor bubble during the heating process, Type II inclusions can be divided into three

subtypes: IIa, IIb, and IIc. Type IIa inclusions are characterized by final homogenization by bubble disappearance following halite dissolution. Type IIb inclusions are characterized by halite dissolution after bubble disappearance during the heating process. Type IIc inclusions are homogenized to the liquid phase, and their temperatures match those of the daughter mineral dissolution and bubble disappearance. Remarkably, both the Stage I garnet and the Stage III quartz contain coexisting Type Ia, Type Ib, and Type II inclusions within a single population (Figs. 9a, 9k). Type III inclusions are limited in number and occur only in Stage III quartz crystals. They mainly contain vapor CO2 and liquid CO2 in addition to an aqueous liquid and, occasionally, one or more solids (transparent and/or opaque minerals) at room temperature (Fig. 9l). CO2 vapor appears to account for 10 % ~ 20 % of the total volume. Type III inclusions typically coexist with Type II inclusions as fluid inclusion assemblages (Fig. 9l). Type IV inclusions are characterized by a single vapor phase (Figs. 9a, 9h) and are predominant in garnet and quartz with elliptical, round, and irregular forms. 5.3. Fluid inclusion microthermometry Microthermometric measurements were performed on the four types of inclusions in garnet, tremolite, quartz, epidote and calcite. The results are presented in Table 2 and Fig. 10. The salinities (wt.% NaCl equivalent) of the Type I and Type IV fluid inclusions were calculated based on their final ice-melting temperatures (Tm, ice) (Hall et al., 1988). Meanwhile, the salinities of the Type II and Type III fluid inclusions were calculated based on the halite dissolution temperatures (Tm, s) (Hall

et al., 1988) and the clathrate melting temperatures (Collins, 1979), respectively. Stage I

The homogenization temperatures of the Type Ia inclusions in garnet range from 466 °C to 482 °C (average 476.8 °C), and they are finally homogenized to the liquid phase. Their final ice-melting temperatures range from –11.2 °C to –6.8 °C (average –8.7 °C), corresponding to salinities ranging from 10.2 wt.% to 15.2 wt.% NaCl equivalent (average 12.8 wt.% NaCl equivalent). The homogenization temperatures of the Type Ib inclusions in garnet range from 446 °C to 592 °C (average 549.1 °C) but are higher (> 600 °C) in some cases. These inclusions are finally homogenized to the vapor phase. The final ice-melting temperatures range from –14.2 °C to –6.0 °C (average –9.3 °C), corresponding to salinities ranging from 9.2 wt.% to 16.1 wt.% NaCl equivalent (average 12.5 wt.% NaCl equivalent). The homogenization temperatures of the Type IIa inclusions in the garnet range from 546 °C to 577 °C (average 567.3 °C) and consistently show final homogenization by bubble disappearance after halite dissolution. The halite dissolution temperatures range from 258 °C to 279 °C (average 271.1 °C), corresponding to salinities ranging from 35.2 wt.% to 36.6 wt.% NaCl equivalent (average 36.1 wt.% NaCl equivalent). The homogenization temperatures of the Type IIb inclusions in garnet range from 486 °C to 551 °C (average 515.5 °C) and consistently exhibit final homogenization by halite dissolution after bubble disappearance. The salinities calculated from these fluid inclusions range from 59.5 wt. % to 63.2 wt. % NaCl equivalent (average 61.9 wt.% NaCl equivalent). As explained above, petrographic evidence reveals the coexistence of Type I, II, and IV fluid inclusions in the Stage I garnet, and these coexisting

inclusions exhibit similar ranges of homogenization temperatures but different salinities. Specifically, the opaque minerals and minor transparent minerals did not disappear by heating. Thus, we assume that the bubble disappearance temperatures are the homogenization temperatures. Stage II

The homogenization temperatures of the Type Ia inclusions in tremolite range from 249 °C to 308 °C (average 275.4 °C) with peaks between 260 °C and 300 °C. Their final ice-melting temperatures range from –17.2 °C to –4.7 °C (average –6.0 °C), corresponding to salinities ranging from 7.4 wt.% to 10.7 wt.% NaCl equivalent (average 9.2 wt.% NaCl equivalent). Stage III

The homogenization temperatures of the Type Ia inclusions in quartz range from 412 °C to 462 °C (average 438.3 °C). The final ice-melting temperatures range from –10.2 °C to –7.4 °C (average –9.1 °C), corresponding to salinities ranging from 12.4 wt.% to 14.1 wt.% NaCl equivalent (average 12.7 wt.% NaCl equivalent). The Type Ib inclusions in quartz exhibit homogenization temperatures of 325 °C ~ 540 °C (average 449.3 °C), with peaks between 420 °C and 440 °C. The final ice-melting temperatures of the Type Ib inclusions range from –9.6 °C to –3.1 °C (average –5.7 °C), corresponding to salinities ranging from 5.1 wt.% to 13.5 wt.% NaCl equivalent. The homogenization temperatures and corresponding salinities of the Type IIb inclusions in quartz range from 381 °C to 496 °C (average 441.6 °C) and from 36.5 wt. % to 57.1 wt. % NaCl equivalent (average 49.9 wt.% NaCl equivalent), respectively. The coexisting Type I and Type II inclusions in the same region (Figs. 9g,

9i) have similar homogenization temperatures but different salinities (Fig. 10). Specifically, some of the halite within the halite-bearing inclusions (characterized by halite volume occupancies of 90 % ~ 95 %) in this stage does not melt below 520 °C. The homogenization temperature of the Type III inclusion in quartz is 438 °C, and the clathrate melting temperature is 4.2 °C. Consequently, the salinity of Type III inclusion is 10.2 wt.% NaCl equivalent. Stage IV

Type Ia inclusions are the most abundant inclusions in the quartz and epidote from Stage IV. These inclusions are homogenized to the liquid phase between 215 °C and 322 °C (average 254.1 °C) with temperature peaks between 240 °C and 260 °C. The final ice-melting temperatures of the Type Ia inclusions range from –11.6 °C to –2.5 °C (average –7.1 °C), corresponding to salinities ranging from 4.2 wt.% ~ 15.6 wt.% NaCl equivalent. The homogenization temperatures of the Type IIa inclusions in quartz range from 243 °C to 318 °C (average 270.7 °C), and their halite dissolution temperatures and corresponding salinities range from 176 °C to 276 °C (average 255.8 °C) and from 30.7 wt.% to 36.4 wt.% NaCl equivalent (average 35.1 wt.% NaCl equivalent), respectively. The homogenization temperatures of the Type IIc inclusions in quartz range from 250 °C to 269 °C (average 258.8 °C), and their halite dissolution temperatures and corresponding salinities range from 251 °C to 270 °C and from 34.7 wt.% to 35.9 wt.% NaCl equivalent (average 35.3 wt.% NaCl equivalent), respectively. The Stage IV calcite veins contain only Type Ia fluid inclusions, which are homogenized at temperatures ranging from 158 °C to 286 °C (average 203.6 °C), with peaks appearing in the range of 200 °C ~ 220 °C. The final

ice-melting temperatures of the Type Ia inclusions range from –7.7 °C to –0.9 °C, corresponding to salinities ranging from 1.5 wt.% to 11.3 wt.% NaCl equivalent. 5.4. Stable isotopes 5.4.1. H–O isotopes The hydrogen and oxygen isotopic data of the host minerals and fluid inclusions are summarized in Table 3. The δDV-SMOW values range from –107.9‰ to –76.7‰ for garnet, from –117.3‰ to – 112.6‰ for tremolite, from –115.2‰ to –83.2‰ for quartz, from –97.7‰ to –93.2‰ for K-feldspar, and from –99.4‰ to –58.5‰ for epidote. The δ18OV-SMOW values range from 3.5‰ to 7.8‰ for garnet, from 3.6‰ to 6.0‰ for tremolite, from 7.7‰ to 11.9‰ for quartz, from 11.5‰ to 11.7‰ for K-feldspar, and from 3.4‰ to 6.4‰ for epidote. The δ18Omgt values of magnetite sampled from massive and banded ores range from –0.6‰ to 4.1‰. The δ18Ofluid values were calculated using oxygen-isotope fractionation equilibrium formulas (Zheng et al., 1993a, 1993b; Zheng, 2000) (Table 4), where T is the homogenization temperature in Kelvin (K). The calculated δ18Ofluid values range from 6.4‰ to 10.7‰ for garnet, from 3.3‰ to 5.7‰ for tremolite, from –1.3‰ to 4.7‰ for quartz, and from 1.5‰ to 4.5‰ for epidote. The calculated δ18Ofluid values of magnetite range from 7.6‰ to 12.6‰. 5.4.2. S isotopes The Hongyuntan deposit yielded a single sulfur mineral (pyrite) for sulfur isotopic analysis. The forty-four δ34S values of the pyrite samples vary widely from

–3.8‰ to 4.7‰ (Table 5). The sulfur isotopic compositions of the massive, veined, and banded ores are symmetrical and cluster near zero. The δ34S values of the pyrites in the volcanic rocks are negative, ranging from –3.5‰ to –2.0‰, whereas those of the pyrite samples obtained from breccia ores, mineralized rocks, and some veined ores are highly positive, ranging from 1.4‰ to 4.7‰. 6. Discussion 6.1 Mineralization age and its significance To date, the genesis of the Hongyuntan magnetite deposit remains controversial. A skarn-type deposit origin is supported by the skarn alteration observed in the Hongyuntan district (Zhang et al., 2012; Zhang et al., 2013a). However, the Hongyuntan deposit has also been proposed to be a typical volcanic-hydrothermal deposit (Zheng, 2015). The mineralization age of a deposit can help interpret its genesis and thus could help resolve this controversy (Chai et al., 2014). Our dating of pyrite samples from the Stage II ores yields an isochron age of 324 ± 31 Ma, which suggests that the deposit was formed during the Early Carboniferous. This new Re-Os age is consistent with the U-Pb age (324.1 ± 3.1 Ma) of the volcanic rocks from the Yamansu Formation (Zheng, 2015) rather than the previously reported ages of intrusions (Wu et al., 2006; Zheng, 2015), suggesting a genetic relationship between mineralization and volcanism. Moreover, no geological field evidence has demonstrated that the formation of magnetite ores is related to intrusive activity in the Hongyuntan district. We conclude that Hongyuntan magnetite mineralization may be related to the volcanic activity that occurred in the Carboniferous period. This

hypothesis is similarly supported by the accordant ages of volcanic rocks and mineralization in other magnetite deposits throughout the Aqishan–Yamansu belt, including the Yamansu magnetite deposit (Hou et al., 2014; Huang et al., 2018), which also suggest that the magnetite mineralization is related to volcanic activity. 6.2 Origin of ore-forming fluids As shown in the δD versus δ18Ofluid plot (Fig. 11), the calculated δ18Ofluid values in Stage I range from 6.4‰ to 10.7‰, and these values are consistent with those of fluids derived from magmatic sources (6.0 ‰ ~ 10.0 ‰; Taylor, 1974). The δD values vary widely from –107.9 ‰ to –76.7 ‰, showing a pronounced tendency of fractionation. This remarkable decrease in δD may have resulted from the boiling of ore-forming fluid (Taylor, 1974). Experimental studies have suggested that D preferentially fractionates into the vapor phase relative to H during fluid boiling, reducing the δD values of the residual fluid by approximately 28 ‰ while scarcely changing the δ18O values (Shmulovich et al., 1999; Driesner and Seward, 2000). The petrographic evidence for the coexistence of Type I, II, and IV fluid inclusions (Fig. 9a) reveals that fluid boiling occurred in Stage I, further supporting our interpretation of fractionation. In Stage II, the δ18Ofluid values of tremolites deviate from those of magmatic water. Their lower values may indicate water–rock interactions and/or the mixing of magmatic fluids with shallow non-magmatic fluids (Shmulovich et al., 1999). The Stage III quartz samples fall into the lower-left area associated with magmatic water on the plot of δD versus δ18Ofluid. According to previous studies, the

δD and δ18O values in a closed system can continue to increase during the interactions between fluid and magmatic fluid (Taylor, 1997). Accordingly, multiple volcanic eruptions have occurred in the Hongyuntan magnetite deposit, and its energy and materials may have been replenished during a second eruption (Corriveau et al., 2016), resulting in these increased δD and δ18Ofluid values (Rieger et al., 2012). This concept of replenishment is supported by the high homogenization temperatures (440 °C ~ 460 °C) of the fluid inclusions in the Stage III quartz. The δ18Ofluid values of the quartz samples decreased in Stage IV (average 2.1 ‰). The δD values of these samples, especially those of epidote, vary widely from –115.2 ‰ to –58.5 ‰ (Fig. 11). Three possible mechanisms can account for this phenomenon (Taylor, 1977): (1) the fluids may have undergone boiling; (2) magmatic water may have mixed with fluids such as seawater containing heavier hydrogen isotopes; or (3) the fluids may represent metamorphic or sedimentary formation water. We found no evidence of fluid boiling in the epidote. Furthermore, no regional metamorphism or metamorphic rocks have been identified in the Hongyuntan district. However, Type IIc fluid inclusions were found in the Stage IV quartz, which may reflect fluid boiling (Lu et al., 2004). Therefore, we consider that the wide-ranging δD values were caused by the incorporation of seawater into these fluids in conjunction with fluid boiling. The oxygen isotope (δ18Omgt) values of the Stage II magnetite in the Hongyuntan deposit cover the typical ranges of other volcanic iron deposits, including the Kiruna-type and volcanic-type iron deposits in the Western Tianshan (Fig. 12). The calculated δ18Ofluid values of magnetite are higher than those of magmatic origin

(Taylor, 1967), suggesting that the ore-forming fluids did not equilibrate with a primary magma or with a high-temperature magmatic fluid. Based on the above discussion, we infer that the magnetite ores in Stage II predominantly precipitated from mixed fluids through reactions between magmatic fluids and seawater. 6.3 Origin of ore-forming materials The wide range and multiple peaks of sulfur isotopic values in the Hongyuntan magnetite deposit (Fig. 13) suggest that the metallogenic materials therein were not sourced from a single reservoir. The pyrite crystals from the main ore types (massive, veined, and banded ores) have δ34S values with peaks ranging from –1.0 ‰ to 0 ‰, indicating that most of the sulfides were deeply sourced and suggesting the involvement of magmatic sulfur (Hoefs, 1997). The δ34S values of the pyrite crystals from volcanic rocks are negative but close to those of mantle-sourced sulfur, implying that the primary sulfur experienced decompression and degasification through volcanic activity (Ren, 1985). Previous studies have shown that the δ34S values in submarine hydrothermal deposits are elevated (Plimer and Finlow-Bates, 1978). The elevated δ34S values (approximately 15 ‰ lower than those of the oceans) of the pyrite crystals sampled from breccia ores, mineralized rocks, and some veined ores are consistent with the δ34S values of submarine hydrothermal deposits (Sangster, 1968). This change in sulfur concentration was probably caused by the restricted degree of seawater circulation (Plimer and Finlow-Bates, 1978; Song et al., 2003).

6.4 Nature and evolution of the fluids The microthermometry results reveal that the various types of inclusions in different stages exhibit wide ranges of temperatures and salinities. The fluid inclusions in the Stage I garnet have relatively high homogenization temperatures (with peaks ranging from 540 °C to 560 °C) and high halite dissolution temperatures, indicating the direct involvement of a magmatic–hydrothermal fluid source (Meinert et al., 2003). Moreover, the garnet contains coexisting Type I, II, and IV fluid inclusions, which have similar homogenization temperatures but different salinities (Figs. 9, 10), suggesting that fluid boiling occurred in Stage I (Baker and Lang, 2003). We observed black opaque minerals (presumably magnetite) in the garnet, which are generally interpreted to represent the abundant metallogenic materials carried by fluids. The Type Ia inclusions in the Stage II tremolite are characterized by medium temperatures (average 275.4 °C) and medium salinities (average 9.2 wt. % NaCl equiv). Fluid mixing, potentially combined with cooling, is of great importance in the precipitation of magnetite. This explains the formation of massive and disseminated ores during Stage II. As mentioned above, multiple volcanic eruptions occurred in the Hongyuntan district. The episodic ascent and emplacement of magma could have injected additional heat and ore metals into the ore-forming system (Sáez et al., 1999; Chen et al., 2012). The peak homogenization temperatures of the fluid inclusions in Stage III quartz range from 420 °C to 440 °C. The coexistence of halite-bearing and vapor-rich inclusions suggests that fluid boiling occurred during Stage III and may have formed

the magnetite and magnetite + quartz ± actinolite veins. Although the cooling of vapor-rich fluids could also have led to the same result (Chu et al., 2010), Type II inclusions are much more abundant than Type Ib inclusions in the Hongyuntan deposit, and vapor-rich Type II inclusions are also trapped in Stage III quartz (Fig. 9k). As these highly saline, vapor-rich inclusions could not have been formed by condensation, they must have been produced by trapped boiling fluids. In addition, we observed black opaque minerals, Type III inclusions, and hematite in Stage III (Figs. 9l, 9m), further indicating that mineral precipitation was triggered by the boiling of fluids containing ore metals (Soloviev, 2011). The fluid inclusions in the quartz and calcite from Stage IV are characterized by medium-low temperatures and medium-low salinities (Fig. 10). In this stage, the presence of Type IIc fluid inclusions in quartz may reflect fluid boiling (Lu et al., 2004), which would have driven the precipitation of sulfides. The homogenization temperatures clearly decrease from Stage III to Stage IV, implying a cooling process. The first melting temperatures of the Type Ia fluid inclusions in Stage IV range from –52.3 °C to –30.5 °C, suggesting that the fluids contained variable amounts of NaCl, CaCl2, and MgCl2 (Shepherd et al., 1985; Lu et al., 2004) or were involved with circulating seawater. Hence, fluid boiling and mixing both occurred in Stage IV. Our conclusion is also supported by the previously discussed isotopic data obtained from Stage IV. 6.5 Trapping pressure and mineralization depth The trapping pressure can be estimated when the actual trapping temperature is

known or if fluid boiling or immiscibility occurred at the time of entrapment in the system (Roedder and Bodnar, 1980). In the Hongyuntan deposit, we found evidence of immiscibility in the Stage I garnet and Stage III quartz. Fluid boiling occurred during Stages I, III, and IV. Therefore, we interpret the homogenization temperatures as approximate trapping temperatures, and we determine the pressure based on the microthermometry data from the fluid inclusions (Roedder, 1984). However, we can only estimate the minimum pressures of the fluids in Stage II because of the availability of indirect evidence of fluid boiling (Bouzari and Clark, 2006). The homogenization temperatures of the fluid inclusions in the Stage I garnet are governed by the final homogenization through both bubble disappearance and halite dissolution. Therefore, we cannot estimate the trapping pressures using the diagram developed by Cline and Bodnar (1994), but we can apply the diagrams of Ususova (1975), Haas (1976), and Bodnar et al. (1985). As shown in Fig. 14, the calculated trapping pressures in Stages I, III, and IV range from 32.0 to 91.1 MPa, from 11.5 to 64.2 MPa, and from 1.8 to 11.4 MPa, respectively. In addition, the estimated minimum trapping pressures of Stage II range from 3.5 to 9.2 MPa. Considering that the ore-hosting rocks of the Hongyuntan deposit are mainly composed of andesite and andesitic tuff, and assuming a closed fluid system with a rock density of 2.6 g/cm3, we calculated the fluid depths of Stages I, II, III, and IV to be 1.19 ~ 3.37 km, 0.13 ~ 0.34 km, 0.43 ~ 2.38 km, and 0.07 ~ 0.41 km, respectively.

6.6 Ore-forming process and genetic model The Hongyuntan magnetite deposit experienced multistage alteration and mineralization processes linked to volcanic-hydrothermal activity. The fluid inclusion microthermometry data and hydrogen, oxygen, and sulfur isotopic composition data suggest that the ore-forming fluids were derived from two fluid end-members: magmatic–hydrothermal fluid and externally circulating seawater. These various types of fluid inclusions indicate a sequence of events, with boiling, mixing, or cooling occurring throughout the mineralization process (Fig. 15). Similar ore-forming fluid characteristics have been observed in previous studies of volcanic-type deposits (Degens and Ross, 1969; Rona, 1984; Damm, 1990). Based on the microthermometry data and trapping pressures as well as the fluid sources for different stages, we propose a genetic model for the formation of the Hongyuntan magnetite deposit (Fig. 16). In previous studies of submarine volcanic-hydrothermal fluids, high-temperature and high-salinity fluids were believed to favor the transportation of ore-forming metals (Huston et al., 2010). We suggest that the formation of the Hongyuntan magnetite deposit may have been associated with an initially supercritical fluid that was exsolved from a magma chamber. The exsolved fluid, which carried an abundance of metals, was characterized by high temperature, high salinity, and high volatility. When this fluid ascended to a depth of approximately 3 km, regional brittle faults caused abrupt changes in the pressure and physicochemical conditions of the fluid. These changes induced fluid boiling in Stage I. After boiling, the temperature and solubility of the ascending hot fluid were reduced by reacting with wall rock

and/or cooler, non-magmatic fluids. Hence, mineral assemblages comprising chlorite, tremolite, and considerable amounts of magnetite precipitated in Stage II. The heat and materials of the ore-forming system were recharged during periodic volcanic activity (Ohmoto, 1996; Sáez et al., 1999). The fluid then boiled due to decompression and separated into a low-salinity vapor phase and a hypersaline liquid phase (Shepherd et al., 1985) during Stage III. Meanwhile, magnetite and magnetite + quartz ± actinolite veins began to form over the mineral assemblages that developed during the early period. The ore-forming fluid system ascended to a shallow depth and progressively mixed with and was cooled by the circulating seawater in Stage IV, creating favorable conditions for the accumulation of sulfide minerals. Hence, as the fluid temperature dropped and the fluid boiled, the sulfides were ultimately deposited. 7. Conclusions (1) The Hongyuntan magnetite deposit is hosted in submarine volcanic rocks, which contain orebodies exhibiting banded or lenticular forms. Re-Os pyrite geochronology suggests that the mineralization has an affinity to volcanism. The ore-forming process can be divided into three periods (comprising five stages), and magnetite mainly formed during Stage II. (2) The temperatures and salinities of the ore-forming fluids fluctuated from the early stage to the late stage. Magnetite most likely formed by fluid mixing (Stage II) and boiling (Stage III). The boiling and mixing of fluids during Stage IV played an important role in triggering the deposition of sulfides. (3) The ore-forming fluids were of magmatic origin during the early stage and

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Figure Captions Fig. 1 Sketched tectonic map of the Central Asia Orogenic Belt (a. after Jahn et al., 2000) and geological map with thedistribution of deposits associated with the Eastern Tianshan (b. modified from Wang et al., 2006).

Fig. 2 Geological map of the Hongyuntan magnetite deposit (XJBGMR, 2009).

Fig. 3 Borehole log and lithostratigraphic description of the ZK502 drill hole in the Hongyuntan magnetite deposit. Fig. 4 NE-trending orebody II3, which is banded and lies parallel (looking northeast) to the ore blocks in the Yamansu Formation.

Fig. 5 Photographs of hand specimens (a–d) and micrographs of various magnetite samples (e–h) representing different types of ores. a – massive ore dominated by magnetite. b – disseminated ore containing both magnetite and chlorite. c – banded ore, magnetite and epidote developed by interphase.

d – chlorite + magnetite replaced by quartz + magnetite + pyrite + actinolite vein. e – platy magnetite replaced by quartz + epidote + pyrite (reflected light). f – disseminated magnetite distributed in chlorite (plane-polarized light). g – magnetite replaced by epidote + pyrite (reflected light). h – coexisting magnetite and pyrite (reflected light). Abbreviations: Mag – magnetite; Py – pyrite; Chl – chlorite; Ep – epidote; Qtz – quartz; Act – actinolite.

Fig. 6 Alteration and mineralization paragenesis of the Hongyuntan magnetite deposit.

Fig. 7 Photographs of representative samples and their relationships in the Hongyuntan magnetite deposit. a – granular garnets replaced by quartz + actinolite veins. b – garnet skarn cut by magnetite vein. c – diopside + magnetite replaced by quartz + epidote. d – garnet + magnetite cut by quartz + epidote vein. e – massive magnetite + chlorite ore cut by albite + actinolite vein. f – disseminated magnetite in chlorite skarn cut by quartz + magnetite + pyrite vein. g – massive magnetite + chlorite ore cut by quartz + epidote + garnet + pyrite vein. h – disseminated magnetite in chlorite skarn cut by quartz + epidote+ pyrite vein. i – garnet replaced by magnetite + quartz with concentric zoning rims (plane-polarized light). j – garnet enclosed by magnetite, both cut by quartz + epidote vein (plane-polarized light). k – diopside replaced by magnetite (cross-polarized light). l – magnetite replaced by quartz + pyrite + chalcopyrite (reflected light).

Abbreviations: Grt – garnet; Act – actinolite; Qtz – quartz; Mag – magnetite; Di – diopside; Ep – epidote; Py – pyrite; Chl – chlorite; Ab – albite; Ccp – chalcopyrite.

Fig. 8 187Re-187Os isochron plot for pyrite from the Hongyuntan magnetite deposit.

Fig. 9 Photomicrographs showing different types of fluid inclusions in the Hongyuntan magnetite deposit. a – coexisting Type Ia, Ib, IIa, and IV fluid inclusions distributed in garnet. b – Type Ib fluid inclusion containing an opaque daughter mineral in garnet. c – Type Ib fluid inclusion in garnet. d – Type II fluid inclusion in garnet. e – Type II fluid inclusion containing multiple daughter minerals in garnet. f – Type Ia fluid inclusion in tremolite. g – coexistence of representative Type Ia and Ib fluid inclusions in quartz. h – Type IV fluid inclusion in quartz. i – coexisting Type Ia, Ib and II fluid inclusions in quartz, where the Type II solid contains an opaque and a transparent daughter mineral. j – Type II fluid inclusion in quartz, where the daughter mineral most is likely halite, as inferred from its cubic shape. k – coexisting vapor-rich Type II and liquid–rich Type II fluid inclusions in quartz. l – Type III fluid inclusion containing multiple daughter minerals in quartz. m – Type II fluid inclusion containing a vapor, a liquid, a solid, and a hematite daughter mineral. n – Type Ia fluid inclusion in epidote. o – Type Ia fluid inclusion in calcite. Abbreviations: L – liquid phase; V – vapor phase; S – daughter phase; Op – opaque phase; Tr – transparent phase; Hem – hematite.

Fig. 10 Frequency histogram of the total homogenization temperatures (Th) and salinities of fluid inclusions in different stages in the Hongyuntan magnetite deposit.

Fig. 11 δ18O vs δD plot for the ore-forming fluids in the Hongyuntan magnetite deposit (magmatic and metamorphic water after Taylor, 1974; meteoric water line after Craig, 1961; boiling and mixing trend after Shmulovich et al., 1999).

Fig. 12 Magnetite δ18O values from the Hongyuntan magnetite deposit compared with those of other iron deposits (modified from Jonsson et al., 2013). The range for ortho–magmatic magnetites is from Taylor, 1967. The Dunde and Beizhan data are from Li, 2012. The Chagangnuoer data are from Hong, 2012; Zhang, 2013b. The Zhibo data are from Wang, 2013; Jiang, 2014. The Kiruna and El Laco data are from Jonsson et al., 2013; Nyström et al., 2008. The Ningwu data are from Yuan et al., 1997. Other data are from Hong, 2012; Zhao et al., 2012.

Fig. 13 Histogram of the sulfur isotopic compositions of pyrites in the Hongyuntan magnetite deposit.

Fig. 14 Diagram for estimating the pressures of the boiling fluid inclusions from the Hongyuntan magnetite deposit (modified from Bouzari and Clark, 2006).

Fig. 15 Plot of the homogenization temperature (°C) vs. salinity (% NaCleqv) of the fluid inclusions from the Hongyuntan magnetite deposit.

Fig. 16 Schematic illustration of the proposed genetic model of the Hongyuntan magnetite deposit

Table Titles Table 1 Re-Os isotopic compositions of pyrite samples from the Hongyuntan magnetite deposit

Table 2 Microthermometric data of different types of fluid inclusions in the Hongyuntan magnetite deposit

Table 3 Hydrogen and oxygen isotopic analyses and the calculated isotopic compositions of fluids from the Hongyuntan magnetite deposit

Table 4 Isotope fractionation equilibrium formulas used in the mineral–water system (1000lnα = A*106/T2 + B*103/T + C) (Kelvin (K) for T)

Table 5 Analytical results of the sulfur isotopic compositions of pyrites from the Hongyuntan magnetite deposit

Fig. 1

Fig. 2

Fig. 3

Fig. 4

Fig. 5

Fig. 6

Fig. 7

Fig. 8

Fig. 9

Fig. 10

Fig. 11

Fig. 12

Fig. 13

Fig. 14

Fig. 15

Fig. 16

Table 1 Re-Os isotopic compositions of pyrite samples from the Hongyuntan magnetite deposit Sampl

Weigh

Re(n

e

t(g)

g/g)

H692

0.600

68.55

4-16

36

H692

0.200

4-20

20

H692

0.200

4-23

16

H692

0.200

4-24

22

H692

0.200

3-4

35

66.56

16.96

72.87

7.31

2σ

C

Os(n

2σ

g/g)

187

Re(n

2σ

g/g)

0.

0.063

0.00

51

00

049

0.

0.042

0.00

49

89

033

0.

0.015

0.00

13

15

015

0.

0.071

0.00

54

55

054

0.

0.041

0.00

05

84

032

43.08

187

Os(n

g/g) 0.

0.2375

32 41.83

0.

0.

0.2175

0.

0.0577

0. 03

2

187

88

Os

σ

88

5235

5

28.80

0.00

0.00

0.2537

0.00

7489

0.00 02

7

5291

6

5

38.89

8. 5

0. 07

28.13

0. 18

27.09

0 844.8

0. 06

1 4906

2σ

Os

6

19 0.0268

Os/1

4

04

34 14.59

Re/1

16

08 45.80

0.00

187

18

31 10.66

2σ

0. 04

4.93

0. 01

Note: COs is common Os; 187Re decay constant = 1.666×10-11 a-1 with a relative uncertainty of ± 0.31% (Smoliar et al., 1996).

Table 2 Microthermometric data of different types of fluid inclusions in the Hongyuntan magnetite deposit Stag

Host

Type(

e

mineral

n)

Th (°C)

Tm, ice (°C)

Tm, cla (°C)

Tm, s (°C)

Salinity (wt. %. NaCl eqv)

Stag

Garnet

Ia(5)

eI

Rang

Mea

Rang

Mea

Rang

Mea

Rang

Mea

Rang

Mea

e

n

e

n

e

n

e

n

e

n

466 −

476.8

−

− 8.7

10.2

12.8

482

11.2

−

to −

15.2

6.8 Ib(24)

446

>549.

−

− >60

1

14.2

− 9.3

9.2 −

12.5

16.1

to −

0

6.0 IIa(6)

546 −

567.3

577

258 −

271.

35.2

279

1

−

36.1

36.6 IIb(10)

486 −

515.5

551

486 −

515.

59.5

551

5

−

61.9

63.2 Stag

Tremoli

e II

te

Ia(16)

249 −

275.4

− 7.2

− 6.0

7.4 −

to −

308

9.2

10.7

4.7 Stag

Quartz

Ia(4)

e III

412 −

438.3

462

−

− 9.1

12.4

10.2

−

to −

14.1

12.7

7.4 Ib(31)

325 −

449.3

− 9.6

− 5.7

5.1 −

to −

540

8.6

13.5

3.1 IIb(9)

381 −

441.6

496

278 −

417.

36.5

480

4

−

49.9

57.1

Stag

Quartz

III(1)

438

438

Ia(65)

223 −

262.3

e IV

322

4.2 −

4.2

− 7.3

11.6

10.2

10.2

4.2 −

10.7

15.6

to − 2.5 IIa(14)

243 −

270.7

318

176 −

255.

30.7

276

8

−

35.1

36.4 IIc(14)

250 − 269

258.8

251 −

259.

34.7

270

6

−

35.3

35.9 Epidote

Ia(56)

215 −

238.4

259

−

− 6.0

10.6

6.2 −

9.1

14.6

to − 3.8 Calcite

Ia(35)

158 − 286

203.6

− 7.7

−3.1

to −

1.5 −

5

11.3

0.9 Ia – Liquid–rich inclusions. Ib – Vapor–rich inclusions. IIa – inclusions showing the characteristics of homogenization by bubble disappearance after halite dissolution. IIb – inclusions are confined by halite dissolution after bubble disappearance during heating. IIc – inclusions are homogenized to the liquid phase with the same temperature for the dissolution of the daughter minerals and the disappearance of bubbles. III – CO2 multi–phase inclusions. Th – homogenization temperature. Tm, ice – temperature of final ice melting. Tm, cla – the melting temperatures of clathrate. Tm, s – halite dissolution temperature. ω – salinity. (n) – number of inclusion tests. All temperatures in ° C. Salinity is expressed as wt.% NaCl equivalent.

Table 3 Hydrogen and oxygen isotopic analyses and the calculated isotopic compositions of fluids from the Hongyuntan magnetite deposit No

Sample No.

Mineral

δDV-SMOW(‰)

δ18OV-SMOW(‰)

δ18Ofluid(‰)

1

H6924–19

Garnet

– 88.5

3.8

6.7

2

ZK502–14

Garnet

– 83.3

3.5

6.4

3

H6923–2

Tremolite

– 117.3

3.6

3.3

4

H6924–27

Tremolite

– 112.6

6.0

5.7

5

H6924–26

Epidote

– 65.1

5.7

3.8

6

ZK503–9

Epidote

– 61.9

3.4

1.5

7

ZK503–10

Epidote

– 58.5

3.8

1.9

8

H6924–20

Quartz

– 85.5

8.6

4.8

9

H6926–3

Quartz

– 83.2

7.7

3.9

10

H6924–23

Quartz

– 87.9

8.6

0.1

11

H6923–1

Massive magnetite

–

– 0.5

7.7

12

H6923–2

Massive magnetite

–

– 0.6

7.6

13

H6924–6

Banded magnetite

–

– 0.4

7.8

14

H6924–9

Massive magnetite

–

0.2

8.4

15

H6924–19

Massive magnetite

–

– 0.2

8.0

16

H6924–31

Massive magnetite

–

4.1

12.6

17

H6924–25

Massive magnetite

–

1.9

10.1

18

H6924–28

Massive magnetite

–

3.2

11.4

19

Garnet

– 76.7

7.8

10.7

20

Garnet

– 104.4

5.4

8.3

21

Garnet

– 106.1

6.3

9.2

22

Garnet

– 107.9

4.6

7.5

23

Epidote

– 96.8

6.4

4.5

24

Epidote

– 99.4

4.9

3.0

25

K–feldspar

– 73.2

11.7

3.7

26

K–feldspar

– 97.7

11.5

3.5

27

Quartz

– 95.1

10.3

–1.3

28

Quartz

– 106.8

10.3

1.8

29

Quartz

– 98.3

9.9

1.4

30

Quartz

– 96.2

9.3

0.8

31

Quartz

– 115.2

10.6

2.1

32

Quartz

– 103.2

11.9

3.4

33

Quartz

– 96.2

11.5

3.0

.

Reference

This study

Zheng, 2015

Table 4 Isotope fractionation equilibrium formulas used in the mineral–water system (1000lnα = A*106/T2 + B*103/T + C) (Kelvin (K) for T) Minerals

Fractionation formula

Garnet–water

6

Reference

2

3

1000lnα = 3.76 * 10 /T – 9.05 * 10 /T + 2.52 2

3

Zheng et al., 1993a

6

2

3

Zheng et al., 1993b

6

2

3

Zheng et al., 1993a

6

2

3

Zheng et al., 1993a

1000lnα = 4.48 * 10 /T – 4.77 * 10 /T + 1.71

Quartz–water

1000lnα = 3.95 * 10 /T – 8.28 * 10 /T + 2.38

Tremolite–water

1000lnα = 4.05 * 10 /T – 7.81 * 10 /T + 2.29

Epidote–water

1000lnα = 4.32 * 10 /T – 6.27 * 10 /T + 2.00

K–feldspar–water Magnetite–water

Zheng et al., 1993a

6

6

2

3

1000lnα = 2.88 * 10 /T – 11.36 * 10 /T + 2.89

Zheng, 2000

Table 5 Analytical results of the sulfur isotopic compositions of pyrites from the Hongyuntan magnetite deposit No.

Sample No.

Description

Mineral

δ34SCDT (‰)

1

H6923–1

Massive ore

Pyrite

– 0.6

2

H6923–2

Massive ore

Pyrite

– 1.0

3

H6923–6

Massive ore

Pyrite

1.0

4

H6924–8

Disseminated ore

Pyrite

1.3

5

H6924–11

Vein ore

Pyrite

– 0.3

6

H6924–12

Massive ore

Pyrite

– 0.9

7

H6924–17

Banded ore

Pyrite

– 2.2

8

H6924–18

Massive ore

Pyrite

– 3.2

9

H6924–20

Vein ore

Pyrite

4.7

10

H6924–21

Pyrite vein

Pyrite

– 3.3

11

H6924–22

Massive ore

Pyrite

1.0

12

H6924–24

Quartz–pyrite vein

Pyrite

– 1.3

13

H6924–26

Quartz–pyrite vein

Pyrite

– 0.4

14

ZK503–6

Quartz–pyrite vein

Pyrite

– 1.1

15

HYT–76–2

Massive ore

Pyrite

– 3.1

16

HYT–77

Massive ore

Pyrite

– 3.2

17

HYT–113

Breccia ore

Pyrite

1.4

18

HYT–132

Breccia ore

Pyrite

– 2.4

19

HYT–86

Massive ore

Pyrite

– 1.9

20

HYT–129

Breccia ore

Pyrite

– 3.8

21

HYT–1

Banded ore

Pyrite

– 0.4

22

HYT–86

Banded ore

Pyrite

– 0.8

23

HYT–114

Breccia ore

Pyrite

4.3

24

HYT–117

Massive ore

Pyrite

– 1.3

25

HYT–148

Banded ore

Pyrite

1.2

26

TS1055

Massive ore

Pyrite

0.4

27

TS1061

Massive ore

Pyrite

0.4

Reference

This study

Zheng, 2015

Zhang et al., 2012

28

TS1062

Massive ore

Pyrite

0.3

29

TS1066

Massive ore

Pyrite

– 0.1

30

TS1070

Mineralized tuff

Pyrite

1.6

31

HZ–21

Pyrite vein

Pyrite

– 1.8

32

HZ–29

Pyrite vein

Pyrite

– 0.3

33

HZ–46

Pyrite vein

Pyrite

– 0.2

34

HZ–471

Pyrite vein

Pyrite

– 1.0

35

HZ–268

Pyrite vein

Pyrite

– 1.7

36

HZ–434

Pyrite vein

Pyrite

– 0.6

37

HZv167

Pyrite vein

Pyrite

– 1.1

38

HZ–389

Massive ore

Pyrite

0.7

39

HZ–150

Massive ore

Pyrite

1.2

40

HZ–162

Mineralized tuff

Pyrite

1.4

41

HZ–105

Pyrite vein

Pyrite

– 0.6

42

HZ498

Pyrite in volcanics

Pyrite

– 2.0

43

644

Pyrite in volcanics

Pyrite

– 2.5

44

645

Pyrite in volcanics

Pyrite

– 3.5

Wang, 1981

Ren, 1985

Highlights ► The Hongyuntan magnetite deposit, NW China, formed at 324 ± 31 Ma. ► Four types of fluid inclusions can be identified during the metallogenic stages of the Hongyuntan deposit. ► The temperatures and salinities of the ore-forming fluids fluctuate from the early to the late stages. ► The ore-forming fluids were magmatic in origin and mixed with seawater during the late stage. ► Fluid mixing and fluid boiling were main potential mechanisms for the deposition of magnetite.

Graphic abstract