Rb–Sr Pyrite Dating and S–Pb Isotopes in the Fang’an Gold Deposit, Wuhe Area, Eastern Anhui Province

Abstract

The Fang’an gold deposit in the Wuhe area, Anhui Province, is located in the area adjacent to the Bengbu Uplift and Wuhe Platform Depression in the southeastern part of North China. This study aimed to determine the deposit’s mineralization age and the source of its metallogenic materials and mineralization processes through investigations into its geological characteristics, Rb–Sr isotopes, and S–Pb isotopes. The orebodies of the Fang’an gold deposit in the Neoarchean Xigudui Formation primarily exhibit a vein-type structure. The ore-forming process can be divided into four stages: (i) the quartz stage (Py1); (ii) the quartz–pyrite stage (Py2); (iii) the polymetallic sulfide stage (Py3); and (iv) the carbonate stage. Of these, the main mineralization stage is also the main period in which gold mineralization occurs. In situ sulfur isotope results of pyrite (Py1 to Py3) in the first three mineralization stages, suggesting a contribution of sulfur from crust–mantle magmatic fluids. The δ³⁴S values for Py2 (average 5.51‰) are higher than Py1 (average 4.45‰) and showed that the magmatic fluids mixed with meteoric waters. The δ³⁴S values for Py3 (average 5.18‰) are lower than Py2 (average 5.51‰), revealing that it related fluid immiscibility. The lead isotopic compositions of sulfides within the ores possessed ²⁰⁶Pb/²⁰⁴Pb ratios ranging from 16.759 to 16.93, ²⁰⁷Pb/²⁰⁴Pb ratios ranging from 15.311 to 15.402, and ²⁰⁸Pb/²⁰⁴Pb ratios ranging from 37.158 to 37.548. These lead data were plotted close to the Xigudui Formation, relatively distant from the Mesozoic granites, indicating that the Xigudui Formation was the source of lead for the Late Mesozoic ores of the deposit. Taken together, due to the degassing of mantle-derived magma in the shallow parts of the crust, it can be determined that the sources of ore-forming sulfur and lead were crust–mantle magmatic activities in the Wuhe area. Rb–Sr dating of pyrite from Fang’an gold deposit reveals that the mineralization occurred at 126.89 ± 0.58 Ma. Considering the previous research into the dating of magmatic rocks in the Wuhe area, we propose that the genesis of the Fang’an gold deposit is closely associated with magmatic activities in the area at around 130 Ma.

1. Introduction

The Wuhe area is located on the southeastern margin of the North China Craton (NCC) (Figure 1a) and is adjacent to the Bengbu Uplift and the Wuhe Platform Depression [1,2]. This area is characterized by ductile and brittle–ductile tectonics and magmatic activity, which are favorable for mineralization [3]. To date, quartz-vein gold deposits occur in this area [4,5,6,7,8,9], including Dagongshan, Hekou, Rongdu, Zhu Ding, and Fang’an. Previous studies have been conducted on Wuhe gold deposits, contributing to the advancement of gold mineralization theory and practical mineral prospecting in the area [4,5,6,7,8,9,10,11,12]. The Wuhe gold deposits are hosted in the Neoarchean Xigudui Formation [11]. This Formation is a source bed and provided ore-forming materials [13,14,15]. In addition, previous study of S isotope data for Wuhe gold deposits have shown that the ore-forming materials were derived from the mantle and wall rocks [14]. Recently, gold mineralization at Wuhe area has been generally constrained as 113–134 Ma corresponding to the Early Cretaceous as deduced from Rb–Sr, Re–Os, K–Ar, and Ar–Ar geochronology [4,5,7,9]. The ore-forming materials were related to the Early Cretaceous magmatism [5,9,16]. Despite many previous studies, the ore-forming materials sources remain debated in the Wuhe gold deposit, including mantle-derived materials, wall rock sources, and magmatic sources [4,5,9,13,14,15,16].

The Fang’an gold deposit was first explored in 2018 by Geological Team 312 of the Anhui Geological and Mineral Exploration Bureau, China. Nevertheless, research on the geological characteristics of the Fang’an deposit and its mineralization process remains relatively limited. In this study, we present new Rb–Sr isotope data of pyrite with a view to elucidating the timing of magmatism and gold mineralization. Additionally, we present in situ results on the sulfur isotopes of pyrite and the lead isotopes of sulfide in order to identify the sources of the ore-forming materials. By combining our new results with the published data, we attempt to formulate the process of mineralization of the Fang’an deposit.

2. Geological Characteristics

2.1. Regional Geology

The Wuhe area is located at the intersection of the south margin of the NCC and the Qinling–Dabie orogenic belt (Figure 1b). The NNE-trending Tan–Lu Fault Zone (TLFZ) passes through this area, with the Bengbu Uplift to the west. The Wuhe area is mainly covered by Quaternary sediments, and the exposed metamorphic basement comprises the Neoarchean–Paleoproterozoic Wuhe group (Ar₃wh) (Figure 1c) with a total thickness of over 6400 m. Based on its lithological characteristics, the Wuhe group is divided into five subunits: the Xigudui, Zhuangzili, Fengshanli, Xiaozhangzhuang, and Yinjiajian Formations [17]. The Wuhe group is composed substantially of greenschist-to-granulite-facies metamorphosed flysch-type sedimentary rocks and volcanic rocks [8]. Tectonically, the complex is part of the NCC basement [18] The Xigudui Formation is chiefly source bed in the Wuhe group, with the thickness (>1535 m) [14]. The Xigudui Formation primarily consists of amphibolite, amphibole plagioclase gneiss, and biotite plagiogneiss, which has undergone amphibolite facies metamorphism. Moreover, the Phanerozoic strata consist of the Cretaceous Qingshan Group (andesite), Xinzhuang Formation (clastic sediments), and Neogene–Quaternary sediments (Figure 1c).

The granitic rocks exposed in the central-to-southern part of the study area primarily consist of granites, diorites, granodiorites, and potassic granites [19], such as the Shanzhangjia, Baishishan, Nvshan, and Zhuangzili plutons (Figure 1c). These granitic rocks show clear intrusive contact with the wall rocks (the Xigudui Formation). Only a few veins have been observed in the north of the study area. Yang et al. (2010) determined that the Zhuangzili Formation comprises potassic granite and granitic amphibolite with an emplacement age of 2.1 Ga [17]. The age of the Nvshan pluton is 130.1 ± 3 Ma by zircon U–Pb dating analysis [20]. Wang et al. (2012) suggested that these magmatic intrusions are part of the metamorphic basement of the Xigudui Formation, with the remainder being products of Mesozoic magmatism [6] with an age of 120–160 Ma. Mesozoic intrusions in the Wuhe area have been divided into Late Jurassic and Early Cretaceous granitoids. Moreover, intermediate–acid and mafic dikes primarily occur in the forms of granite porphyry, granodiorite porphyry, lamprophyre, granodiorite porphyry, and diabase. The gold deposits in the Wuhe area are usually closely spatially related to Yanshanian magmatic veins [13]. The structures in the Wuhe area substantially comprise E- and NE-to-NNE-trending faults, with the nearly NNE-trending continental-scale TLFZ being the most important structure at the regional scale. The Hefei Fault and Dongzhuang–Zhouzhuang Fault mark its western and eastern boundaries, respectively (Figure 1c).

2.2. Deposit Geology

The Fang’an gold deposit located in the eastern Wuhe district (Figure 1b) [1,3], is a small-sized quartz vein-type deposit influenced by the Xigudui Formation and the NNE-trending fault structure. It contains gold reserves that are 2.269 tons, with an industrial grade of 3.67 g/t; lead reserves are 13.38 tons, with an average grade of 2.26%; silver reserves are 12.682 tons, with an average grade of 28.81 g/t.

Based on drill hole data, the Fang’an deposit is covered by the Quaternary sediments, with a depth of about 60.33~90.70 m. Strata in the Fang’an deposit consist of the Neoarchean Xigudui Formation, Late Mesozoic Qingshan and Xinzhuang Formations, and Quaternary sediments. The Fang’an deposit is located between the Wuhe–Hongxinpu fault and the Zhuding–Shimenshan fault (Figure 1c) [13], which are secondary faults of the TLFZ. Plutonic rocks, primarily composed of granitic porphyry, diorite porphyry, and diorite, have been identified within the Fang’an deposit through drill hole investigations.

The Fang’an deposit comprises two ore zones, I and II (Figure 2a). The ore-bearing structures consist of ductile–brittle shear zones with nearly N–S strike and a SE dip at 30–60°. Gold-bearing quartz veins are primarily hosted within Xigudiu Formation.

The crosscutting and replacement relation of ore minerals and gangue minerals show different mineralization stages. The photographs of ore in Figure 3 show gold mineralization assemblages and alterations from the Fang’an deposit. The ore minerals are mainly pyrite, followed by halcopyrite, sphalerite, and galena (Figure 3 and Figure 4). The gangue minerals primarily consist of quartz, calcite, and sericite. The wall rocks alterations are common in the Fang’an deposit, including pyritization, silicification, sericitization, chloritization, and carbonation (Figure 3 and Figure 4), showing medium–low-temperature hydrothermal alteration combinations. The vein minerals are dominated by quartz and calcite.

A combination of drill core observation and microscopic analysis was used to determine mineral assemblages, structures, texture, and crosscutting relationships between ore veins. The detailed paragenesis was described by Wang et al. (2022) [16]. Four stages of mineralization were noted in the Fang’an deposit (Figure 5), including the quartz–sericite stage (I), which is defined by sericite–quartz altered rocks, or pyrite–milky white quartz veins in places (Figure 3a,b). Small amounts of euhedral-to-subhedral pyrite (Py1) grains are developed (Figure 4a) and they have relatively clean surfaces (Figure 4g). In the quartz–pyrite stage (II), quartz veins containing pyrite are identified. Pyrite (Py2) occurs as coarse euhedral cubes and subhedral aggregates (Figure 3c,d). Chalcopyrite and galena generally present as irregular inclusions, or are hosted in open fractures in pyrite (Figure 4c,h). In the polymetallic sulfide stage (III), the mineral assemblage is quartz, pyrite, galena, chalcopyrite, and sphalerite. Quartz is usually dark gray (Figure 3e). Pyrite (Py3) occurs as fine-grained subhedral and anhedral aggregates (Figure 4d,e,i). In the carbonate stage (IV), the mineral assemblage is calcite, quartz, and sericite and forms veinlet structures that cut through the earlier stages (Figure 3f and Figure 4f).

3. Sample Description and Analytical Methods

3.1. Sampling and Petrographic Observations

In this study, 15 samples were collected from drill cores (ZK141 and ZK1665) and prepared as polished into sections for petrographic study. The quartz-pyrite stage ore samples, in which pyrites were coeval with galena, were selected from drill cores (ZK141) for Pb–Rb–Sr isotopic analyses. Four samples of different stages were selected from drill cores (ZK141 and ZK1665) for in situ sulfur isotope analyses. From these, single mineral selections were performed at Hebei ChengPu Testing Technology Co. (Langfang, China). The samples were washed, dried, crushed, and sieved to select single mineral samples of pyrite and galena with a purity of over 99% using binoculars. The Pb–Rb–Sr isotope testing on the samples was completed in an analytical laboratory at Beijing Research Institute of Uranium Geology (BRIUG), China. In situ sulfur isotope testing analysis of thin sections of different stages of pyrite was carried out using LA–MC–ICP–MS at the Institute of Geology, Chinese Academy of Geological Sciences (CAGS).

3.2. Rb–Sr Isotope

For the Rb–Sr isotope analyses, the selected pyrite grains were ground into powdery form, then ultrasonically cleaned for 30 min with Milli–Q water. After the drying process at a low temperature, pyrites samples (each weighing 200 mg) were completely dissolved in mixed acid, which consists of 1 mL HCl and 3 mL HNO₃ at 120 °C on an electric hot plate. Subsequently, the sample was mixed with hydrochloric acid after being dried with distillation, and the centrifugal separation of Rb, Sr, and other elements was performed with an ion-exchange column from the liquid supernatant. A Phoenix thermal surface ionization mass spectrometer was used for the determination of the Rb and Sr concentrations of the samples and their isotopic ratios. Analyses of the NBS987 standard yielded a mean ⁸⁷Sr/⁸⁶Sr ratio of 0.710233 ± 6 (2σ). The total procedure blank for Rb and Sr analyses was 3 × 10⁻⁹ g. The accuracy of the Rb/Sr ratio and Sr concentration were 1% and 10⁻⁵, respectively. More details about the analysis can be found in Li et al. (2008) and Yang et al. (2010) [22,23]. Mass spectrometry was performed using a Phoenix thermal surface ionization mass spectrometer [24], and regression analysis for the age calculation was performed using ISOPLOT 3.0 [25].

3.3. In–Situ Sulfur Isotope

In situ sulfur isotope analyses were carried out using LA–MC–ICP–MS, following the methods of Li et al. (2019) [26]. An MC–ICP–MS (Neptune Plus) equipped with a laser ablation system (Resolution S155) was used for the analyses. Helium was used as the carrier gas for the ablation cell in the laser ablation system. For single-spot analysis, the laser diameter was 32 μm with a laser repetition rate of 6 Hz. The energy fluency of the laser was approximately 5 J/cm². The Neptune Plus instrument was outfitted with nine Faraday cups (FCs) fitted with 10¹¹ Ω resistors. To reduce polyatomic interferences, all measurements were performed at medium resolution with a revolving power that was always greater than 7000. The FCs (L3, C, and H3) concurrently and statically collected signals for ³²S, ³³S, and ³⁴S. Sulfur isotope mass fractionation was corrected using the standard-sample bracketing method (SSB). The analysis used certified international standards (including Pyrite of Balmat) and working reference samples (pyrite YP-136), with the reference standard values of δ³⁴S_V–CDT reported by Crowe and Vaughan (1996) and Li et al. (2019) [26,27]. The analytical precision calculated from the replicate analyses of the samples was better than 0.5 per mil.

3.4. Lead Isotope

An ISOPROBE–T thermal surface ionization mass spectrometer was adopted for the analysis of the Pb isotopes. About 200 mg pyrite or 3 mg galena powder was transferred into a low-pressure digestion tank (PFA) and a mixture of HF–HNO₃–HClO₄ was added for 24 h. Subsequently, the solution was evaporated to dryness and then dissolved by adding HCl (6 mol/L) to transform the Pb into Pb–chloride and evaporated to dryness again. Then, Pb was separated and purified in a Bio–Rad AG–1 × 8 anion exchange resin column using HBr solution (0.5 mol/L). A total of 1 mL HCl (6 mol/L) was used to resolve the Pb in a Teflon beaker followed by distilling to dryness; the mixture was then coated on a Re filament using the H₃PO₄ and silicon gel method for analysis using an ISOPROBE–T thermal surface ionization mass spectrometer. The measured Pb isotopes were calibrated for instrument mass fractionation by reference to replicate analyses of the standard NBS 981. The determined ²⁰⁸Pb/²⁰⁶Pb, ²⁰⁷Pb/²⁰⁶Pb, and ²⁰⁴Pb/²⁰⁶Pb ratios for NBS 981 were 2.164940 ± 0.000015, 0.914338 ± 0.000007, and 0.0591107 ± 0.0000002, respectively. The total blank for Pb was <100 pg. The measurement procedures and reference materials were described in detail by Wang et al. (2017) [28].

4. Results

4.1. Rb and Sr Isotope Analysis

A total of three Rb–Sr isotope samples were selected from the Fang’an deposit, composed of pyrites in the quartz–pyrite stage. The Rb–Sr isotope analytical results for the pyrite samples are listed in

Table 1. The Rb/Sr isotope contents and ratios of pyrite in the Fang’an gold deposit.

Sample	Rb (×10−6)	Sr (×10−6)	87Rb/86Sr	87Sr/86Sr	Std Err
ZK141–H8	0.182	0.482	1.0941	0.715128	0.000029
ZK141–H11	2.53	0.905	8.0946	0.727487	0.000024
ZK141–N2	1.75	9.46	0.5361	0.714043	0.000023

. The Rb and Sr contents in the samples were 0.182–2.53 ppm and 0.482–9.46 ppm, respectively, with ⁸⁷Rb/⁸⁶Sr ratios of 0.5361–8.0946 and ⁸⁷Rb/⁸⁶Sr analytical uncertainty (2σ) values of 0.714043–0.727487, indicating that the range of ⁸⁷Rb/⁸⁶Sr ratios is relatively large, while the distribution of the analytical uncertainty is relatively concentrated. ISOPLOT 3.0 software was adopted to analyze the test results [25], with a 2σ index for the dating error. The value of ⁸⁷Rb/⁸⁶Sr yielded an Rb–Sr isochron age of 126.89 ± 0.58 Ma (MSWD = 6.6) (Figure 6c), with an initial ⁸⁷Sr/⁸⁶Sr value of 0.713127 ± 0.00004.

4.2. In Situ Sulfur Isotope of Pyrite

The δ³⁴S values of hydrothermal pyrite from the Fang’an deposit have a narrow range of 3.73‰~6.81‰ (n = 25). The isotopic data (n = 25) for in situ pyrite isotope analyzed (δ³⁴S) using LA–MC–ICP–MS are provided in

Table 2. The sulfur isotopic composition of pyrite from the Fang’an deposit and related.

Site	Sample No.	Testing Object	δ34Sv–CDT/‰	References
Fang’an	ZK141–H11–1	Py1	3.73	This study
	ZK141–H11–2		4.95
	ZK141–H11–3		4.81
	ZK141–H11–4		3.84
	ZK141–H11–5		4.93
	ZK141-K3–1–1	Py2	5.86
	ZK141-K3–1–2		6.81
	ZK141-K3–1–3		6.09
	ZK141-K3–1–4		5.69
	ZK141-K3–1–5		6.12
	ZK1665–F4–2–2		5.31
	ZK1665–F4–2–3		4.55
	ZK1665–F4–2–4		4.69
	ZK1665–F4–2–5		4.84
	ZK1665–F4–2–6		5.11
	ZK1665–F4–1–4	Py3	4.78
	ZK1665–F4–1–5		4.49
	ZK1665–F4–1–3		4.09
	ZK1665–F4–1–6		4.5
	ZK141–H1–1		5.66
	ZK141–H1–2		5.53
	ZK141–H1–3		5.66
	ZK141–H1–4		5.49
	ZK141–H1–5		5.9
	ZK141–H1–6		5.7
Rongdu		Galena	2.3	Wei (2011) [14]
		Pyrite	2.5
			6.8
		Galena	3.1
	DQT–Py–1	Pyrite	6.8	Zong et al. (2022) [29]
	DQT–Py–2		5.8
	DQT–Gn–1	Galena	3.7
	DQT–Gn–2		1.3
Hekou	1	Barren Pyrite	6.47	Liu et al. (2023) [30]
	2		6.33
	3		6.29
	4		6.37
	5		6.12
	6		6.24
	7		6.07
	8		6.02
	9		5.98
	10		6
	11		6.2
	12		6.19
	13		6.47
	14		6.58
	15		6.51
	16		6.73
	1	Auriferous Pyrite	6.08
	2		6.1
	3		5.81
	4		5.77
	5		5.83
	6		6.17
	7		5.74
	8		5.73
	9		5.65
	10		5.74
	11		5.97
	12		6.08
	13		6.4
	14		6.18
Dykes	2011–47	Lamprophyre	2.1	Yang (2020) [9]
	2011–50		1.1
Xigudui	RD–10	Plagiogneiss	4.1
	DGS–17		4
	2207–5		4.4
	2011–18	Amphibolite	2
		Pyrite	4.22	Chen (2021) [31]
			4.77
			4.33
			5.63
			4.65

, as shown in Figure 7. The δ³⁴S values of Py1 range between 3.73‰ and 4.95‰ with an average of 4.45‰ (n = 5). The δ³⁴S values of Py2 range between 4.55‰ and 6.81‰ with an average of 5.51‰ (n = 10). The δ³⁴S values of Py3 range between 4.09‰ and 5.9‰ with an average of 5.18‰ (n = 10).

4.3. Lead Isotopic Composition of Sulfide

The lead isotopic compositions of galena and pyrite from the Fang’an gold deposit are relatively uniform, as shown in Table 3 and illustrated in Figure 8 and Figure 9. The values of ²⁰⁶Pb/²⁰⁴Pb, ²⁰⁷Pb/²⁰⁴Pb, and ²⁰⁸Pb/²⁰⁴Pb of galena and pyrite, together with the ²⁰⁶Pb/²⁰⁷Pb values of galena and pyrite, are shown in Table 3, with ²⁰⁶Pb/²⁰⁴Pb values of 16.759~16.93, ²⁰⁷Pb/²⁰⁴Pb values of 15.311~15.402, ²⁰⁸Pb/²⁰⁴Pb values of 37.158~37.548, and ²⁰⁶Pb/²⁰⁷Pb values of 1.0917~1.2022. The values of μ, ω, and Th/U were determined using the single-stage Pb evolution model calculation. The average values for these parameters were 9.24, 39.18, and 4.1, respectively (Table 3). Table 3 also lists the calculated Δβ and Δγ values, including data from previous studies.

Table 3. Lead isotopic compositions of pyrite and galena from the Fang’an gold deposit and related.

Sample	Test Object	Pb Isotopic Ratios				Parameters					References
		208Pb/204Pb	207Pb/204Pb	206Pb/204Pb	206Pb/207Pb	μ	ω	Th/U	Δβ	Δγ
ZK141–H1	Galena	37.158	15.331	16.759	1.0931	9.18	38.29	4.04	0.07	−6.47	This study
ZK141–H13		37.453	15.372	16.924	1.101	9.23	39.02	4.09	2.74	1.42
ZK141–H8	Pyrite	37.47	15.397	16.881	1.0964	9.29	39.67	4.13	4.37	1.87
ZK141–H6		37.548	15.402	16.93	1.0992	9.29	39.76	4.14	4.7	3.96
BB9–1	Feldspar	38.4123	15.5937	17.9498	1.1511				17.2	27.07	Yang et al., (2019) [32]
BB9–2		38.3915	15.5896	17.9098	1.1488				16.93	26.51
BB9–3		38.3976	15.5919	17.9175	1.1492				17.08	26.67
	Whole rock	36.869	15.376	17.03	1.1076				3	−14.2	Wang et al., (2012) [6]
		37.492	15.397	17.186	1.1162				4.37	2.46
		34.54	15.659	18.826	1.2022				21.46	−76.47
		37.23	15.38	16.79	1.0917				3.26	−4.55	Sang et al., (1990) [33]
		37.7472	15.5027	17.2879	1.1152				11.27	9.28	Dong et al., (1995) [34]
		38.1016	15.5563	17.4453	1.1214				14.76	18.76
		37.6212	15.4679	17.1807	1.1107				9	5.91
		38.5558	15.4608	17.3121	1.1197				8.53	30.9
		37.8292	15.5158	17.4061	1.1218				12.12	11.48
		37.69	15.345	16.946	1.1043				0.98	7.75	Tu et al., (1992,1993) [35,36]
		37.692	15.431	17.134	1.1104				6.59	7.81
		37.692	15.384	17.273	1.1228				3.52	7.81

Abbreviations are as follows: μ = initial ²³⁸U/²⁰⁴Pb ratio; ω = initial ²³²Th/²⁰⁴Pb ratio. Δβ = (²⁰⁷Pb/²⁰⁴Pb/15.33 − 1) × 1000; Δγ = (²⁰⁸Pb/²⁰⁴Pb/37.4 − 1) × 1000. All parameters and Th/U values were determined using Geokit [37,38].

Table 3. Lead isotopic compositions of pyrite and galena from the Fang’an gold deposit and related.

Sample	Test Object	Pb Isotopic Ratios				Parameters					References
		208Pb/204Pb	207Pb/204Pb	206Pb/204Pb	206Pb/207Pb	μ	ω	Th/U	Δβ	Δγ
ZK141–H1	Galena	37.158	15.331	16.759	1.0931	9.18	38.29	4.04	0.07	−6.47	This study
ZK141–H13		37.453	15.372	16.924	1.101	9.23	39.02	4.09	2.74	1.42
ZK141–H8	Pyrite	37.47	15.397	16.881	1.0964	9.29	39.67	4.13	4.37	1.87
ZK141–H6		37.548	15.402	16.93	1.0992	9.29	39.76	4.14	4.7	3.96
BB9–1	Feldspar	38.4123	15.5937	17.9498	1.1511				17.2	27.07	Yang et al., (2019) [32]
BB9–2		38.3915	15.5896	17.9098	1.1488				16.93	26.51
BB9–3		38.3976	15.5919	17.9175	1.1492				17.08	26.67
	Whole rock	36.869	15.376	17.03	1.1076				3	−14.2	Wang et al., (2012) [6]
		37.492	15.397	17.186	1.1162				4.37	2.46
		34.54	15.659	18.826	1.2022				21.46	−76.47
		37.23	15.38	16.79	1.0917				3.26	−4.55	Sang et al., (1990) [33]
		37.7472	15.5027	17.2879	1.1152				11.27	9.28	Dong et al., (1995) [34]
		38.1016	15.5563	17.4453	1.1214				14.76	18.76
		37.6212	15.4679	17.1807	1.1107				9	5.91
		38.5558	15.4608	17.3121	1.1197				8.53	30.9
		37.8292	15.5158	17.4061	1.1218				12.12	11.48
		37.69	15.345	16.946	1.1043				0.98	7.75	Tu et al., (1992,1993) [35,36]
		37.692	15.431	17.134	1.1104				6.59	7.81
		37.692	15.384	17.273	1.1228				3.52	7.81

Abbreviations are as follows: μ = initial ²³⁸U/²⁰⁴Pb ratio; ω = initial ²³²Th/²⁰⁴Pb ratio. Δβ = (²⁰⁷Pb/²⁰⁴Pb/15.33 − 1) × 1000; Δγ = (²⁰⁸Pb/²⁰⁴Pb/37.4 − 1) × 1000. All parameters and Th/U values were determined using Geokit [37,38].

5. Discussion

5.1. Timing of Gold Mineralization

Two main methods are used to determine the mineralization age of hydrothermal deposits. One method involves the ages of wall rocks (strata or intrusions) to constrain the mineralization age. Another method is directly measuring the ages of ore minerals or altered minerals [39]. After the first attempt at geochronological Rb–Sr dating of sulfide minerals was made by Nakai et al. (1990) [40], many studies have been conducted in this regard, proving the effectiveness of using Rb–Sr methods for dating hydrothermal deposits [41,42]. This suggests that Rb–Sr isotope pyrite dating is an effective and direct method for defining the formation ages of hydrothermal sulfide deposits.

As proposed by previous studies on the binary mixed system of ⁸⁷Sr/⁸⁶Sr ratios, a set of samples participating in the isochron fit by their ⁸⁷Sr/⁸⁶Sr–Sr⁻¹ and ⁸⁷Rb/⁸⁶Sr–Rb⁻¹ values were plotted to determine whether they were mixed lines. When the samples were positively correlated, they were considered to be mixed lines; otherwise, the obtained isochronous ages were considered to be representative of the age of mineralization during the corresponding phase [43]. When the ⁸⁷Sr/⁸⁶Sr–Sr⁻¹ and ⁸⁷Rb/⁸⁶Sr–Rb⁻¹ values of the rubidium–strontium isotopes of the Fang’an gold deposit were plotted (Figure 6a,b), it was found that there was no correlation between the elements in the Fang’an gold deposit. The isochronous age obtained via calculations based on the three samples through ISOPLOT software was 126.89 ± 0.58 Ma (MSWD = 6.6) (Figure 6c), and the initial value of ⁸⁷Sr/⁸⁶Sr was 0.713127 ± 0.00004 with an MSWD value of 6.6. Then, the data were plotted using ISOPLOT software, and we found that the ⁸⁷Sr and ⁸⁷Rb values did not co-vary with ⁸⁶Sr–Sr⁻¹ and ⁸⁶Sr–Rb⁻¹, irrespective of samples (Figure 6a,b), indicating that the age was isochronal and therefore able to reflect the metallogenic age of gold mineralization. The Rb–Sr isotopic age of 126.89 ± 0.58 Ma for pyrite samples was considered to be reliable.

Previous work on the mineralization age of the gold deposits in the Wuhe area utilized various radioisotope dating methods, such as Rb–Sr, Re–Os, K–Ar, and Ar–Ar ore and/or gangue minerals, resulting in distinct isotope ages of 113–134 Ma [4,5,7,9,44]. Furthermore, the Wuhe magmatic rocks are composed of a large number of intermediate–felsic intrusions (115~130 Ma) and a few mafic dikes (116~130 Ma) [44]. Due to their close spatial and temporal relationship, the gold mineralization in the Wuhe area is associated with Early Cretaceous (115~130 Ma) magmatism.

Therefore, we conclude that the Rb–Sr isochron ages of pyrite (126.89 ± 0.58 Ma) constrain the mineralization age of the Fang’an deposit, coeval with emplacement ages of Early Cretaceous intrusions in the Wuhe area, indicating a genetic link between them.

5.2. Ore–Forming Material Source

Sulfur is one of the key elements in gold deposits. Consensus among studies suggests that gold was transported in ore-forming fluids, primarily as bisulfide complexes, and co-precipitated with sulfides, particularly pyrite [45,46]. Consequently, understanding the sulfur isotopic composition of pyrite is crucial for enhancing our knowledge of the sources of ore-forming materials and deposits [47,48,49].

The δ³⁴S values of hydrothermal pyrite in the three stages of the Fang’an deposit have a narrow range from 3.73‰ to 6.81‰ (Figure 7). There is no obvious difference in the δ³⁴S values of each stage (average 4.45‰–5.51‰–5.18‰) (Figure 7a), indicating that they came from the same source. In addition, the δ³⁴S values of the whole rock and sulfides (δ³⁴S = 2‰~5.63‰) from the Neoarchean Xigudui Formation, and lamprophyre (δ³⁴S = 1.1‰~2.1‰) from the Mesozoic intermediate–basic dikes (Figure 7b), have been previously documented. Furthermore, the δ³⁴S values fall within the narrow ranges of these rocks. It is still difficult to establish whether the sources of ore-forming sulfur were magmatic, metamorphic, or mantle.

Figure 7. (a). Histograms of δ³⁴S values of hydrothermal pyrites from the Fang’an deposit; (b). Natural sulfur isotope reservoirs (modified after [50]) and sulfur isotope composition of sulfide from deposits in the Wuhe area. The Fang’an gold deposit data are from this study; the data of other gold deposits in the Wuhe area are from Table 2.

Figure 7. (a). Histograms of δ³⁴S values of hydrothermal pyrites from the Fang’an deposit; (b). Natural sulfur isotope reservoirs (modified after [50]) and sulfur isotope composition of sulfide from deposits in the Wuhe area. The Fang’an gold deposit data are from this study; the data of other gold deposits in the Wuhe area are from Table 2.

Although the average δ³⁴S values in Fang’an are close to the value for the Xigudui Formation (δ³⁴S = 2‰~5.63‰), the source of ore-forming sulfur cannot be attributed to the metamorphic rocks of the Neoarchean Xigudui Formation, because the gold event occurred about 2 billion years later than the regional high-grade metamorphism. Consequently, these essential components would have been lost from the rocks during this extended period. Furthermore, the age of the gold mineralization (126.89 Ma) is nearly 4 Ma younger than the ages of the regional Mesozoic granites, such as the Nvshan (130 Ma) granitoids [20], thus raising the possibility of magmatic sources of ore-forming fluids.

However, our results show that the δ³⁴S values of these geologic units are comparable; in particular, the δ³⁴S (1.1‰~2.1‰) values of Mesozoic mantle-derived basic–intermediate dikes deviate from the mantle values (δ³⁴S = ~0‰). Mao et al. (2008) proposed that the similar sulfur isotopic compositions of the Mesozoic rocks imply the homogenization of the sulfur isotopic system through crust–mantle interactions [51]. Furthermore, Zhu et al. (2015), using He–Ar isotopes, proposed that during the westward subduction of the paleo–Pacific plate in the Early Cretaceous, parts of the subducted slab remained stagnant in the mantle transition zone under the eastern NCC [52]. In other words, during the Mesozoic mineralization events, it is implied that the ore-forming fluids sourced from a common reservoir were likely associated with processes of magmatic crust–mantle interactions.

In this study, our lead isotope results also confirm that the ore-forming fluids were derived from crust–mantle magmatic activity. Lead isotope compositions (particularly those of sulfide minerals, such as pyrite and galena) are an excellent proxy for constraining the sources of lead in ore-forming fluids [53,54]. Lead isotopes can fingerprint metal sources, reveal the nature of Pb reservoirs, and determine rough tectonic settings for ore formation [55,56].

The Pb isotopic compositions of sulfide minerals in the Fang’an deposit, despite their slight variations (Figure 8), define a linear array in the ²⁰⁷Pb/²⁰⁴Pb vs. ²⁰⁶Pb/²⁰⁴Pb and ²⁰⁸Pb/²⁰⁴Pb vs. ²⁰⁶Pb/²⁰⁴Pb diagrams. The samples from the Fang’an deposit are plotted near or below the average lower crust line in Figure 8a, and between the mantle and lower crust lines in Figure 8b. They were plotted close to the Xigudui Formation, relatively distant from the Mesozoic granites, indicating that the Xigudui Formation was the source of lead for the Late Mesozoic ores of the deposit.

Figure 8. Lead isotopic tectonic model for the Fang’an gold deposit, (a,b) ²⁰⁶Pb/²⁰⁴Pb vs. ²⁰⁷Pb/²⁰⁴Pb and ²⁰⁶Pb/²⁰⁴Pb vs. ²⁰⁸Pb/²⁰⁴Pb diagrams showing the comparison of the Pb isotope of the Fang’an gold deposit and regional rocks in the Wuhe area and general earth Pb reservoirs (after Zartman and Doe, 1981) [57]. Literature for the Pb isotope data of the regional metamorphic and granite rocks are listed in Table 3.

Figure 8. Lead isotopic tectonic model for the Fang’an gold deposit, (a,b) ²⁰⁶Pb/²⁰⁴Pb vs. ²⁰⁷Pb/²⁰⁴Pb and ²⁰⁶Pb/²⁰⁴Pb vs. ²⁰⁸Pb/²⁰⁴Pb diagrams showing the comparison of the Pb isotope of the Fang’an gold deposit and regional rocks in the Wuhe area and general earth Pb reservoirs (after Zartman and Doe, 1981) [57]. Literature for the Pb isotope data of the regional metamorphic and granite rocks are listed in Table 3.

The samples from the Fang’an gold deposit yielded ²⁰⁶Pb/²⁰⁷Pb values of 1.0931~1.101, with an average ²⁰⁶Pb/²⁰⁷Pb value of 1.097 (Table 3), which is similar to that of the Xigudui Formation (1.12) but contrasting with that of the Mesozoic granites (1.149). This indicates that the ore-forming lead was mainly sourced from common lead (²⁰⁶Pb/²⁰⁷Pb = 0.89~1.20) [58,59] rather than radiogenic lead. The growth in radiogenic Pb resulting from the decay of U and Th after the formation of these sulfides was negligible. The lead isotopic coefficients (U, Th, and Pb) were different in the upper crust, lower crust, and mantle. Therefore, the source of Pb can also be determined using Pb isotope parameters, such as the μ (²³⁸U/²⁰⁴Pb), ω (²³²Th/²⁰⁴Pb), and κ (Th/U) values [60,61,62,63,64]. The μ values of the Fang’an sulfides (9.18~9.29, average of 9.24) are slightly higher than the average for the mantle (8.92) [61,62,63,64,65]. In contrast, the ω values of the Fang’an sulfides (38.29~39.76, average of 39.18) are distinctly higher than the average for the crust (36.84) [62]. Low μ and high ω values are characteristic of the lower crust [60,61,62,63,64]. Moreover, the κ values of the Fang’an sulfides (4.04~4.14, average of 4.1) are also highly similar to the average for the crust (4) [66,67]. This further supports the assertion that magmatic crust–mantle sources played an important role in Fang’an ore formation. In Figure 9, the samples from the Xigudui Formation are chiefly plotted in the mantle domains (labelled 1 in Figure 9), with a few in the magmatic crust–mantle mixture in a subduction setting, i.e., the magmatic arc (labeled 3a in Figure 9). This is consistent with the previous understanding that crust–mantle magmatism resulted in the formation of the granites, followed by gold mineralization.

Figure 9. The Δβ − Δγ tectonic discrimination (after Zhu, 1998) [68]. Notation of fields: 1, mantle; 2, upper crust; 3, subduction setting mixed by upper crust and mantle (3a = magmatism, 3b = sedimentation); 4, chemical sediment; 5, seafloor hydrothermal systems; 6, middle– to high–grade metamorphism; 7, high–grade metamorphosed lower crust; 8, orogenic belts; 9, old shales in upper crust; 10, retrograde metamorphism.

Figure 9. The Δβ − Δγ tectonic discrimination (after Zhu, 1998) [68]. Notation of fields: 1, mantle; 2, upper crust; 3, subduction setting mixed by upper crust and mantle (3a = magmatism, 3b = sedimentation); 4, chemical sediment; 5, seafloor hydrothermal systems; 6, middle– to high–grade metamorphism; 7, high–grade metamorphosed lower crust; 8, orogenic belts; 9, old shales in upper crust; 10, retrograde metamorphism.

In summary, due to the degassing of mantle-derived magmas in the shallow parts of the crust, it can be determined that the source of ore-forming sulfur and lead was crust–mantle magmatic activity in the Wuhe area.

5.3. Ore Genesis

In the Early Cretaceous period, the westward subduction of the paleo–Pacific plate caused parts of the subducted slab to remain stagnant in the mantle transition zone under the eastern NCC [69,70,71]. This led to the intense metasomatism of the lithospheric mantle and extensive crust–mantle magmatism in the Wuhe area, increasing oxygen fugacity (ƒ_O2) and leading to the enrichment of ³⁴S [52,72]. The δ³⁴S values of early Py1 (average 4.45‰) are obviously higher than that of the primitive mantle (δ³⁴S = ~0‰) and contemporaneous mantle-derived lamprophyre (δ³⁴S = 1.1‰~2.1‰). This indicates the presence of a more ³⁴S-rich source, which is related to the degassing and intense metasomatism of the lithospheric mantle [49,73]. And this change can be seen in the in situ S isotope analysis on pyrite from different stages (Figure 7a), revealing an increase in δ³⁴S values from Py 1 (average 4.45‰) to Py 2 (average 5.51‰). When ore-forming fluids transformed to stage II, the magmatic fluids mixed with meteoric waters characterized by higher ƒ_O2 and lower temperatures, resulting in decreasing temperature and increasing ƒ_O2 of the ore-forming fluids [49,74]. Recently, O and C-O isotope data have confirmed that the initially magmatic ore-forming fluids were subsequently mixed with meteoric water [15]. Meanwhile, the Rb–Sr isochron dating results (126.89 ± 0.58 Ma) are consistent with Early Cretaceous (115~130 Ma) magmatism in the Wuhe area [44]. This age coincidence allows us to propose that the sources of the ore-forming fluids were closely related to crust–mantle magmatism. Notably, the decline in δ³⁴S values from Py 2(average 5.51‰) to Py 3 (average 5.18‰) was related to fluid immiscibility [75,76]. The study by Liu (2023) on the Hekou deposit finds that the coexistence of liquid-rich and CO₂-rich fluid inclusions in quartz near pyrite provides evidence of fluid immiscibility in NaCl-H₂O-CO₂ fluids during the polymetallic sulfide stage [30,75,76]. When fluid immiscibility occurs, pyrite not only captures more precious and base metals, but also exhibits characteristics of sulfide precipitation [77,78,79]. As shown in Figure 4, the surfaces of pyrite (Py2 and Py3) display abundant sulfide inclusions. When ore-forming fluid transformed to Stage II and Stage III, characterized by temperature decrease, pressure fluctuation, and fluid immiscibility, the H₂S was fractionated strongly into the vapor phase and the activity of HS⁻ lowered. This process might have led to the decomposition of Au(HS)—2 and enrichment of the gold [74,80,81].

Taken together, we can conclude that the formation of the Fang’an deposit was associated with crust–mantle magmatism in the 115~130 Ma period, and thus, it can be classified as a magmatic (crust–mantle)–hydrothermal-type deposit formed during the Early Cretaceous. During ca. 130 to 115 Ma, significant lithospheric thinning in the NCC was caused by the subduction of the paleo–Pacific plate. This period of tectonic activity along the TLFZ was marked by asthenospheric underplating and the emplacement of numerous mantle- and crust-related intrusions into the shallow crust. These magmatic fluids migrated upwards along fractures or through highly permeable strata to a lower pressure level. Meanwhile, meteoric water entered the magmatic fluid, leading to fluid mixing, a decrease in temperature and pressure, and the fluid immiscibility within the ore-forming hydrothermal systems, give rise to the precipitation of native Au, and polymetallic sulfides in the Fang’an deposit.

6. Conclusions

Based on a study of the geological characteristics of the Fang’an deposit, the S–Pb isotopes of the sulfides, and the Rb–Sr isotopic composition of the pyrite, the following conclusions are drawn:

(1) The S–Pb isotope data indicates that the ore-forming material sources were derived from crust–mantle magmatic activity. Py1 has higher δ³⁴S values, suggesting degassing and intense metasomatism of the lithospheric mantle during the early mineralization stage. Py2 (average 5.51‰) are higher than Py1 (average 4.45‰), indicating that the magmatic fluids mixed with meteoric waters. The δ³⁴S values for Py3 (average 5.18‰) are lower than Py2 (average 5.51‰), revealing that it is related to fluid immiscibility.

(2) The mineralization age of the Fang’an deposit is 126.89 ± 0.58 Ma, closely associated with magmatic activity.

(3) The Fang’an deposit and other deposits in the Wuhe area have similar geochronology, ore-forming source systems, and stable isotope compositions, indicating that they belong to the same magmatic–hydrothermal system.