Mineralogical and Geochemical Characteristics of Carbonates and Their Geological Signiﬁcance to the Fuli Pb-Zn Deposit, Yunnan Province

: Mississippi Valley-Type (MVT) deposits are among the main types of Pb-Zn deposits that feature carbonate minerals as the main gangue minerals; their formation runs through the entire metallogenic process of MVT deposits. Therefore, carbonate minerals contain rich information on metallogenic fluid evolution and are thus important prospecting indicators. The Fuli Pb-Zn deposit in eastern Yunnan is located in the southeast of the Sichuan-Yunnan-Guizhou downward trend. Fe and Mn entered the alkaline environment of the carbonate minerals, while Fe and Mn were released into the acidic fluid, indicating that due to the metasomatism from strong to weak, their metallogenic environment evolved from alkaline to acidic. From the early to the late stage of mineralization, Sr showed an upward trend, which might indicate that the continuous reaction between the hydrothermal fluid and the wall rock continuously released Sr into the fluid. The Fe-Sr and Mn-Sr diagrams show that two kinds of fluid mixing occurred in the ore-forming process. The Fuli Pb-Zn deposit may have formed from mineral precipitation caused by the mixing of the metal-rich, oxidized acidic fluid and the sulfur-rich, reduced alkaline fluid. The results show that the Fuli Pb-Zn deposit belongs to MVT deposits.


Introduction
The Sichuan-Yunnan-Guizhou (SYG) metallogenic province is located on the southwestern margin of the Yangtze craton in South China. It hosts world-class concentrations of lead (Pb), zinc (Zn), silver (Ag), and germanium (Ge) deposits, which are mostly associated with low-temperature fluids (<300 • C) [1][2][3][4][5][6]. The Pb-Zn deposits in this area mainly occur in a range of carbonate rocks from the Neoproterozoic Sinian Dengying Formation to the Late Paleozoic Permian Yangxin Formation. These deposits are characterized by high Pb and Zn grades and are enriched in Cd, Ge, Ga, etc. Mineralization, which is generally structurally controlled by faults and folds, is interpreted as epigenetic [7] and controlled by the fault-fold structural system. Previous studies generally agree on its nature as an epigenetic deposit [8][9][10][11][12][13][14][15][16].
The Fuli Pb-Zn deposit is located in the southeast of the Sichuan-Yunnan-Guizhou Pb-Zn ore concentration area in Fuyuan County, Yunnan Province ( Figure 1). The deposit occurs in the interlayer fractured zone of Middle Permian Yangxin Formation Dolomite, and the orebody shape is stratiform. Studies and reports on this deposit are rare, particularly those focused on carbonate minerals. Carbonate minerals are gangue minerals in many deposits and are especially important in Mississippi Valley-Type (MVT) deposits [17]. Carbonate minerals run through the entire metallogenic process of MVT deposits. Therefore, carbonate minerals carry rich metallogenic information that is important for understanding the genesis of MVT deposits. Compared to traditional methods for trace element analysis, LA-ICP MS (laser ablation-inductively coupled plasma mass spectrometry) is an in situ analysis method that can overcome the problem of sample contamination caused by sample selection and effectively judge the spatial distributions of trace elements in the sample [18,19].
In this article, with the help of LA multi-collector ICP MS (LA-MC-ICP MS), we performed a detailed mineralogical and in situ trace element composition study on gangue minerals (dolomite and calcite) from the Fuli Pb-Zn deposit. The composition and distribution characteristics of trace elements in the gangue minerals of the ore deposit and their geological significance were investigated to determine the source and properties of the ore-forming fluid forming this deposit.  [3,6,20]).
In this article, with the help of LA multi-collector ICP MS (LA-MC-ICP MS), we performed a detailed mineralogical and in situ trace element composition study on gangue minerals (dolomite and calcite) from the Fuli Pb-Zn deposit. The composition and distribution characteristics of trace elements in the gangue minerals of the ore deposit and their geological significance were investigated to determine the source and properties of the ore-forming fluid forming this deposit.

Tectonic Features
The study area is located between the Qujing and Mile-Shizong faults ( Figure 1). The exposed basement of the study area is the Kunyang Group of the Meso-Neo Proterozoic era, while the Sinian, Cambrian, Ordovician, and Silurian are missing above the Kunyang Group. The Haikou Formation of the Middle Devonian is in direct contact with the Kunyang Group in an angular unconformity, and the Kunyang Group is only exposed in  [3,6,20]).

Tectonic Features
The study area is located between the Qujing and Mile-Shizong faults ( Figure 1). The exposed basement of the study area is the Kunyang Group of the Meso-Neo Proterozoic era, while the Sinian, Cambrian, Ordovician, and Silurian are missing above the Kunyang Group. The Haikou Formation of the Middle Devonian is in direct contact with the Kunyang Group in an angular unconformity, and the Kunyang Group is only exposed in the southwest of the study area. The Proterozoic structure framework is dominated by the east-west trend. After the Jinning, Caledonian, and Hercynian movements, the structure in the area was transformed into a north-south-trending structure supplemented by a north-east-trending structure, both north-west and east-west structures inherited. During the Yanshan and Himalayan tectonic movements, the northwest-trending structure was dominantly active and gradually formed the current north-south-trending structure followed by the north-east- trending structure. The east-west and northwest-trending structures were also developed at this time. In general, carbonate and clastic rocks of the Proterozoic and Paleozoic eras were deposited in the platform fold belt with a total thickness of more than 10,000 m. They show the characteristics of long-term depression deposition in this area. The strata in the Sinian to Permian eras are relatively well developed but exhibit frequent discontinuities. The late Hercynian period witnessed the large-scale eruption of basalt magma and its intrusion; thus, it is the only magmatic rock in the area. The structural framework is characterized by parallel and almost equidistant north-south-trending faults. The northeast-trending structures are also relatively clear with large scale and a long developmental history. The north-south-and northwest-trending structures play important roles in controlling the formation, distribution, and enrichment of Zn-Pb deposits in the region.
The Mile-Shizong fault is the main fault in the region. The Tuoniu-Duza anticline on the west side of the mining area is the most important fold structure in the region ( Figure 1b). Together, both the fault and the anticline control the distribution of regional strata along with the distribution of secondary structures and mineralization in the area. The Mile-Shizong fault is generally northeast trending (north-northeast trending near the mining area). The fault zone is composed of multiple high-angle and steep faults, and the faults are generally compressional reverse faults. The Tuoniu-Duza anticline is a gentle structure with a stratum dip angle of 10-12 • on both flanks and a north-northeast-trending axis, consistent with the Mile-Shizong fault.
This area experienced a near-SN-trending extension in the Late Permian. During the Late Triassic, it was affected by tectonic events such as the collisional orogeny of the Indosinian and South China Plates and the Yidun arc collisional orogen, forming a series of near SN-NNE compressive torsional fractures (Figure 2a). Influenced by the Late Triassic compressional structure and the Tuoniu-Duza anticline, a series of extensional fissures and interlayer fracture zones were generated in the middle Yangxin Formation (P 2 y 2 ), which provided structural traps for circulating metal-rich fluids. The NE-NNE-trending structures developed in the mining area ( Figure 2a) are post-metallogenic structures with dip angles of 50-80 • . These structures have deformed and modified the continuity of the mineral body.

Deposit Geology
The Fuli Pb-Zn deposit is located in the southeastern section of the SYG metallogenic province in the southwestern margin of the Yangtze craton. The Early-Middle Permian era, which was the largest transgressive cycle in this area during the Late Paleozoic, mainly led to the formation of carbonate deposits. Under the influence of continental uplift, a set of neritic Mg-bearing carbonate assemblages was developed. The strata of the Middle Permian Yangxin Formation (P 2 y) are exposed in the mining area (Figure 2b), and the lithology of the lower section (P 2 y 1 ) is light-grey limestone with dolomite. The middle section (P 2 y 2 ) is mainly interbedded with light-grey limestone and dolomite with local siliceous dolomite; it is the main type of stratum occurring in the Pb-Zn polymetallic deposits. The upper section (P 2 y 3 ) is grey medium-thick layered crystalline limestone with a small amount of dolomitic limestone and dolomite containing flint bands. The Pb-Zn ore in the mining area is hosted in dolomitic limestone, and the shape of the orebodies is stratiform. At present, two Pb-Zn mineralization bodies have been found in the Fuli Pb-Zn mine.
Mineral body #1: This mineral body is 210 m in length with an inclination extension of 314 m, true thickness of 1.11 m, overall inclination of 181 • , and dip angle of 15 • . The body is distributed in the southern part of the mining area. The average grades of Pb and Zn are 1.51% and 8.43%, respectively. The mineral body is thick in the middle and thin at both ends. In addition, the surface is thin and low grade, while the body becomes thicker and the metal grade increases with greater depth.

Deposit Geology
The Fuli Pb-Zn deposit is located in the southeastern section of the SYG metallogenic province in the southwestern margin of the Yangtze craton. The Early-Middle Permian era, which was the largest transgressive cycle in this area during the Late Paleozoic, mainly led to the formation of carbonate deposits. Under the influence of continental uplift, a set of neritic Mg-bearing carbonate assemblages was developed. The strata of the Middle Permian Yangxin Formation (P2y) are exposed in the mining area (Figure 2b), and the lithology of the lower section (P2y 1 ) is light-grey limestone with dolomite. The middle section (P2y 2 ) is mainly interbedded with light-grey limestone and dolomite with local siliceous dolomite; it is the main type of stratum occurring in the Pb-Zn polymetallic deposits. The upper section (P2y 3 ) is grey medium-thick layered crystalline limestone with a small amount of dolomitic limestone and dolomite containing flint bands. The Pb-Zn ore in the mining area is hosted in dolomitic limestone, and the shape of the orebodies is stratiform. At present, two Pb-Zn mineralization bodies have been found in the Fuli Pb-Zn mine.
Mineral body #1: This mineral body is 210 m in length with an inclination extension of 314 m, true thickness of 1.11 m, overall inclination of 181°, and dip angle of 15°. The body is distributed in the southern part of the mining area. The average grades of Pb and Mineral body #II: The mineral body is distributed in the northern part of the mining area with an unexposed surface. The mineral body features a layered output with a 220-mlong engineering control trend, an inclination extension of 103-215 m, an overall inclination of 144 • , and a dip angle of 12 • . The thickness of the mineral body is 1.00-1.20 m, with an average of 1.11 m. The average grades of Pb and Zn are 5.20% and 13.29%, respectively. Characteristically, the mineral body is thick in the middle, where the mineral grade is high, and thin at both ends, where the mineral grade is low. In the hanging wall of the normal fault, the mineral body shows thickening and enrichment.
Through field surveys, observations of ore texture and mineral association, and previous petrographic work, the formation process of the Fuli Pb-Zn deposit was divided into three stages: the early metallogenic period, main metallogenic period, and late metallogenic period (Figures 3 and 4).    Early metallogenic period: The main feature of the ore, which was formed in the early metallogenic period, is hydrothermal alteration dolomitization. The dolomite formed during this period can be divided into two stages: the D1 and D2 stages. The D1 dolomite appears as grey-black, fine-grained xenomorphic granules. The sphalerite formed at this stage is mainly reddish-brown to brown-black, and the crystals are mainly subhedral to euhedral in shape. The pre-altered dolomite is wrapped in grey-white and grey-black fine-grained mesocrystalline dolomite that coexists with the sphalerite. It is disseminated and speckled. Agglomerated sphalerite grains are distributed in the dolomite. Galena is present in a cubic crystal form. During the main mineralization period, two phases of dolomite can be identified: grey dolomite (HD3) and white dolomite (HD4). These two kinds of dolomite are hydrothermal bright crystalline dolomite types that are closely related to the ore and form a dolomite halo around the mineral body (Figs. 3a and b). The HD3 dolomite is a grey-white, meso-coarse-grained semi-autotype dolomite with a grain size between 0.1 and 0.8 mm; it is distributed along the edges of sphalerite and galena (Figure 3c,e). The iron content of this dolomite is high, and it was formed at the same time as the sulfide. The HD4 dolomite usually presents saddle-like, euhedral coarse crystals that are whitemilky white in color. The HD4 dolomite fills open spaces or cement fissures. The HD3 dolomite is in direct contact with metasomatic wall rock, while the HD4 dolomite is mostly in contact with the grey dolomite and was formed later than the grey dolomite ( Figure 3f). The HD3 dolomite presents bright red luminescence under CL, while white dolomite shows dark red luminescence. Early metallogenic period: The main feature of the ore, which was formed in the early metallogenic period, is hydrothermal alteration dolomitization. The dolomite formed during this period can be divided into two stages: the D1 and D2 stages. The D1 dolomite appears as grey-black, fine-grained xenomorphic granules. (Figure 3f The sphalerite formed at this stage is mainly reddishbrown to brown-black, and the crystals are mainly subhedral to euhedral in shape. The pre-altered dolomite is wrapped in grey-white and grey-black fine-grained mesocrystalline dolomite that coexists with the sphalerite. It is disseminated and speckled. Agglomerated sphalerite grains are distributed in the dolomite. Galena is present in a cubic crystal form. During the main mineralization period, two phases of dolomite can be identified: grey dolomite (HD3) and white dolomite (HD4). These two kinds of dolomite are hydrothermal bright crystalline dolomite types that are closely related to the ore and form a dolomite halo around the mineral body (Figure 3a,b). The HD3 dolomite is a grey-white, meso-coarsegrained semi-autotype dolomite with a grain size between 0.1 and 0.8 mm; it is distributed along the edges of sphalerite and galena (Figure 3c,e). The iron content of this dolomite is high, and it was formed at the same time as the sulfide. The HD4 dolomite usually presents saddle-like, euhedral coarse crystals that are white-milky white in color. The HD4 dolomite fills open spaces or cement fissures. The HD3 dolomite is in direct contact with metasomatic wall rock, while the HD4 dolomite is mostly in contact with the grey dolomite and was formed later than the grey dolomite ( Figure 3f). The HD3 dolomite presents bright red luminescence under CL, while white dolomite shows dark red luminescence. Late metallogenic period: The euhedral calcite veins (C5), which are the last stage of gangue minerals, are massively thick veins that enclose the sphalerite and galena formed in the main metallogenic period. This kind of calcite is non-luminescent or presents dark crimson luminescence under CL.

Elemental Measurement by LA-ICP MS
In this study, trace element analysis of minerals was conducted by LA-ICP MS at the Wuhan Sample Solution Analytical Technology Co., Ltd., Wuhan, China. The operating conditions for the LA system and the ICP MS instrument and data reduction were the same as those described by Zong et al. [21]. Laser sampling was performed using a Geolas Pro LA system (Coherent Inc., Santa Clara, CA, USA) 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 7900 (Agilent Technologies, Inc., Santa Clara, CA, USA) ICP- Late metallogenic period: The euhedral calcite veins (C5), which are the last stage of gangue minerals, are massively thick veins that enclose the sphalerite and galena formed in the main metallogenic period. This kind of calcite is non-luminescent or presents dark crimson luminescence under CL.

Elemental Measurement by LA-ICP MS
In this study, trace element analysis of minerals was conducted by LA-ICP MS at the Wuhan Sample Solution Analytical Technology Co., Ltd., Wuhan, China. The operating conditions for the LA system and the ICP MS instrument and data reduction were the same as those described by Zong et al. [21]. Laser sampling was performed using a Geolas Pro LA system (Coherent Inc., Santa Clara, CA, USA) 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 7900 (Agilent Technologies, Inc., Santa Clara, CA, USA) 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 instrument. A "wire" signal smoothing device was included in the LA system [22]. The spot size and frequency of the laser were set to 44 µm and 5 Hz, respectively. The trace element compositions of minerals were calibrated against various reference materials (NIST 610, BHVO-2G, BCR-2G, BIR-1G, and MACS-3) without using an internal standard [23,24]. Each analysis incorporated a background acquisition of approximately 20-30 s followed by 50 s of data acquisition from the sample. ICP-MS-DataCal v12.2, an Excel-based software, was used to perform off-line selection and integration of the background; the signals, time-drift correction, and quantitative calibration for trace element analysis [23].

C-O In Situ Isotopes
Sixteen representative samples were selected, and microdrill sampling was carried out in the proposed area (wall rock, metallogenic area, and non-metallogenic area) for C-O isotopic analysis. The dolomite C-O isotope test and analysis were performed in the Stable Isotope Analysis Laboratory of Kunming University of Science and Technology, China. The test adopted the 100% phosphoric acid method. The sample reacted with phosphoric acid for more than 1 h at 90 • C. Under the protection of high-purity helium, the released CO 2 gas was introduced into an ISO Prime 100 gas-phase stable isotope mass spectrometer (Elementar Analysensysteme GmbH, Langenselbold, Hessian, Germany) for the determination of C and O isotope composition. The experiment used Vienna Pee Dee Belemnite (V-PDB) as the standard, where δ 18 O SMOW = 1.03086 × δ 18 O PDB + 30.86.

CL Analysis
CL petrography was performed on a Leica DM2700P binocular petrographic microscope (Leica Microsystems, Wetzlar, Hessian, Germany) with a CLF-2 CL system (BII Inc., Markham, ON, Canada) at the Key Laboratory of Mineralogy and Metallogeny, GIGCAS. The voltage for the systems was held between 13 and 15 kV, the current ranged from 240 to 260 µA, and chamber pressure was between 50 and 60 millitorrs.

CL Images of Dolomite
The CL images for representative dolomite samples are illustrated in Figure 6. The different stages of dolomite exhibited different characteristics under CL. In the early metallogenic period, the D1 dolomite was non-luminescent or presented a dark crimson luminescence, while the D2 dolomite presented a dark red luminescence under CL. In the main metallogenic period, the grey dolomite (HD3) presented bright red luminescence, and the white dolomite (HD4) was dark red. In the late metallogenic period, the euhedral calcite veins (C5) were non-luminescent or presented a dark crimson luminescence.

Textural and Mineralogical Analyses
The LA-ICP-MS in situ trace element analysis results of carbonate minerals for the Fuli deposits are shown in Table 1, and the variations in these trace elements are shown in Figure 7.

C-O Isotopic Characteristics
The analytical results of the C-O isotopic compositions are listed in Table 2. In the early metallogenic period, the δ 18 O SMOW value for dolomite ranged from 16.91‰ to 20.55‰, with an average value of 17.87‰. The δ 13 C PDB value ranged from 2.59 to 3.5‰, with an average value of 2.91‰.
In the main metallogenic period, the δ 18 O SMOW value for the dolomite ranged from 13.29‰ to 16.76‰, with an average value of 15.77‰. The δ 13 C PDB value ranged from −4.13‰ to 3.02‰, with an average value of 0.76‰.
In the late metallogenic period, the δ 18 O SMOW value for calcite was 22.48‰, and the δ 13 C PDB value was −0.85‰.

Origin of Hydrothermal Dolomite
Y and Ho have been reported to show similar geochemical behaviors in fluid migration [25][26][27][28]. Therefore, the Y/Ho ratios can effectively reflect the source of fluid [29,30]. The Y/Ho ratios of different fluids (seawater, brine, etc.) are significantly different. The Y/Ho ratios of deep brines (and unaffected by seawater) tend to be relatively low because these brines might interact with basement strata formations or sediments; the Y/Ho ratios are close to those of the upper crust (27.5) or chondrites (25)(26)(27)(28) [31][32][33]. In this study, the Y/Ho ratios of hydrothermal dolomite found in the different stages of the Fuli Pb-Zn deposits ranged between 33.21 and 72.56 (mean value: 48.72). These Y/Ho ratios are higher than those of typical brines, and some of the ratios are lower than those of marine carbonates. This implies that the fluid that formed the hydrothermal dolomite was mixed with the carbonate wall rock, which is consistent with the overall horizontal distribution of dolomite (rock) on the Y/Ho-La/Ho map (Figure 8). Meanwhile, the distribution of stone (rock) in the Y/Ho-La/Ho diagram shows obvious stratification, implying that the formation of dolomite was affected by the mixing of various fluids.
posits ranged between 33.21 and 72.56 (mean value: 48.72). These Y/Ho ratios are higher than those of typical brines, and some of the ratios are lower than those of marine carbonates. This implies that the fluid that formed the hydrothermal dolomite was mixed with the carbonate wall rock, which is consistent with the overall horizontal distribution of dolomite (rock) on the Y/Ho-La/Ho map (Figure 8). Meanwhile, the distribution of stone (rock) in the Y/Ho-La/Ho diagram shows obvious stratification, implying that the formation of dolomite was affected by the mixing of various fluids.

Sources of C and O
The C and O isotopic compositions of hydrothermal dolomite can effectively trace the source of mineralizing materials [6,34,35]. There are three main sources of C and O in a hydrothermal system: (1) the mantle (δ 13 C = −8‰ to −4‰, δ 18 O = +6‰ to 10‰) [36][37][38]; (2) marine carbonate rocks (δ 13 C = −4‰ to +4‰, δ 18 O = +20‰ to +30‰) [39,40]; and (3) sedimentary organic matter (δ 13 C = −30‰ to + 10‰, δ 18 O = +24‰ to +30‰) [41,42]. As shown in Figure 9, the hydrothermal dolomite of the Fuli deposit is significantly different from the sedimentary organic matter. The δ 13 C PDB and δ 18 O SMOW values for the wall rocks of the Fuli deposit are slightly smaller than those of the marine carbonate rocks. The C and O isotopic composition range of hydrothermal dolomite is between the values for the primary and marine carbonate rocks; this indicates that the C in the ore-forming fluids of the Fuli deposit might originate from the sedimentary contamination of the mantle multiphase system, effects of high temperature, or the dissolution of marine carbonates. According to the microscopic thermometry of fluid inclusions in the Fuli deposit (Liang, unpublished data), the ore-forming fluids in this area comprised medium-and low-temperature fluids, suggesting that high-temperature differentiation could not be the main factor. In Figure 9, most of the projection points of δ 13 C PDB and δ 18 O SMOW are distributed nearly horizontally, while only a few suggest the depletion of δ 13 C PDB and δ 18 O SMOW at the same time. These patterns indicate the influence of dissolution of marine carbonates, while the lower δ 18 O could be due to the O isotope exchange between the wall rock and the depleted δ 18 O ore-forming fluid.

Properties of Ore-Forming Fluids
During geological processes, rare earth elements (REEs) have similar geochemical behavior and often migrate as a whole. In addition, as the radii of rare earth ions are similar to those of Ca 2+ ions, REEs are often concentrated into Ca-bearing mineral lattices via isomorphic substitutions, resulting in negligible rare earth contents in sulfides [43][44][45][46]. Therefore, the rare earth compositions of Ca-bearing minerals may reflect the REE-induced characteristics of ore-forming fluids; therefore, the rare earth distribution model has become a reliable tool to explore the source and evolution of ore-forming fluids [47][48][49][50][51][52][53][54].
published data), the ore-forming fluids in this area comprised medium-and low-temperature fluids, suggesting that high-temperature differentiation could not be the main factor. In Figure 9, most of the projection points of δ 13 CPDB and δ 18 OSMOW are distributed nearly horizontally, while only a few suggest the depletion of δ 13 CPDB and δ 18 OSMOW at the same time. These patterns indicate the influence of dissolution of marine carbonates, while the lower δ 18 O could be due to the O isotope exchange between the wall rock and the depleted δ 18 O ore-forming fluid.

Properties of Ore-Forming Fluids
During geological processes, rare earth elements (REEs) have similar geochemical behavior and often migrate as a whole. In addition, as the radii of rare earth ions are similar to those of Ca 2+ ions, REEs are often concentrated into Ca-bearing mineral lattices via isomorphic substitutions, resulting in negligible rare earth contents in sulfides [43][44][45][46]. Therefore, the rare earth compositions of Ca-bearing minerals may reflect the REE-induced characteristics of ore-forming fluids; therefore, the rare earth distribution model has become a reliable tool to explore the source and evolution of ore-forming fluids [47][48][49][50][51][52][53][54].
Eu and Ce anomalies can reflect the redox conditions of ore-forming fluids. Eu has two ionic states in nature: Eu 3+ and Eu 2+ . Eu 3+ has similar properties to other REE trivalent ions. Eu 2+ is relatively active, and it often enters the Ca 2+ ion lattices of minerals, leading to positive Eu anomalies [55]. The Eu anomaly is determined by the temperature, Eu and Ce anomalies can reflect the redox conditions of ore-forming fluids. Eu has two ionic states in nature: Eu 3+ and Eu 2+ . Eu 3+ has similar properties to other REE trivalent ions. Eu 2+ is relatively active, and it often enters the Ca 2+ ion lattices of minerals, leading to positive Eu anomalies [55]. The Eu anomaly is determined by the temperature, pressure, the pH of the ore-forming fluid, and the state of occurrence of the REEs; among these factors, the Eu anomaly is most sensitive to temperature [56]. As demonstrated by Moller et al. [57], at higher temperatures (>200 • C), Eu 3+ is reduced to Eu 2+ , which does not readily enter the crystalline structure of minerals; hence, a negative Eu anomaly in carbonate forms in equilibrium with the fluids if the temperature is above 200 • C. Furthermore, preservation of positive Eu anomalies requires that the fluid cools sufficiently to stabilize Eu 3+ . The ionic radius of Eu 3+ is 0.99, similar to that of Ca 2+ ; hence, it is easy for Eu 3+ to displace Ca 2+ in the mineral lattice, resulting in a positive Eu anomaly [58,59]. The Eu anomalies of D1 and D2 dolomites in the Fuli deposit are insignificant, indicating that the fluids were primarily low in temperature with high oxygen fugacity. The HD3 grey hydrothermal dolomite exhibited a weak positive Eu anomaly (Figure 10), which might be indicative of a gradual decrease in the temperature of the ore-forming fluid with the progress of mineralization. Furthermore, the fluid gradually became enriched in Eu 2+ . Sulfide precipitation and the formation of hydrothermal dolomite were accompanied by the reduction of Eu 3+ to Eu 2+ , indicating that the precipitation of sulfide and dolomite may have resulted from the decreases in oxygen fugacity and temperature of the fluid; the positive Eu anomaly reflects the relatively reducing environment. Ce anomalies are determined by the pH and ƒO 2 of the ore-forming fluids and are more sensitive to pH [60]. Ce 3+ is oxidized to Ce 4+ , which has low solubility and is easily lost from the ore-forming fluid via hydroxide adsorption, resulting in a negative Ce anomaly [61]. The negative Ce anomalies of the D1 and D2 dolomites of the Fuli deposit were significant, and Ce 4+ adsorbed hydroxide in the ore-forming fluids, indicating that the fluid was alkaline when the wall rocks were altered by hydrothermal fluids. The HD3 grey hydrothermal dolomite showed a strong negative Ce anomaly, while the HD4 white hydrothermal dolomite and C5 late calcite showed weakly negative Ce anomalies ( Figure 10). These patterns indicate that the basicity of the ore-forming environment weakened followed by an evolution of the ore-forming environment to neutrality and finally weak acidity.
reduction of Eu to Eu , indicating that the precipitation of sulfide and dolomite may have resulted from the decreases in oxygen fugacity and temperature of the fluid; the positive Eu anomaly reflects the relatively reducing environment. Ce anomalies are determined by the pH and ƒO2 of the ore-forming fluids and are more sensitive to pH [60]. Ce 3+ is oxidized to Ce 4+ , which has low solubility and is easily lost from the ore-forming fluid via hydroxide adsorption, resulting in a negative Ce anomaly [61]. The negative Ce anomalies of the D1 and D2 dolomites of the Fuli deposit were significant, and Ce 4+ adsorbed hydroxide in the ore-forming fluids, indicating that the fluid was alkaline when the wall rocks were altered by hydrothermal fluids. The HD3 grey hydrothermal dolomite showed a strong negative Ce anomaly, while the HD4 white hydrothermal dolomite and C5 late calcite showed weakly negative Ce anomalies ( Figure 10). These patterns indicate that the basicity of the ore-forming environment weakened followed by an evolution of the oreforming environment to neutrality and finally weak acidity. As shown in Figure 11, the D1 and D2 dolomites of the Fuli deposit are enriched in Fe and Mn. Moving from the HD3 grey hydrothermal dolomite to the HD4 white hydrothermal dolomite, Fe and Mn show a downward trend. Fe and Mn entered the carbonate in an alkaline environment, whereas Fe and Mn were released into the fluid in an acidic environment. Thus, as the metasomatism changed from strong to weak, the ore-forming environment predominantly changed from alkaline to acidic. The Sr content increased from D2 dolomite to HD3 grey hydrothermal dolomite, HD4 white hydrothermal dolomite, and late calcite, which might indicate that the continuous reaction between the hydrothermal fluid and the wall rock releases Sr in the wall rock into the fluid. The changes in the Fe, Mn, and Sr contents suggest that mineralization might result from the mixing of the two fluids. environment . Thus, as the metasomatism changed from strong to weak, the ore-forming environment predominantly changed from alkaline to acidic. The Sr content increased from D2 dolomite to HD3 grey hydrothermal dolomite, HD4 white hydrothermal dolomite, and late calcite, which might indicate that the continuous reaction between the hydrothermal fluid and the wall rock releases Sr in the wall rock into the fluid. The changes in the Fe, Mn, and Sr contents suggest that mineralization might result from the mixing of the two fluids. Figure 11. Plots of LA-ICP-MS data from carbonate depicting relative changes in Mn, Fe, and Sr in the Fuli Pb-Zn deposit. Figure 11. Plots of LA-ICP-MS data from carbonate depicting relative changes in Mn, Fe, and Sr in the Fuli Pb-Zn deposit.

Ore Genesis Model
There are three mechanisms that may lead to sulfide precipitation: the reduced sulfur mechanism, the sulfate reduction mechanism, and the mixing mechanism. In the reduced sulfur mechanism, that metallogenic metals and reduced sulfur coexist in the same fluid, and sulfides precipitate out due to pH changes, cooling, or dilution of the fluid [58,[63][64][65]. The most critical problem in this model is that the solubility of sulfide in the fluid is considered to be very low; therefore, it is difficult for the fluid to allow the coexistence of a large amount of metal and reduced sulfur. The sulfate reduction mechanism requires that the ore-forming metals are transported to the ore-forming area in the form of sulfate; reduced sulfur is then formed via bacterial reduction or the thermochemical reduction of sulfate. Finally, the reduced sulfur and ore-forming metals react and precipitate out [66][67][68]. In the mixing mechanism, a fluid transports ore-forming metals, and another fluid transports the reduced sulfur. The mixing of the two fluids results in the precipitation of metal sulfides [37,38,[69][70][71][72][73][74].
The Permian Yangxin Formation is a stratum of limestone mixed with dolomite. The water in this formation is stored between the stratum and remains in balance with calcium carbonate. Limestone mixed with the dolomite stratum contains a large amount of organic matter and pyrite formed during the depositional period; hence, the formation water occurs in an alkaline environment with low oxygen fugacity and low temperature, and it can produce reduced S from the marine sulfates hosted in the regional rocks by the TSR reaction to form a fluid rich in reduced sulfur-bearing. Based on the reduction in oxygen fugacity and temperature of the environment during mineralization along with the evolution from alkaline to neutral to weakly acidic, one can infer that mineralization might be the result of the mixing of the fluid containing metal elements and water in the sedimentary formation. Studies on fluid inclusions in the Fuli deposit (unpublished data) have shown homogeneous temperatures distributions in the fluid inclusions in the sphalerite of the Fuli Pb-Zn deposit, with temperatures between 110 • C and 210 • C. The salinity distribution has two intervals: one between 2% and 11% NaCleqv and a second between 19% and 23% NaCleqv. The ore-forming fluid may be a mixture of medium-low-temperature, and low-salinity fluids with medium-low temperature, and high-salinity fluids.
Integrating all available information, the mineralization model of the Fuli deposit can be described as follows ( Figure 12): Thermal activity related to the Indosinian Orogeny resulted in the circulation of hydrothermal fluids within basement rocks (Kunyang Group clastic rocks). These fluids carried up abundant metals and were likely acidic with high oxygen fugacity and high temperature. In the early mineralization stage, compressional tectonics drove the hydrothermal fluids upward along regional faults to the surface, and the wall rocks were subjected to hydrothermal alteration to form D1 and D2 dolomites. During syn-mineralization, the hydrothermal fluids (carrying abundant metals) migrated upwards and mixed with H 2 S-rich fluid from the regional strata. This resulted in the precipitation of metallic sulfides and hydrothermal crystalline dolomites in the middle Permian Yangxin Formation, where the faults and contact zones are well developed. In the post-mineralization period, the circulating process between dissolution and re-precipitation in carbonate rocks can cause calcite precipitation [13,75]. Moving from the early stage to the late stage, the oxygen fugacity of the fluid evolved from high to low, and the metallogenic environment changed from alkaline to neutral/weakly acidic. Mineralization occurred due to the changes in the physical and chemical conditions of the ore-forming fluid. Therefore, we suggest that the Fuli Pb-Zn deposit might have formed due to the mixing of metal-rich, oxidized acidic fluid with sulfur-rich, reduced alkaline fluid.

Conclusions
(1) The dolomite from the Fuli deposit exhibits both dark and bright luminescence under CL, with primarily bright luminescence. Moving from early dolomite to late calcite, the luminescence shifts from dark to bright and back to dark. (2) The C-O isotopic data suggest that the source of C in the Fuli Pb-Zn deposit might be Therefore, we suggest that the Fuli Pb-Zn deposit might have formed due to the mixing of metal-rich, oxidized acidic fluid with sulfur-rich, reduced alkaline fluid.

Conclusions
(1) The dolomite from the Fuli deposit exhibits both dark and bright luminescence under CL, with primarily bright luminescence. Moving from early dolomite to late calcite, the luminescence shifts from dark to bright and back to dark. (2) The C-O isotopic data suggest that the source of C in the Fuli Pb-Zn deposit might be the dissolution of marine carbonates. In contrast, the O isotopic signature is related to the sedimentary rock contamination of the mantle multiphase system, while the relatively low δ 18 O might be attributed to O isotope exchange between the wall rock and the depleted δ 18 O ore-forming fluid.