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Article

Evolution of the Ore-Bearing Fluid of Alin Sb–Au Orebodies in Shuixie Cu–Co Orefield, SW China: Constraints on the Rare Earth Element and Trace Element Components of Auriferous Pyrite and Host Rock

by
Guo Li
1,
Shanshan Ru
1,2,*,
Chuandong Xue
1,* and
Wei Wang
1
1
Department of Earth Sciences, Kunming University of Science and Technology, Kunming 650093, China
2
City College, Kunming University of Science and Technology, Kunming 650093, China
*
Authors to whom correspondence should be addressed.
Minerals 2025, 15(5), 491; https://doi.org/10.3390/min15050491
Submission received: 17 March 2025 / Revised: 23 April 2025 / Accepted: 2 May 2025 / Published: 6 May 2025
(This article belongs to the Special Issue Structure and Origin of Gold Mineralization: From Primary to Placer Gold Deposits)
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Abstract

:
The Shuixie Cu–Co polymetallic orefield, located in western Yunnan Province (southeastern margin of the Qinghai–Tibet Plateau), is renowned for its Cu–Co mineralization. A recent resource reassessment identified Sb–Au and Cu–Co–Bi (Sb–Au) orebodies as genetically associated with primary Cu–Co mineralization. The mineralization characteristics and microscopic observations indicate that gold mineralization in the Sb–Au orebodies follow a pulsating fluid injection model. The model includes four pulses: (1) euhedral gold-poor pyrite (PyI1) precipitation; (2) margin-parallel growth of gold-rich pyrite (PyI2) on PyI1; (3) continued growth of gold-rich pyrite (PyI3) along PyI2; and (4) outermost concentric gold-rich pyrite (PyI4) formation. This study examined gold-bearing pyrite in orebodies and host rocks. In situ laser ablation inductively coupled plasma mass spectrometry (LA–ICP–MS) analysis of pyrite and inductively coupled plasma mass spectrometry (ICP–MS) whole-rock trace element analysis were conducted to track the ore-forming fluid evolution. Compared with CI chondrite, pyrites from all pulses were enriched in LREEs over HREEs. The pyrite REE distribution curves exhibited right-skewed patterns, reflecting LREE enrichment. The Hf/Sm, Nb/La, and Th/La ratios were generally below 1, indicating high-field-strength element depletion. These results suggest a Cl-rich, F-poor ore-forming fluid. The pyrite trace elements showed enrichment in the chalcophile elements (e.g., Cu and Pb) and exceptionally high Bi levels compared with the continental crust. The chalcophile elements (e.g., Zn and Cd) were depleted, whereas iron-group elements (e.g., Co) were enriched and Ni was depleted. The pyrite δCe values (0.87–1.28, mean = 1.01) showed weak anomalies, indicating a reducing ore-forming environment. The δEu values of pyrite during pulses 1 to 4 ranged widely, from 0.2–3.01 (mean of 1.17), 0.27–1.39 (0.6), and 0.41–1.40 (0.96) to 0.4–1.36 (0.84), respectively, suggesting an initial temperature decline and subsequent increase in the ore-forming fluid. Significant variations were found in the Y/Ho, Zr/Hf, and Nb/Ta ratios across pulses, indicating the potential involvement of high-temperature hydrothermal fluids or late-stage alteration during ore formation. The Y/Ho ratio of pyrite overlapped most closely with that of the continental crust of China, indicating a close relationship between the ore-forming fluids and the crust.

1. Introduction

The Sanjiang Composite Orogenic Belt (SCOB), located on the southeastern edge of the Tibetan Plateau, is a key part of the global Tethyan metallogenic domain. Its Cenozoic tectonic–magmatic activities and polymetallic mineralization have long been a focus of international mineral deposit studies [1,2,3] (Figure 1). The Lanping Basin, a mid-Cenozoic intracontinental strike–slip basin in the southern segment of the orogenic belt, records the distant effects of the India–Eurasia plate collision. Additionally, its complex hydrothermal mineralization system has produced polymetallic resources, including copper, lead, zinc, antimony, and gold, making it a natural laboratory for exploring the metallogenic mechanisms of continental collision zones [1,3]. The recent discovery of the Alin gold deposit in the western part of the basin revealed the potential for the superposition of orogenic gold deposits and shallow low-temperature hydrothermal systems. This provides a key case for understanding the metallogenic diversity along the eastern Tethyan margin [4,5]. Despite extensive global research on orogenic gold deposits (e.g., craton greenstone belts) and shallow, low-temperature hydrothermal systems (e.g., the circum-Pacific belt), the coupling mechanisms between the two in intraplate strike–slip environments remain controversial [6]. The morphology, crystal features, and trace element content of pyrite are closely linked to the physicochemical properties of ore-forming fluids including sulfur fugacity, temperature, and cooling rate [7,8,9,10]. Hydrothermal pyrite, a direct sample for studying mineralization fluids, reflects the properties of ore-forming fluids and their material sources through its REE distribution, trace element content, and ratios [11,12,13,14].
The uniqueness of the Alin gold deposit lies in the unclear source of the mineralizing fluid. Existing isotopic data only preliminarily suggest the involvement of magma fluids [5], but lack quantitative constraints on crust–mantle interactions. The mechanism of element occurrence is unclear. As the primary gold-bearing mineral, pyrite’s rare earth element (REE) distribution pattern and trace element ratios have not yet been revealed in terms of their role in fluid evolution. This study used micro-region in situ analytical techniques to systematically analyze the distribution patterns of REE and trace element combinations in pyrite from the Alin gold deposit. The aims were to (1) reveal the material sources and evolution of metallogenic fluids, filling the geochemical gap in fluid studies in the western margin of the Lanping Basin, and (2) establish the coupling relationships between REEs and elements such as Au, As, and Sb, providing new evidence for understanding the extraordinary element enrichment mechanism in collision-zone gold deposits.

2. Geological Setting

2.1. Regional Geological Background

The Shuixie Cu–Co polymetallic orefield is situated at the southwestern edge of the Meso–Cenozoic Lanping Basin, on the southeastern Xizang–Qinghai Plateau. It is an important part of the Sanjiang Composite Orogenic Belt (SCOB) (Figure 1a). The tectonic framework of the Lanping Basin is shaped by the Baoshan–Tengchong Block to the west, the Yangtze Platform to the east, and the Meso–Cenozoic Simao Basin to the south. Permian to Cretaceous sedimentary strata are exposed within and along the margins of the basin (Figure 1b). Geological surveys have shown that the SCOB, including the Lanping Basin, evolved through several phases: the Paleo–Tethys continental marginal arc basin, the Neo–Tethys back-arc foreland basin, and the ongoing Cenozoic intracontinental collisional belt between the Indian and Eurasian continents. These processes led to the widespread development of Triassic arc magmatic belts, Mesozoic red-bed basins, and Eocene to Oligocene alkali-rich porphyry belts [1,2,3,15,16,17]. These processes are associated with tectonic, sedimentary, metamorphic, and magmatic activities, creating a distinct polymetallic metallogenic zone.
The Yongping–Weishan Cu–Co–Sb–Au polymetallic mineralization belt lies in the southwest of the Lanping Basin and is rich in Cu, Co, Sb, Au, and associated minerals such as siderite, barite, and fluorite (Figure 1b). Notable deposits include the Baiyangchang Cu–Ag–Pb–Zn deposit in Yunlong County, the Gongguo Cu–Co–Au, Reshuitang Cu–Au–Pb–Zn, Qingyangchang and Changjie Cu–Co–Ag, and Laoyingpo and Shuixie Cu–Co–Au deposits in Yongping County, and the Tianbaxin–Bijianshan–Dugupi Sb–fluorite and Zhacun Au–Sb–Pb–Zn deposits in Weishan County (Figure 1b). Himalayan–aged (38–34 Ma) potassic intrusive bodies are associated with hornfels development in the surrounding sandy siltstones and small-scale Cu–Pb–Zn–Au polymetallic mineralization at the Zhuopan pluton in Yongping County, Lianhuashan and Gejiumu plutons in Weishan County, and Guduo and Jiedingshan plutons in Midu County (Figure 1b).
The Shuixie orefield is part of the Yongping–Weishan region, a prospective area for gold, copper, cobalt, antimony, mercury, arsenic, iron, and rare earth elements. Over more than fifty years of geological exploration, more than ten medium-to small-sized Cu–Co and Cu deposits have been discovered including the Shuixie Jie, Xiaotuanshan, Caiyuanzi, Shangbielie, Qingtian, Zhifanghe, and Alin deposits [4,5].

2.2. Deposit Geology

The Alin deposit lies on the eastern flank of the Alin anticline. The deposit contains outcrops of the Upper Triassic Maichuqing Formation (T3m), Lower Jurassic Yangjiang Formation (J1y), and Middle Jurassic Huakaizuo Formation (J2h) (Figure 2). A detailed stratigraphic description is provided in Figure 2. The main faults in the area are F1, F2, F3, F4, F5, and F6. Of these, F2 and F4 are NNE-trending, F1 is NNW-trending, and F3 is EW-trending. F1 is the primary controlling fault of the Alin deposit, with a total length of 1.7 km, and is a reverse fault. F2, the second largest fault after F1, also controls mineralization in the area. It is displaced by F3 in the northern part and has a total length of 1.8 km. F3 is a right–lateral reverse fault, 2.9 km long, extending from Mingcaoshan in the west to beyond the mining area in the east [18]. No igneous rock outcrops are observed. The Alin anticline is the largest fault-derived structure in the mining area. It is a tilted syncline, with its hinge rising toward the southwest and extending near the western side of the Alin mineralization belt. The Alin syncline lies southeast of the Alin anticline, roughly parallel to it. Gold is currently mainly mined from the 1760 and 1700 tunnels at the Alin deposit. The orebodies occur in vein-like and layered forms, primarily in the T3m, which consists of grayish-white feldspar-quartz sandstone and gray-black to black carbonaceous mudstone in fault breccia zones. It also occurs in the J2h and J1y, containing purple to purplish-red, dark medium-grained sandstone, quartz sandstone, and feldspar-quartz sandstone. The mineralized breccia consists mainly of feldspar-quartz sandstone fragments and fault clay (Figure 3). The thickness of the orebodies ranges from 6 to 11.98 m, with a length of up to 130 m. The ore exhibits dense, massive, banded, impregnated, vein-like, networked, and breccia structures, with euhedral, subhedral, and anhedral grains as well as textures of replacement, dissolution, and cataclasis (Figure 3). The gold (Au) grade ranges from 0.26 to 36.35 g/t, with a maximum of 120.0 g/t and an average of 4.02 g/t [18]. This study focused on the Sb–Au orebodies in the 1700-, 1715-, and 1760-level tunnels (Figure 3). The main metal minerals are pyrite, arsenopyrite, stibnite, sphalerite, and trace chalcopyrite. The gangue minerals are primarily barite, rhodochrosite, dolomite, calcite, quartz, feldspar, and biotite (Figure 4). Secondary minerals include limonite and azurite.
The Shuixie Alin Sb−Au orebodies follow a pulsating fluid injection model, based on field geological observations, petrographic identification, and BSE morphology analysis. The mineralization consists of five pulses, four of which exhibit gold mineralization, characterized by alternating precipitation of pyrite and arsenopyrite: (I) pulses 1, with barite (minor)–pyrite–arsenopyrite; (II) pulses 1, with pyrite—arsenopyrite; (III) pulse 3, with pyrite–arsenopyrite– stibnite (minor); (IV) pulse 4, with pyrite–arsenopyrite–quartz; and (V) pulse 5, characterized by a large amount of stibnite [5].
During pulse 1, barite (trace), pyrite (PyI1), and arsenopyrite (ApyI1) are precipitated, with weak gold mineralization. Arsenopyrite and pyrite in the host rock occur as euhedral−subhedral crystals (Figure 4c–e). In the orebodies, they predominantly appear as subhedral crystals, often with a pyrite ring–core structure (Figure 4f–j). During pulse 2, pyrite (PyI2) and arsenopyrite (ApyI2) crystallize, showing stronger gold mineralization. Without arsenopyrite and pyrite crystals from this pulse can be observed in the host rock. In the orebodies, pyrite mainly appears as euhedral or fragmented crystals, growing along arsenopyrite (ApyI2) and forming the primary ring in a ring−core pyrite structure (Figure 4f,h–j). Pulse 3 precipitates pyrite (PyI3), arsenopyrite (ApyI3), and stibnite (trace), with the strongest gold mineralization. These pyrites are identified only in the orebodies, mainly as euhedral pyrite, cut or intergrown with arsenopyrite (Figure 4f and g), forming secondary rings in a ring–core pyrite structure (Figure 4h–j). The final pulse of gold mineralization (IV) consists of pyrite (PyI4), arsenopyrite (ApyI4), and quartz, with relatively strong mineralization. They mainly occur as the outermost rings of the ring–core structure, with crystal forms similar to those of pulses 2 and 3 (Figure 4h–j). For further details on the ore mineral features, see Li et al. [5].

3. Sample and Analytical Methods

The samples from the Alin Sb–Au orebodies in this study were collected during the horizontal mining stage. Systematic samples from both the orebodies and host rock were collected. The samples were taken from the 1700 tunnel. All samples were cut into thin sections for microscopic observation and backscattered electron (BSE) imaging. The pyrite mineralization stages in different samples were divided into pulses. Each pulse of pyrite was delineated with an oil-based pen, followed by microscopic photography. Finally, in situ LA–ICP–MS analysis was conducted. The host rock samples were ground and analyzed for whole-rock trace elements and REEs.

3.1. LA-ICP-MS Analysis

The trace element analysis of pyrite was conducted by LA–ICP–MS at Sample Solution Analytical Technology Co. Ltd., Wuhan, China. Detailed operating conditions for the laser ablation system and the ICP–MS instrument and data reduction were the same as described by [19]. Laser sampling was performed using a GeolasPro laser ablation system that consisted of a COMPexPro 102 ArF excimer laser (wavelength of 193 nm and maximum energy of 200 mJ) and a MicroLas optical system. An Agilent 7700e ICP–MS instrument was used to acquire ion-signal intensities. Helium was applied as a carrier gas. Argon was used as the make-up gas and mixed with the carrier gas via a T-connector before entering the ICP. A “wire” signal smoothing device is included in this laser ablation system [20]. The spot size and frequency of the laser were set to 32 µm and 5 Hz, respectively. The trace element compositions of the sulfides were calibrated against various reference materials (NIST 610 and NIST 612) without using an internal standard [21]. The sulfide reference material of MASS–1 (USGS) was used as the unknown sample to verify the accuracy of the calibration method. Each analysis incorporated a background acquisition of approximately 20–30 s, followed by 50 s of data acquisition from the sample. An Excel–based software ICPMSDataCal 11.8 was used to perform offline selection and integration of the background and analyzed signals, time-drift correction, and quantitative calibration for trace element analysis [21].

3.2. Trace Element Analysis of Whole Rock

Trace element analysis of the whole rocks was conducted on an Agilent 7700e ICP–MS at Wuhan Sample Solution Analytical Technology Co. Ltd., Wuhan, China. The detailed sample-digesting procedure was as follows: (1) Sample powders (200 mesh) were placed in an oven at 105 °C and dried for 12 h; (2) 50 mg of the sample powder was accurately weighed and placed in a Teflon bomb; (3) 1 mL of HNO3 and 1 mL of HF were slowly added into the Teflon bomb; (4) the Teflon bomb was put into a stainless steel pressure jacket and heated to 190 °C in an oven for >24 h; (5) after cooling, the Teflon bomb was opened, placed on a hotplate at 140 °C, and evaporated to incipient dryness, then, 1 mL of HNO3 was added and evaporated to dryness again; (6) 1 mL of HNO3, 1 mL of MQ water, and 1 mL of 1 ppm internal standard solution In were added, and the Teflon bomb was resealed and placed in the oven at 190 °C for >12 h; and (7) the final solution was transferred to a polyethylene bottle and diluted to 100 g via the addition of 2% HNO3.

4. Results

4.1. Rare Earth Elements

The in situ rare earth element (REE) results for pyrite and the host rock are presented in Table S1. The measured results were standardized using the CI chondrite standard [22], as illustrated in Figure 5. Significant differences in the REE geochemical characteristics were observed between the pyrite of each pulse and the host rocks. (I) The ƩREE content of the host rocks ranged from 260.26 × 10−6 to 773.11 × 10−6, with an LREE/HREE ratio of 2.92 to 3.85, indicating LREE enrichment. The (La/Yb)N values ranged from 8.14 to 11.36, (La/Lu)N from 8.19 to 10.92, and (Ce/Yb)N from 5.84 to 8.47. The LREE−HREE fractionation was minimal, with (La/Sm)N values ranging from 3.43 to 4.05, and (Gd/Yb)N from 1.47 to 1.80. The LREEs exhibited slightly more differentiation than the HREEs. The δEu values ranged from 0.68 to 0.82, with an average of 0.73, and the δCe values ranged from 0.92 to 1.00, with an average of 0.96 (Figure 5e, Table S1). (II) For pulse 1 pyrite, the ƩREE content ranged from 4.49 × 10−6 to 21.48 × 10−6 with an LREE/HREE ratio of 0.96 to 3.54, indicating LREE enrichment. The (La/Yb)N values ranged from 3.36 to 121.21, (La/Lu)N from 1 to 43.62, and (Ce/Yb)N from 1.85 to 91.12. A significant LREE−HREE fractionation was observed, with (La/Sm)N between 0.52 and 4.29, and (Gd/Yb)N between 1.12 and 15.71. The LREEs exhibited less differentiation than the HREEs. The δEu values ranged from 0.20 to 3.01, with an average of 1.17, and the δCe values ranged from 0.69 to 1.10, with an average of 0.99 (Figure 5a, Table S1). (III) For pulse 2 pyrite, the ƩREE range was 7.8 × 10−6~57.94 × 10−6. Compared with pulse 1, both the ƩREE content and range increased. The LREE/HREE ratio ranged from 0.87 to 6.73, with the LREEs generally being more enriched. The (La/Yb)N value ranged from 1.46 to 28.53, (La/Lu)N from 1.15 to 23.85, and (Ce/Yb)N from 0.39 to 23.21. A large fractionation between the LREEs and HREEs was observed, with (La/Sm)N values between 2.15 and 4.25, and (Gd/Yb)N between 0.15 and 4.91. The variation in the LREEs was smaller than in the HREEs. The δEu value ranged from 0.27 to 1.39, with an average of 0.60, and the δCe value ranged from 0.35 to 1.52, with an average of 0.99 (Figure 5, Table S1). (IV) For pulse 3 pyrite, the ƩREE content ranged from 18.77 × 10−6 to 84.40 × 10−6. Compared with pulse 2, both the ƩREE content and range increased. The LREE/HREE ratio ranged from 1.7 to 5.55, with the LREEs being more enriched. The (La/Yb)N value ranged from 3.0 to 16.55, (La/Lu)N from 2.83 to 14.51, and (Ce/Yb)N from 2.83 to 14.51. A significant fractionation between the LREEs and HREEs was observed, with (La/Sm)N values ranging from 1.92 to 5.85 and (Gd/Yb)N from 0.46 to 1.42. The variation in LREEs was greater than that in HREEs. The δEu value ranged from 0.41 to 1.40, with an average of 0.96, and the δCe value ranged from 0.85 to 1.72, with an average of 1.28 (Figure 5c, Table S1). (V) For pulse 4 pyrite, the ƩREE content ranged from 3.19 × 10−6 to 54.79 × 10−6. Compared with pulse 3, both the ƩREE content and ranges decreased. The LREE/HREE ratio ranged from 1.65 to 6.22, with the LREEs being more enriched. The (La/Yb)N value ranged from 2.22 to 17.59, (La/Lu)N from 2.25 to 9.75, and (Ce/Yb)N from 2.35 to 12.59. Significant fractionation between the LREEs and HREEs was observed, with (La/Sm)N values ranging from 2.18 to 4.08 and (Gd/Yb)N from 0.45 to 1.19. The variation in LREEs was greater than that in the HREEs. The δEu value ranged from 0.40 to 1.36, with an average of 0.84, and the δCe value ranged from 0.79 to 1.49, with an average of 1.13 (Figure 5d, Table S1).

4.2. Trace Elements

The trace element concentrations are presented in Table S1. The results were standardized using the continental crust standard [23], as shown in Figure 6. The trace elements were classified into iron-group elements (V, Cr, Co, and Ni), rare elements (Li, Sc, Y, Zr, Nb, Hf, etc.), radioactive elements (Th and U), tungsten–molybdenum group elements (Mo and W), and sulfur-loving elements (Cu, Zn, Cd, Tl, Pb, etc.) according to Zawaritsky’s scheme [24,25]. The iron-group elements in pyrite varied significantly from pulse 1 to 4. V and Cr were depleted at all stages, with depletion being smaller during pulses 1 to 3 compared with pulse 4. Co and Ni were enriched in all pulses, particularly in pulse 1. However, Co was depleted at some points in pulses 1, 2, and 3, while Ni was mostly depleted in pulses 3 and 4. The Co/Ni ratio ranged from 0.18 to 30.35 across all pulses, with nearly half of the data points below 1, and 20 out of 27 points below 5. The rare element concentrations in pyrite were depleted in all pulses. The radioactive elements in pyrite were generally depleted across all pulses. The tungsten–molybdenum group elements exhibited significant variation in all pulses. Mo was depleted, while W was enriched in all pulses, with the enrichment factors (K) mostly greater than 1. Chalcophile elements were the primary ore-forming elements. Cu and Pb were enriched in all pulses, with enrichment factors (K) typically between 1 and 10. Tl was depleted in pulse 1 but enriched in the other pulses, whereas Zn and Cd were depleted across all pulses. The iron-group elements in the host rocks showed little variation across the orebodies. V, Cr, Co, and Ni were depleted in both the upper and lower parts, but their concentrations were similar to those of the continental crust standard. The rare element concentrations showed minor depletion, similar to those of the iron-group elements. The radioactive elements were generally enriched, with the enrichment factors (K) just above 1. Due to experimental limitations, the tungsten–molybdenum group elements (Mo and W) were not tested in the host rocks. The enrichment–depletion characteristics of the chalcophile elements (Cu, Pb, Tl, Zn, and Cd) were similar to those observed in pyrite. Notably, Bi exhibited extraordinary enrichment in both the surrounding rocks and pulses, with their enrichment factors (K) ranging from 10 to 100, and even exceeding 100.

5. Discussion

5.1. The Composition of Ore-Forming Fluid

The average rare earth element (REE) content of each pulse of pyrite is shown in Figure 7. The average ƩREE content increased from pulse 1 (12.78 × 10−6) to pulse 2 (19.63 × 10−6), peaked at pulse 3 (58.59 × 10−6), and then decreased at pulse 4 (21.03 × 10−6). Both individual data points and average values indicated that the ƩREE content was much lower than the host rock’s average (528.21 × 10−6). All pyrite pulses showed enrichment of the LREEs over HREEs, with varying degrees of LREE-HREE fractionation. For instance, pulse 1 exhibited a large fractionation between the LREEs and HREEs, with smaller variation within the LREEs compared with that within the HREEs. In contrast, pulse 3 showed a greater fractionation between the LREEs and HREEs, with more variation within the LREEs than within the HREEs, likely due to temperature changes that alter the physicochemical properties of the hydrothermal fluids or the influence of external fluids. Our arsenopyrite As atom data indicated average precipitation temperatures of 240 °C, 220 °C, 235 °C, and 230 °C for pulses 1 to 4, respectively [5]. This is consistent with previous temperature measurements (142–230 °C) of quartz fluid inclusions from the Shuixie deposit [24]. The average ƩREE content in the host rock (528.21 × 10−6) showed a greater enrichment of the LREEs, with a small LREE-HREE fractionation and slightly more variation within the LREEs than in the HREEs.
Previous studies have shown that REE ions can form complexes with ions such as carbonate (CO32−), chloride (Cl−), fluoride (F−), nitrate (NO3−), and sulfate (SO42−), leading to compounds such as carbonates, sulfates, and chlorides. This represents a key form of REE occurrence in nature [25]. The REE distribution curves for pyrite from various pulses in the Alin Sb–Au orebodies were mostly right-skewed, indicating LREE enrichment (Figure 5). This suggests that the ore-forming fluids contain a significant amount of F− or Cl− [26,27,28]. High field strength elements (HFSEs), including HREEs, Nb, Ta, and Ce, have small ionic radii and high ionic charges, leading to strong ionic fields. The distribution of REEs and HFSEs is primarily controlled by their content in pyrite-forming fluids, with little influence from the crystal structure. If the ore-forming fluid is rich in Cl−, the Hf/Sm, Nb/La, and Th/La ratios are generally less than 1. In contrast, F−-enriched fluids with concentrated HFSEs show Hf/Sm, Nb/La, and Th/La ratios greater than 1 [27]. Table S1 shows that most of the Hf/Sm, Nb/La, and Th/La values for pyrite from the Alin gold deposit were less than 1, with only a few exceeding 1, suggesting that the ore-forming fluid is rich in Cl− and low in F− [24,25,26].
Each pulse of pyrite in the Alin Sb–Au orebodies was enriched in chalcophile elements such as Cu and Pb, and relatively depleted in Zn and Cd, showing an exceptionally high enrichment of Bi, while the iron-group elements were enriched in Co but depleted in Ni. The HFSEs, including Th, U, Ta, Nb, Zr, and Hf, were relatively depleted. The large ion lithophile elements (LILEs), such as Sr, Ba, and Rb, were also relatively depleted (Figure 5). The tungsten–molybdenum group elements (Mo and W) exhibited varying patterns of enrichment and depletion at different pulses. The trace element characteristics of pyrite suggest that the ore-forming fluid is enriched in sulfur-loving elements such as Cu, Pb, and Bi (Figure 6). The field observations, hand specimens, and microscopic analysis revealed that, in addition to pyrite, the ore minerals included small amounts of chalcopyrite, sphalerite, and galena, consistent with the trace element enrichment. In the host rock, the enrichment and depletion patterns of chalcophile elements, such as Cu, Pb, Tl, Zn, and Cd, mirrored those in pyrite. The iron-group elements showed minimal variation between the upper and lower parts of the ore body, with V, Cr, Co, and Ni depleted in both, though their depletion was not significant when compared with the standard crustal values. The REEs exhibited similar depletion patterns to those of the iron-group elements. The radioactive elements typically showed enrichment, with their enrichment coefficients (K) slightly exceeding 1. Notably, Bi showed exceptional enrichment in both the surrounding rock and all pulses, with K values ranging from 10 to 100, occasionally exceeding 100. The trace elements in pyrite displayed characteristics of ore-forming fluid element enrichment.

5.2. Evolution of Ore-Bearing Fluid

Eu2+ is stable in high-temperature reducing hydrothermal fluids, leading to positive Eu anomalies. In contrast, low-temperature reducing conditions result in negative Eu anomalies due to the instability of Eu2+. Under oxidizing conditions, Ce3+ oxidizes to Ce4+, separating from other elements and causing Ce anomalies [26]. The δCe values of pyrite in different pulses varied significantly, with an average around 1 (ranging from 0.87 to 1.28, mean = 1.01), showing no significant Ce anomaly. This finding suggests a reducing environment during pyrite mineralization, consistent with the presence of pyrite and arsenopyrite as the main gold-bearing minerals in the Alin Sb–Au orebodies. Pyrite’s REE profile showed negative Eu anomalies, with δEu values ranging from 0.6 to 1.17 (Figure 8). The δEu average values for pulses 1 to 4 were 1.17, 0.60, 0.96, and 0.84, respectively, indicating considerable variation. For pyrite crystallized during pulse 1 (PyI1), the number of δEu values greater than and less than 1 were roughly equal, suggesting a high formation temperature and a complex formation environment. The pyrite precipitated during pulses 2 and 3 exhibited a general negative Eu anomaly, with only a few points greater than 1, indicating a lower temperature compared with that during pulse 1. The pyrite precipitated during pulse 4 showed a similar number of δEu values greater than and less than 1, suggesting a high temperature. The temperature decreased from pulses 1 to 3, and then increased during pulse 4, consistent with the formation temperatures (240 °C, 220 °C, 235 °C, and 230 °C for pulses 1 to 4, respectively) determined based on the arsenopyrite atomic percentages [5] and quartz fluid inclusions [24], suggesting the possible involvement of external high-temperature hydrothermal fluids during the later stages of mineralization. The host rock of the Alin Sb–Au orebodies showed a narrow range of δCe values (0.92−1.0, mean = 0.96) and δEu (0.68−0.82, mean = 0.73) values, exhibiting an overall negative anomaly. Whether these anomalies can reflect the physicochemical conditions of mineralization requires further investigation.
Co and Ni replace Fe in pyrite mainly through isomorphous substitution. Co, being closer to Fe on the periodic table, is more easily incorporated into the pyrite lattice than Ni. The Co/Ni ratio offers insights into the mineralization conditions and pyrite formation [29,30]. Co is more temperature sensitive than Ni, making the Co/Ni ratio useful for indicating the formation temperature of minerals. A lower Co/Ni ratio typically corresponds to a lower formation temperature [31,32]. During pulse 1 of the Alin Sb–Au orebodies, only one pyrite sample had a Co/Ni ratio below 1, while the others were significantly higher, suggesting a higher crystallization temperature. During pulse 2, the Co/Ni ratio of pyrite ranged from 0.34 to 11.4, with a mean of 3.5. The ratio was roughly balanced between values greater than 1 and those less than 1, indicating a lower precipitation temperature compared with that of pulse 1. During pulse 3, the Co/Ni ratio of pyrite ranged from 0.4 to 2.6, with a mean of 1.33, suggesting a lower precipitation temperature than during pulse 2. During pulse 4, the Co/Ni ratio of pyrite ranged from 1.55 to 5.43, with a mean of 3.56, suggesting a higher precipitation temperature than during pulses 2 and 3. The variation in mineralization temperature in the Aling gold deposit, from high to low and back to high, suggests that higher-temperature hydrothermal fluids may have been mixed in during pulse 4. This hypothesis is supported by large-scale stibnite precipitation during the late stages of mineralization [5].
The Y/Ho, Zr/Hf, and Nb/Ta ratios remained stable within the same hydrothermal system due to the similar ionic radii and valence states of the Y−Ho, Zr−Hf, and Nb−Ta. When disturbed by hydrothermal activity or metasomatism, these element pairs exhibited distinct differentiation. Within the same sample, these ratios showed varying values and wide ranges [32,33,34]. The Y/Ho, Zr/Hf, and Nb/Ta values of the Alin Sb–Au orebodies varied widely across pulses (Figure 8a–c), suggesting that the ore-forming hydrothermal system in each pulse may have been disturbed, possibly due to metasomatism or external hydrothermal fluids.
Figure 8b shows the Y/Ho ratios of pyrite under different pulses of the Alin Sb–Au orebodies. Excluding an outlier of 176.58 from pulse 4, the ratios ranged from 3.24 to 36.38, with an average of 15.60. Pulse 2 exhibited the broadest range of Y/Ho ratios. The average Y/Ho values of the pulses increase, then decreased, and finally increased again. The Y/Ho ratio of chondrites is approximately 28, the mantle ranges from 25 to 30, seawater ranges from 40 to 70, and the Y/Ho ratio in China’s continental crust ranges from 20 to 35 [26,33,35,36]. The Y/Ho ratios of pyrite at all pulses of Alin closely matched those of the continental crust of China, indicating a strong relationship between the ore-forming fluids and the crust. The Zr/Hf and Nb/Ta ratios of pyrite in each pulse of the Aling gold deposit, shown in Figure 8a,c, exhibited large variations. The average Zr/Hf values of the pulses decreased, and then increased, while the Nb/Ta values decreased, then increased, and finally decreased again.

6. Conclusions

  • The ore-forming fluids were low in Cl− and F−, but enriched by sulfur-loving elements such as Cu, Zn, Cd, Pb, and Bi and iron-group elements such as Co.
  • The physical and chemical conditions during mineralization were marked by a reducing environment. Initially, the temperature decreased from pulse 1 to pulse 4, and then increased again.
  • The temperature increase during pulse 4 may have resulted from the influx of Sb-bearing high-temperature hydrothermal fluids, causing a metasomatic effect.

Supplementary Materials

The following supporting information can be downloaded at: https://www.mdpi.com/article/10.3390/min15050491/s1. Table S1: Data of gold-bearing pyrite and host rock rare earth elements and trace elements from the Shuixie Arlin Sb–Au deposit. (All data in this paper are shown in Table S1).

Author Contributions

G.L.: Conceptualization, Methodology, Investigation, Methodology, Writing—original draft, Writing—review and editing, Visualization; S.R.: Conceptualization, Methodology, Data curation, Writing—original draft, Writing—review and editing, Investigation, Formal analysis, Visualization, Supervision; Funding acquisition, Project administration; C.X.: Conceptualization, Methodology, Data curation, Writing—original draft, Writing—review and editing, Investigation, Formal analysis, Visualization, Supervision; Funding acquisition, Project administration; W.W.: Writing—review and editing, Investigation, Formal analysis. All authors have read and agreed to the published version of the manuscript.

Funding

The authors declare that this study received funding from the Scientific Research Foundation of the Yunnan Education Department (grant no. 2023J0131), the Provincial Talent Training Program (grant no. KKSY201556033), the Second Tibetan Plateau Scientific Expedition and Research Project (2021QZKK0301), and the National Natural Science Foundation of China (42163007, 41373049). The authors (Shanshan Ru and Chuandong Xue) were the applicants for funding. The funders (Shanshan Ru and Chuandong Xue) were involved in the study design, collection, analysis, interpretation of data, the writing of this article and the decision to submit it for publication.

Data Availability Statement

Data are contained within the Supplementary Materials.

Conflicts of Interest

The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

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Figure 1. Regional tectonic map (a) (modified from [15]) and simplified geological map (b) (modified from [15,16]) of the Yongping–Weishan Cu–Co–Au–Sb metallogenic belt in western Yunnan Province, SE Xizang–Qinghai Plateau.
Figure 1. Regional tectonic map (a) (modified from [15]) and simplified geological map (b) (modified from [15,16]) of the Yongping–Weishan Cu–Co–Au–Sb metallogenic belt in western Yunnan Province, SE Xizang–Qinghai Plateau.
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Figure 2. Geological map of the Cu–Co–Au–Sb mineralized district in Yongping Changjie–Shuixie (modified from [5,18]). Abbreviation: Fr—Formation.
Figure 2. Geological map of the Cu–Co–Au–Sb mineralized district in Yongping Changjie–Shuixie (modified from [5,18]). Abbreviation: Fr—Formation.
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Figure 3. (a) Representative exploration profile of the Alin deposit [18]. (b) The exploration unit tested the gold (Au) grade of the samples [18].
Figure 3. (a) Representative exploration profile of the Alin deposit [18]. (b) The exploration unit tested the gold (Au) grade of the samples [18].
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Figure 4. Photographs of the mineral assemblage of the Sb–Au orebodies (i,j) (from [5]). (a,b) Orebodies and sampling location corresponding to the sketch map, showing strong barite and stibnite mineralization; (c–e) Euhedral PyI1 and ApyI1 grains in a quartz matrix within the country rock; (f,g) Fine euhedral–subhedral arsenopyrite (ApyI1) as mineral inclusions in subhedral pyrite (PyI1); (h) Level III secondary proliferative ring–core structure, formed subhedral PyI1 overgrown by a subhedral rim of pyrite PyI2, PyI2 overgrown by a subhedral ring of pyrite (PyI3), and arsenopyrite (ApyI2) as mineral inclusions in pyrite (PyI2); (i) Level IV secondary proliferative ring–core structure euhedral PyI1 overgrown by a euhedral–subhedral rim of PyI2 and PyI3, ApyI3 as mineral inclusions in PyI3, and PyI3 overgrown by a euhedral ring of PyI4; (j) Irregular PyI3 wrapping fine ApyI3 cut by long columnar ApyI4, forming a quaternary proliferative ring structure composed of PyI4, ApyI4, and SbnI5; Abbreviations: Apy = arsenopyrite; As − Py = pyrite with high arsenic content; Py = pyrite; Qtz = quartz; Sbn = stibnite; PyI1–PyI4 = pyrite during pulses 1–4, respectively; ApyI1–ApyI4 = arsenopyrite during pulses 1–4, respectively.
Figure 4. Photographs of the mineral assemblage of the Sb–Au orebodies (i,j) (from [5]). (a,b) Orebodies and sampling location corresponding to the sketch map, showing strong barite and stibnite mineralization; (c–e) Euhedral PyI1 and ApyI1 grains in a quartz matrix within the country rock; (f,g) Fine euhedral–subhedral arsenopyrite (ApyI1) as mineral inclusions in subhedral pyrite (PyI1); (h) Level III secondary proliferative ring–core structure, formed subhedral PyI1 overgrown by a subhedral rim of pyrite PyI2, PyI2 overgrown by a subhedral ring of pyrite (PyI3), and arsenopyrite (ApyI2) as mineral inclusions in pyrite (PyI2); (i) Level IV secondary proliferative ring–core structure euhedral PyI1 overgrown by a euhedral–subhedral rim of PyI2 and PyI3, ApyI3 as mineral inclusions in PyI3, and PyI3 overgrown by a euhedral ring of PyI4; (j) Irregular PyI3 wrapping fine ApyI3 cut by long columnar ApyI4, forming a quaternary proliferative ring structure composed of PyI4, ApyI4, and SbnI5; Abbreviations: Apy = arsenopyrite; As − Py = pyrite with high arsenic content; Py = pyrite; Qtz = quartz; Sbn = stibnite; PyI1–PyI4 = pyrite during pulses 1–4, respectively; ApyI1–ApyI4 = arsenopyrite during pulses 1–4, respectively.
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Figure 5. CI chondrite-normalized REE patterns of pyrite (normalization values after [22]).
Figure 5. CI chondrite-normalized REE patterns of pyrite (normalization values after [22]).
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Figure 6. Continental crust-normalized trace element patterns of pyrite (normalization values based on [23]).
Figure 6. Continental crust-normalized trace element patterns of pyrite (normalization values based on [23]).
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Figure 7. Average continental crust-normalized trace element patterns of pyrite (normalization values based on [23]).
Figure 7. Average continental crust-normalized trace element patterns of pyrite (normalization values based on [23]).
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Figure 8. Ore-forming fluid characteristics: (a) Nb/Ta ratios of pyrite in each pulse; (b) Y/Ho ratios of pyrite in each pulse; (c) Zr/Hf ratios of pyrite in each pulse; and (d) δCe–δEu values of pyrite in each pulse.
Figure 8. Ore-forming fluid characteristics: (a) Nb/Ta ratios of pyrite in each pulse; (b) Y/Ho ratios of pyrite in each pulse; (c) Zr/Hf ratios of pyrite in each pulse; and (d) δCe–δEu values of pyrite in each pulse.
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Li, G.; Ru, S.; Xue, C.; Wang, W. Evolution of the Ore-Bearing Fluid of Alin Sb–Au Orebodies in Shuixie Cu–Co Orefield, SW China: Constraints on the Rare Earth Element and Trace Element Components of Auriferous Pyrite and Host Rock. Minerals 2025, 15, 491. https://doi.org/10.3390/min15050491

AMA Style

Li G, Ru S, Xue C, Wang W. Evolution of the Ore-Bearing Fluid of Alin Sb–Au Orebodies in Shuixie Cu–Co Orefield, SW China: Constraints on the Rare Earth Element and Trace Element Components of Auriferous Pyrite and Host Rock. Minerals. 2025; 15(5):491. https://doi.org/10.3390/min15050491

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Li, Guo, Shanshan Ru, Chuandong Xue, and Wei Wang. 2025. "Evolution of the Ore-Bearing Fluid of Alin Sb–Au Orebodies in Shuixie Cu–Co Orefield, SW China: Constraints on the Rare Earth Element and Trace Element Components of Auriferous Pyrite and Host Rock" Minerals 15, no. 5: 491. https://doi.org/10.3390/min15050491

APA Style

Li, G., Ru, S., Xue, C., & Wang, W. (2025). Evolution of the Ore-Bearing Fluid of Alin Sb–Au Orebodies in Shuixie Cu–Co Orefield, SW China: Constraints on the Rare Earth Element and Trace Element Components of Auriferous Pyrite and Host Rock. Minerals, 15(5), 491. https://doi.org/10.3390/min15050491

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