Genesis and Evolution of the Qieliekeqi Siderite Deposit in the West Kunlun Orogen: Constraints from Geochemistry, Zircon U–Pb Geochronology, and Carbon–Oxygen Isotopes

Abstract

The Qieliekeqi siderite deposit, located in the Tashkurgan block of western Kunlun, is a carbonate-hosted iron deposit with hydrothermal sedimentary features. This study integrates whole-rock geochemistry, stable isotopes, and zircon U–Pb–Hf data to investigate its metallogenic evolution. Coarse-grained siderite samples, formed in deeper water, exhibit average Al₂O₃/TiO₂ ratios of 29.14, δEu of 2.69, and δCe of 0.83, indicating hydrothermal fluid dominance with limited seawater mixing. Banded samples from shallower settings show an average Al₂O₃/TiO₂ of 17.07, δEu of 3.18, and δCe of 0.94, suggesting stronger seawater interaction under oxidizing conditions. Both types are enriched in Mn, Co, and Ba, with low Ti and Al contents. Stable isotope results (δ¹³C_PDB = −6.0‰ to −4.6‰; δ¹⁸O_SMOW = 16.0‰ to 16.9‰) point to seawater-dominated fluids with minor magmatic and meteoric contributions, formed under open-system conditions at avg. temperatures of 53 to 58 °C. Zircon U–Pb dating yields an age of 211.01 ± 0.82 Ma, with an average εHf(t) of −3.94, indicating derivation from the partially melted ancient crust. These results support a two-stage model involving Late Cambrian hydrothermal sedimentation and Late Triassic magmatic overprinting.

1. Introduction

Iron ore is a key resource for modern metallurgy and infrastructure, and it remains a major focus in global resource strategies. The West Kunlun Orogen, one of China’s major metallogenic belts, has recently yielded important discoveries. The Zankan magnetite deposit, located in the Tashkurgan region, is one of the most significant and has greatly increased the area’s known resource potential [1]. In addition to magnetite, a north–south-trending belt of copper-bearing siderite has also been identified. The Qieliekeqi siderite deposit is the largest and most economically important in this belt, containing 147 million tonnes of resources at the 333 + 334 classification level, as reported by the No.2 Geological Team of the Xinjiang Geological and Mineral Resources Bureau.

The Qieliekeqi deposit was first discovered in 1959 by the Xinjiang Metallurgical Bureau in the Tashkurgan area. It was initially classified as a medium- to low-temperature hydrothermal iron deposit. In 1980, further exploration by the Xinjiang Geological and Mineral Resources Bureau suggested that the deposit might be of sedimentary–metamorphic origin. Since the early 21st century, continued studies on the geological features and fluid evolution of the deposit have led to diverse interpretations of its genesis. Some researchers, based on the presence of siliceous bands and high quartz content, have proposed that the deposit formed through a seafloor exhalative process, most consistent with a sedimentary-exhalative (SEDEX-type) hydrothermal system [2]. Others, using fluid inclusion evidence, have suggested a multi-stage mineralization process involving initial hydrothermal sedimentation followed by overprinting by magmatic-hydrothermal fluids [3,4]. A third group has emphasized the role of the host strata, composed of carbonate and clastic rocks, supporting a combined sedimentary–metamorphic model [5,6,7]. These differing views reflect the complex nature of the Qieliekeqi system, but each model has limitations: studies favoring an exhalative origin lack isotopic evidence [3,4], while those supporting a sedimentary–metamorphic model often lack reliable geochronological constraints [5,6,7]. To address these issues, this study employs a combined geochemical and isotopic approach. Carbon and oxygen isotope compositions are used to distinguish among magmatic, basinal, and seawater-derived fluids, whereas rare earth element (REE) patterns serve to evaluate hydrothermal versus seawater influence and constrain fluid evolution pathways. The timing of mineralization is also debated. Traditionally, the deposit was thought to be hosted in the Silurian Wenquangou Formation [8,9,10,11,12]. However, recent zircon U–Pb dating of granitic intrusions cutting this formation has yielded Late Cambrian ages [13], suggesting that the so-called “Wenquangou Formation” may instead date from the Mesoproterozoic to early Paleozoic. To further assess the stratigraphic and tectonic context, this study applies zircon U–Pb geochronology and Lu–Hf isotopic analysis to represent intrusive rocks. While the U–Pb ages reflect a later magmatic–hydrothermal overprint rather than the primary mineralization event, they provide maximum age constraints on the host sequence. In addition, Lu–Hf two-stage model ages help to trace crustal evolution trends and magmatic sources, offering indirect but meaningful insights into the provenance and tectonic framework of the deposit. If the mineralization at Qieliekeqi is indeed coeval with the Zankan–Mokaer system [14,15], this would indicate formation during a regional tectono-magmatic phase potentially linked to back-arc extension or post-collisional collapse in the West Kunlun Orogen.

There is still a lack of systematic geochemical and isotopic evidence to resolve key questions regarding the Qieliekeqi deposit, including the origin of ore-forming fluids, the nature of the hydrothermal system, the timing of mineralization, and the depositional environment. In addition, previous studies have proposed divergent genetic models, and the overall metallogenic process remains poorly understood. To address these issues, this study integrates field observations with petrographic and mineralogical analysis, whole-rock geochemistry, carbon–oxygen isotope analysis, and zircon U–Pb geochronology of associated intrusive rocks. The goals are to (1) determine the source of ore-forming materials and the hydrothermal mechanisms involved; (2) constrain the timing of mineralization and its relationship with regional magmatic–tectonic evolution; and (3) establish a genetic model that can guide future exploration and enhance understanding of regional metallogeny.

2. Regional Geological Setting

The Tashkurgan area is rich in mineral resources and forms an important metallogenic belt within the West Kunlun–Pamir tectonic system. The region is dominated by skarn-type and sedimentary–metamorphic iron deposits, which together constitute the West Kunlun Tashkurgan iron ore belt. The proven reserves exceed 300 million tons, with a resource potential estimated at over 1.5 billion tons. The Qieliekeqi deposit is one of the major iron ore deposits in this region. It is located in Bulunkou Town, Aketao County, Kizil Autonomous Prefecture, Xinjiang, covering an area of approximately 14.6 km². Structurally, the deposit lies in the northern part of the arcuate Tashkurgan block, which trends northwest–southeast (Figure 1). The Tashkurgan block is situated on the eastern margin of the Pamir Plateau in southwestern Xinjiang, China. It represents a relatively isolated and strongly deformed tectonic unit, bounded by the Karakorum Fault to the southwest and the Kangxiwa Fault to the northeast, which separate it from the Mingtiegai and West Kunlun blocks. The block has undergone a complex geological history, including Paleozoic ocean basin closure, continental collision, and multiple phases of magmatic and tectonic activity. These processes led to widespread regional metamorphism, intense magmatism, and repeated structural overprinting, providing a favorable geological setting for mineralization.

The main stratigraphic units exposed in the study area include the Paleoproterozoic Bulunkuole Group (the age of which remains debated) and the previously assigned Silurian Wenquangou Formation. The Bulunkuole Group is in fault contact with the ore-hosting strata and is mainly exposed to the north and south of the orebody. It consists of medium- to high-grade metamorphic rocks rich in garnet. The Wenquangou Formation, which hosts the ore, was originally considered to be of Silurian age [3,5,17,19,20]. However, recent zircon U–Pb dating suggests that its depositional age may be Late Cambrian [13,20], indicating that this unit could be older than previously recognized. This discrepancy in age assignments represents a key unresolved issue in the regional stratigraphy of the Tashkurgan area. Clarifying whether the Wenquangou Formation belongs to the Silurian or an older Cambrian/Mesoproterozoic sequence is crucial for reconstructing the stratigraphic framework and the timing of mineralization. One of the main objectives of this study is to provide new geochronological constraints to address this issue. The Wenquangou Formation is characterized by thick, layered sequences of carbonate and clastic rocks, which have undergone regional metamorphism and later contact metamorphism related to granitic intrusions during the Permian.

The structural framework of the area is relatively simple, with the ore-bearing strata occurring as a monocline. These strata lie within a tectonic setting that experienced Paleozoic ocean closure and subsequent continental collision, which facilitated the development of large-scale fractures and deep-seated fluid pathways. These structures likely controlled both the spatial distribution of ore-bearing horizons and the episodic inflow of mineralizing hydrothermal fluids. The orebody is hosted in the Wenquangou Formation, appearing as stratiform or quasi-stratiform layers, with some portions extending along bedding planes. Additionally, vein-type and vuggy-textured ores formed by late-stage hydrothermal overprinting are present, indicating that the deposit has undergone multiple phases of mineralization and subsequent modification [17]. The stratiform and quasi-stratiform ores are interpreted to reflect early syn-sedimentary or diagenetic siderite precipitation under reducing conditions, whereas the late-stage vein-type and vuggy-textured ores are attributed to hydrothermal overprinting and remobilization during post-depositional tectonothermal events. This mineralization pattern suggests a polygenetic origin involving both sedimentary and hydrothermal processes.

Magmatic activity plays a significant role in the geological evolution of the region. The area is characterized by strong intermediate to acidic magmatic intrusions, with acidic rocks being dominant, while alkaline, intermediate, and mafic to ultramafic rocks are relatively rare. The most active magmatic events occurred during the Variscan and Indosinian periods, followed by less intense episodes in the Yanshanian and Himalayan periods. Caledonian magmatism was relatively weak. These multiphase magmatic events provided sustained thermal energy and fluid sources for the thermal evolution of the strata and associated hydrothermal processes [3,18].

3. Geological Characteristics of the Mining Area and Orebody

The exposed strata in the study area correspond to the Wenquangou Formation, which was previously assigned to the Silurian. The lithology is relatively simple and is dominated by a thick sequence of clastic–carbonate rocks that have undergone greenschist-facies metamorphism, with a total thickness of several kilometers. The main lithologies include biotite–quartz schist, sericite–quartz schist, and marble, which together constitute the primary host rocks for the stratiform siderite orebodies (Figure 2 and Figure 3). Based on lithological assemblages and their spatial distribution, this rock sequence can be subdivided into seven lithological formations, with the orebodies mainly hosted in the upper three, showing clear stratiform mineralization features [17]. The key characteristics of the ore-bearing formations are as follows: (1) The first formation, located in the southern part of the deposit, comprises gray to dark-gray biotite–quartz schist and chloritized muscovite–quartz schist with irregular lenses of marble. Orebody II is hosted in this formation; (2) the second formation, located centrally, consists of gray-white marble rich in quartz and muscovite. Orebodies I and IV are hosted in this formation; (3) the third formation, distributed in the northeastern and northwestern parts of the area, includes gray to dark-gray garnet-bearing biotite–quartz schist and chloritized muscovite–quartz schist (Figure 2). Orebody III is hosted in this formation. These lithological assemblages suggest that the ore-bearing strata have experienced multiple phases of metamorphism and hydrothermal alteration. Local enrichment in carbonate rocks provided favorable material sources and physical space for subsequent ore formation.

Structurally, the deposit exhibits a relatively simple framework [3,4]. The strata display a monoclinal attitude, and no major folds or faults have been observed within the ore zone. Two sets of shear joints are commonly developed in both the host rocks and the ore layers. Some of these fractures are filled with late-stage hydrothermal siderite, indicating that the deposit underwent post-ore structural and hydrothermal overprinting to a certain extent.

The intrusive rocks in the study area are primarily located along the eastern margin of the Qiukutai pluton, which occupies approximately one-third of the mapped area [17,20]. This pluton is interpreted to have formed during the Late Triassic and represents a significant magmatic event in the region. The intrusive suite includes grayish-white biotite plagiogranite, plagiogranite dikes, plagiogranite pegmatite, and quartz diorite dikes. These intrusions occur either concordantly or discordantly with respect to the host strata. Siderite enclaves are frequently observed near the contact zones, suggesting that the magmatic activity contributed to the modification or partial remobilization of the pre-existing siderite orebodies.

Siderite is the dominant ore mineral in the deposit. Common ore textures include massive, banded, vein-type, and disseminated structures [17,22]. The massive and banded ores are primarily the result of original sedimentary processes and display clear stratigraphic conformity. In contrast, the vein-type and disseminated ores are typically products of later metamorphic and hydrothermal overprinting, reflecting a multi-stage mineralization history. The samples collected for this study are mainly from the massive and banded ore types (Figure 4), with the aim of reconstructing the ore-forming environment, identifying the source of ore-forming fluids, and developing a genetic model for the deposit.

4. Samples and Analytical Methods

4.1. Sample Collection

A total of 54 rock samples were collected from the Qieliekeqi siderite deposit, including typical massive coarse-grained siderite, banded siderite, and acidic intrusive rocks exposed within or near the ore-bearing strata. Sampling was conducted at fresh outcrops with intact structures and no visible signs of weathering or alteration. All samples were immediately sealed in plastic bags to minimize weathering and cross-contamination during transportation and storage. In the laboratory, the samples were classified based on ore type and lithological characteristics. Mineral assemblages and textures were preliminarily identified under a polarizing microscope. From this collection, six representative coarse-grained siderite samples (QLIII-8, QLIII-11, QLIII-12, QLIV-2, QLIV-5, QLIV-8), two banded siderite samples (QLIII-15, QLKQ-8), and four acidic intrusive rock samples (QLIII-1, QLIII-2, QLKQ-2, QLIV-7) were selected for subsequent geochemical and stable isotope analyses.

4.2. Analytical Methods

Major and trace element analyses in this study were conducted by ALS Chemex (Guangzhou) Co., Ltd, located at Fengshu American Industrial Park, Huadu District, Guangzhou, China. To ensure the accuracy and representativeness of the results, all samples were manually selected to remove weathered surfaces, then ultrasonically cleaned, dried at low temperature, and ground to a powder finer than 200 mesh. Major elements were analyzed using X-ray fluorescence spectrometry (XRF). For each sample, 0.7 g of powder was mixed with a specialized flux and fused into a glass bead for analysis. Standard reference materials, including BCR-2, GSR-1, and GSR-3, were used for calibration, and analytical precision and accuracy were both better than 5%. Trace elements were determined using inductively coupled plasma mass spectrometry (ICP-MS). For each analysis, 50 g of powdered sample was digested in Teflon containers using high-purity HNO₃, HF, and HClO₄ in two stages. Rhodium (Rh) was added as an internal standard, and the final solution was diluted to 80 g for analysis. AGV-1 and BHVO-1 were used as reference standards, with analytical precision and accuracy better than 10%.

Zircon U–Pb and Lu–Hf isotopic analyses included two main steps: zircon separation and mounting, followed by laser ablation multi-collector inductively coupled plasma mass spectrometry (LA-MC-ICP-MS). Zircon separation and target preparation were conducted by Chengxin Geological Services Co., Ltd., located at No. 3 Yuhua Road, Langfang, Hebei Province, China. Isotopic measurements were performed at the Institute of Mineral Resources, Chinese Academy of Geological Sciences, using a Finnigan Neptune MC-ICP-MS instrument coupled with a NewWave UP213 UV laser ablation system. For U–Pb dating, the Temora zircon standard (417 Ma [23]) was used as the primary reference material. The analytical precision was better than ±1% (2σ). Isotopic ratios and age calculations were processed using the ICPMS DataCal software, and Concordia and weighted mean age diagrams were generated using Isoplot 3.0. Lu–Hf isotope analyses were conducted on the same spots as U–Pb analyses. External standards included GJ-1 and MUD zircons. The analytical results are consistent with the solution values recommended by Woodhead et al. [24] and fall within the accepted analytical uncertainty.

Carbon (C) and oxygen (O) stable isotope analyses were conducted at the Analytical and Testing Center of the Beijing Research Institute of Uranium Geology. Pure siderite samples were selected and ground to below 200 mesh. Approximately 20 g of dried powder was reacted with phosphoric acid in a thermostatic shaking water bath to release CO₂, which was then collected by cryogenic trapping. The δ¹³C and δ¹⁸O values were measured using an isotope ratio mass spectrometer (IRMS). All results were calibrated against working standards and corrected to the international PDB scale using computer software. Each sample was measured 6–8 times, and the average value was used for data interpretation.

5. Results

5.1. Major and Trace Elements

Table 1. Major and trace element contents of samples from the Qieliekeqi deposit.

Sample Number	QLIII-11	QLIV-5	QLIV-8	QLIII-15	QLKQ-8
Ore Type	Coarse-Grained	Coarse-Grained	Coarse-Grained	Banded	Banded
Major Element (wt%)
SiO2	22.20	10.70	7.17	11.50	13.65
TiO2	0.04	0.10	0.06	0.16	0.14
Al2O3	1.20	2.36	2.03	2.63	2.48
Fe2O3 (Total)	50.55	54.10	53.53	54.18	52.90
MnO	1.58	1.82	1.86	1.76	1.75
MgO	2.63	3.07	4.72	2.28	2.99
CaO	0.36	0.74	0.94	0.28	0.76
Na2O	0.02	0.03	0.03	0.03	0.02
K2O	0.37	0.80	0.65	0.86	0.78
P2O5	0.07	0.03	0.03	0.04	0.06
SO3	0.41	0.48	0.22	0.34	0.08
LOI	20.50	25.65	28.62	25.79	24.24
Trace Element (ppm)
Li	0.60	2.50	0.50	1.80	1.10
Be	0.21	0.47	0.50	0.40	0.72
Sc	6.00	2.90	21.00	2.50	3.50
V	15.00	43.00	34.00	29.00	32.00
Cr	10.00	10.00	10.00	10.00	10.00
Co	9.70	4.50	3.90	8.80	3.20
Ni	38.60	33.30	27.30	39.40	11.70
Cu	0.90	4.60	1.00	5.90	0.70
Zn	4.00	4.00	3.00	3.00	2.00
Ga	4.50	5.40	5.10	6.20	5.40
Ge	0.92	0.61	0.66	0.62	0.85
As	0.50	4.70	3.50	4.60	6.10
Rb	10.90	26.70	20.50	28.70	22.30
Sr	16.90	14.10	11.90	36.10	7.60
Y	9.20	10.30	19.20	18.40	14.50
Zr	10.00	29.00	24.00	29.00	25.00
Nb	0.90	3.00	1.60	4.40	3.20
Mo	1.54	1.15	0.50	0.53	1.24
In	0.87	1.30	0.84	0.55	1.44
Sn	1.00	4.00	2.00	5.00	2.00
Sb	0.35	0.50	0.62	0.75	1.52
Cs	0.23	1.02	0.53	0.84	0.54
Ba	17.30	92.00	22.20	108.00	46.00
La	117.00	1.50	0.70	<b.d.l.	4.00
Ce	143.00	2.60	1.60	0.80	10.60
Pr	12.85	0.37	0.27	0.14	1.68
Nd	30.30	1.40	1.40	0.70	7.40
Sm	2.55	0.61	1.08	0.73	2.92
Eu	1.22	0.49	0.89	0.86	2.08
Gd	1.23	1.09	2.51	2.12	3.39
Tb	0.24	0.24	0.54	0.50	0.55
Dy	1.38	1.70	3.21	3.08	2.86
Ho	0.34	0.41	0.72	0.68	0.57
Er	1.02	1.30	1.88	1.57	1.37
Tm	0.22	0.22	0.26	0.20	0.19
Yb	1.46	1.54	1.59	1.12	1.12
Lu	0.22	0.22	0.21	0.12	0.16
Hf	0.20	0.70	0.60	0.80	0.70
Ta	0.30	0.30	0.30	0.50	0.40
W	2.00	24.00	16.00	25.00	23.00
Tl	0.02	0.10	0.04	0.11	0.03
Pb	2.10	1.30	1.60	0.90	1.10
Bi	0.07	0.71	0.28	1.06	0.09
Th	4.28	1.53	1.16	1.71	2.24
U	5.06	4.08	1.33	3.34	1.40

Note: Fe₂O₃ (Total): total iron content calculated as Fe₂O₃; “< b.d.l.” indicates concentrations below the instrument detection limit (b.d.l.), which is 0.5 ppm for La.

presents the results of major and trace element analyses.

Coarse-grained siderite ore samples (QLIII-11, QLIV-5, QLIV-8) show FeO contents ranging from 45.48% to 48.68%, with an average of 47.44%, indicating a distinctly iron-rich composition. SiO₂ contents vary widely from 7.17% to 22.20%, averaging 13.40%, which may reflect the influence of detrital material from wall rocks or later hydrothermal alteration. MgO contents range from 2.63% to 4.72%, with an average of 3.47%. MnO contents are relatively stable, between 1.58% and 1.86%, with an average of 1.75%. Both TiO₂ and Al₂O₃ contents are low, ranging from 0.04% to 0.10% (average 0.07%) and 1.20% to 2.36% (average 1.86%), respectively.

Compared to the coarse-grained siderite ores, the banded siderite ore samples (QLIII-15, QLKQ-8) have slightly higher FeO contents, ranging from 47.60.90% to 48.75%, with an average of 48.18%. SiO₂ contents are relatively consistent between the two groups, with the banded ores showing a narrower range (11.50% to 13.65%) and an average of 12.58%. MgO contents are lower than in the coarse-grained samples, ranging from 2.28% to 2.99% (average 2.64%). MnO contents are nearly identical between the two groups, with the banded ores ranging narrowly from 1.75% to 1.76% (average 1.76%). In contrast, TiO₂ and Al₂O₃ contents in the banded ores tend to be higher, ranging from 0.14% to 0.16% (average 0.15%) and 2.48% to 2.63% (average 2.56%), respectively.

Rare earth element (REE) analysis shows that the coarse-grained siderite samples (QLIV-5 and QLIV-8) contain total REE concentrations (ΣREE) ranging from 13.69 to 16.48 ppm, with an average of 15.28 ppm. Light REE (ΣLREE) contents range from 5.94 to 6.97 ppm (average 6.46 ppm), while heavy REE (ΣHREE) contents range from 6.72 to 10.92 ppm (average 8.82 ppm). The LREE/HREE ratios range from 0.54 to 1.04, averaging 0.79, indicating a pronounced depletion of light REEs. PAAS-normalized REE distribution patterns exhibit a distinct rightward (HREE-enriched) slope (Figure 5), with (La/Yb)_N values ranging from 0.03 to 0.07 (average 0.05). In addition, these samples display a marked positive Eu anomaly (δEu = 2.55 to 2.83; average 2.69) and a slight negative Ce anomaly (δCe = 0.81 to 0.85; average 0.83). Notably, sample QLIII-11 is enriched in LREEs relative to the other coarse-grained samples, with negligible differences observed in MREE–HREE patterns. It shows an anomalously high ΣREE value of 313.03 ppm, with ΣLREE as high as 306.92 ppm and ΣHREE of only 6.11 ppm, yielding an extremely high LREE/HREE ratio of 50.23. The REE pattern shows a slight rightward slope with a (La/Yb)_N value of 5.92, a very strong positive Eu anomaly (δEu = 3.24), and a similarly negative Ce anomaly (δCe = 0.85).

The banded siderite samples (QLIII-15, QLKQ-8) have total REE concentrations (ΣREE) ranging from 12.62 to 38.89 ppm, with an average of 25.76 ppm. LREE contents range from 3.23 to 28.68 ppm (average 15.96 ppm), and HREE contents range from 9.39 to 10.21 ppm (average 9.80 ppm). The LREE/HREE ratios vary from 0.34 to 2.81, with an average of 1.58. Similar to the coarse-grained samples, the PAAS-normalized REE patterns also show a rightward slope (Figure 5), with (La/Yb)_N values ranging from 0.00 (La concentrations were below the detection limit) to 0.26 (average 0.13), indicating distinct fractionation between LREEs and HREEs. The Eu anomaly is more pronounced in these samples (δEu = 3.11 to 3.26; average 3.18). The Ce anomaly, however, was only calculated for sample QLKQ-8 (δCe ≈ 0.94), as La contents in QLIII-15 were below detection limits.

The coarse-grained and banded siderite ores exhibit some compositional differences, particularly in their light rare earth element (LREE) signatures. While major elements and middle-to-heavy REEs (MREEs–HREEs) are broadly comparable, the coarse-grained samples are generally iron-rich, LREE-depleted, and low in Al and Ti. In contrast, the banded ores tend to show higher contents of Si, Al, and Ti; slightly less pronounced LREE depletion; and relatively stronger positive Eu anomalies. Both ore types exhibit positive Eu anomalies and variable Ce anomalies, although sample-level variability is notable due to the limited dataset.

5.2. U–Pb and Lu–Hf Isotopes

Zircon U–Pb dating was conducted on four intrusive rock samples from the Qieliekeqi deposit (QLIII-1, QLIII-2, QLKQ-2, and QLIV-7), and Lu–Hf isotopic analysis was additionally performed on sample QLIII-1. The analytical results are summarized in

Table 2. U–Pb Isotope contents of samples from the Qieliekeqi deposit.

SampleNumber	Pb	U	206Pb/238U	1σ	207Pb/235U	1σ	207Pb/206Pb	1σ	208Pb/232Th	1σ	232Th/238U	206Pb/238U	1σ	207Pb/235U	1σ	207Pb/206Pb	1σ
	Concentration (ppm)		Isotopic Ratio									Age (Ma)
QLIII-1-1	26.9	808.5	0.0334	0.0002	0.2321	0.0047	0.0504	0.0010	0.0094	0.0002	0.3522	211.9	1.2	211.9	4.3	211.9	46.5
QLIII-1-2	12.7	382.4	0.0334	0.0002	0.2313	0.0106	0.0503	0.0022	0.0094	0.0003	0.3534	211.6	1.3	211.2	9.7	206.9	101.1
QLIII-1-3	84.3	2595.0	0.0339	0.0002	0.2349	0.0043	0.0503	0.0009	0.0116	0.0002	0.1641	214.8	1.2	214.2	3.9	207.8	39.5
QLIII-1-4	12.1	355.8	0.0339	0.0003	0.2336	0.0154	0.0500	0.0032	0.0086	0.0003	0.4131	214.8	1.6	213.2	14.0	195.5	149.7
QLIII-1-5	11.1	322.7	0.0340	0.0002	0.2349	0.0142	0.0501	0.0030	0.0108	0.0002	0.3698	215.5	1.5	214.2	12.9	200.3	139.5
QLIII-1-6	23.3	707.0	0.0337	0.0002	0.2347	0.0060	0.0505	0.0013	0.0075	0.0001	0.3561	213.8	1.2	214.1	5.5	216.6	57.7
QLIII-1-7	25.8	798.3	0.0338	0.0002	0.2346	0.0050	0.0504	0.0011	0.0094	0.0002	0.1982	214.1	1.2	214.0	4.5	212.8	48.7
QLIII-1-8	13.3	405.6	0.0336	0.0002	0.2333	0.0091	0.0504	0.0019	0.0098	0.0003	0.2682	212.9	1.2	212.9	8.3	213.5	89.0
QLIII-1-9	26.4	825.0	0.0335	0.0002	0.2348	0.0041	0.0508	0.0009	0.0087	0.0001	0.2003	212.5	1.2	214.1	3.7	231.7	39.3
QLIII-1-10	18.0	548.3	0.0334	0.0002	0.2335	0.0059	0.0506	0.0013	0.0072	0.0001	0.3947	212.0	1.2	213.1	5.4	224.9	57.2
QLIII-1-11	71.1	2230.4	0.0338	0.0002	0.2345	0.0034	0.0503	0.0007	0.0093	0.0001	0.1428	214.2	1.2	213.9	3.1	210.6	32.1
QLIII-1-12	14.4	416.9	0.0335	0.0002	0.2328	0.0078	0.0504	0.0017	0.0105	0.0002	0.4449	212.4	1.2	212.5	7.1	214.0	75.9
QLIII-1-13	70.8	975.4	0.0697	0.0004	0.5336	0.0073	0.0556	0.0007	0.0162	0.0000	0.6344	434.1	2.4	434.2	6.0	434.8	28.8
QLIII-1-14	14.0	410.3	0.0334	0.0002	0.2334	0.0096	0.0507	0.0020	0.0089	0.0002	0.4649	212.0	1.3	213.0	8.8	225.0	92.4
QLIII-1-15	10.1	287.1	0.0336	0.0002	0.2346	0.0094	0.0507	0.0019	0.0106	0.0003	0.4826	212.8	1.3	214.0	8.6	226.8	88.0
QLIII-1-16	58.8	1708.6	0.0340	0.0002	0.2346	0.0052	0.0501	0.0010	0.0172	0.0006	0.2356	215.3	1.2	214.0	4.8	199.2	48.2
QLIII-1-17	9.5	281.0	0.0335	0.0002	0.2329	0.0104	0.0505	0.0022	0.0089	0.0002	0.4223	212.2	1.3	212.6	9.5	216.7	101.9
QLIII-1-18	11.3	313.3	0.0334	0.0002	0.2333	0.0105	0.0506	0.0023	0.0104	0.0001	0.6259	212.0	1.3	212.9	9.6	222.8	103.2
QLIII-1-19	40.3	1250.5	0.0336	0.0002	0.2362	0.0038	0.0509	0.0008	0.0103	0.0001	0.1798	213.3	1.2	215.3	3.5	236.5	36.0
QLIII-1-20	55.1	1677.4	0.0335	0.0002	0.2354	0.0033	0.0509	0.0007	0.0113	0.0001	0.2354	212.5	1.2	214.6	3.0	238.1	31.6
QLIII-1-21	22.8	600.2	0.0338	0.0002	0.2345	0.0113	0.0504	0.0022	0.0270	0.0016	0.3075	214.1	1.4	213.9	10.3	212.0	101.4
QLIII-1-22	17.6	527.5	0.0335	0.0002	0.2340	0.0070	0.0506	0.0015	0.0124	0.0003	0.2666	212.5	1.2	213.5	6.4	224.2	66.5
QLIII-1-23	27.5	831.3	0.0334	0.0002	0.2349	0.0045	0.0510	0.0009	0.0121	0.0001	0.2565	211.9	1.2	214.3	4.1	240.3	42.5
QLIII-1-24	20.2	579.3	0.0333	0.0002	0.2340	0.0093	0.0510	0.0020	0.0118	0.0002	0.4550	210.9	1.2	213.5	8.5	243.0	90.3
QLIII-1-25	10.1	291.7	0.0335	0.0002	0.2342	0.0106	0.0507	0.0023	0.0134	0.0003	0.3612	212.2	1.3	213.6	9.7	229.4	103.5
QLIII-1-26	14.8	422.4	0.0338	0.0002	0.2340	0.0072	0.0503	0.0015	0.0133	0.0003	0.3637	214.1	1.3	213.5	6.6	207.0	69.4
QLIII-1-27	22.0	647.2	0.0336	0.0002	0.2335	0.0058	0.0504	0.0012	0.0136	0.0002	0.2868	213.1	1.2	213.1	5.3	212.9	56.3
QLIII-1-28	12.2	357.2	0.0333	0.0002	0.2329	0.0118	0.0507	0.0025	0.0111	0.0003	0.3889	211.3	1.4	212.6	10.7	226.4	113.4
QLIII-1-29	17.2	476.4	0.0338	0.0002	0.2333	0.0078	0.0500	0.0016	0.0149	0.0003	0.4193	214.5	1.2	212.9	7.1	195.4	76.1
QLIII-1-30	21.5	626.4	0.0337	0.0002	0.2345	0.0068	0.0505	0.0014	0.0121	0.0002	0.3516	213.7	1.2	213.9	6.2	216.5	65.2
QLIII-2-1	43.6	1220.6	0.0335	0.0002	0.2317	0.0046	0.0502	0.0010	0.0100	0.0001	0.6026	212.3	1.1	211.6	4.2	204.7	45.6
QLIII-2-2	44.5	1213.6	0.0336	0.0002	0.2340	0.0040	0.0505	0.0009	0.0082	0.0001	0.8581	213.1	1.1	213.5	3.6	217.8	39.0
QLIII-2-3	66.0	1889.0	0.0338	0.0002	0.2348	0.0041	0.0503	0.0009	0.0095	0.0001	0.4984	214.5	1.1	214.2	3.7	210.6	39.3
QLIII-2-4	81.4	2317.1	0.0335	0.0002	0.2328	0.0032	0.0503	0.0007	0.0093	0.0000	0.5709	212.6	1.1	212.5	2.9	211.2	30.9
QLIII-2-5	60.8	1703.9	0.0333	0.0002	0.2315	0.0081	0.0505	0.0017	0.0071	0.0000	0.8699	211.0	1.1	211.4	7.4	215.9	79.8
QLIII-2-6	59.3	1670.2	0.0333	0.0002	0.2320	0.0034	0.0506	0.0007	0.0082	0.0000	0.7311	210.9	1.1	211.8	3.1	222.4	33.3
QLIII-2-7	40.6	1063.7	0.0334	0.0002	0.3098	0.0060	0.0673	0.0013	0.0101	0.0001	0.8286	211.7	1.1	274.0	5.3	846.9	38.9
QLIII-2-8	34.8	952.7	0.0335	0.0002	0.2348	0.0042	0.0508	0.0009	0.0109	0.0001	0.6314	212.6	1.2	214.2	3.8	232.0	41.8
QLIII-2-9	24.0	652.6	0.0333	0.0002	0.3430	0.0104	0.0746	0.0021	0.0125	0.0003	0.5316	211.4	1.1	299.4	9.0	1058.6	56.3
QLIII-2-10	46.5	1304.4	0.0336	0.0002	0.2339	0.0044	0.0506	0.0009	0.0104	0.0001	0.5666	212.8	1.1	213.4	4.0	220.4	41.5
QLIII-2-11	42.5	1071.5	0.0333	0.0002	0.4017	0.0148	0.0874	0.0028	0.0116	0.0003	0.8223	211.3	1.3	342.9	12.7	1370.3	61.6
QLIII-2-12	68.1	1950.4	0.0336	0.0002	0.2339	0.0030	0.0505	0.0006	0.0102	0.0000	0.4896	213.0	1.1	213.4	2.8	218.6	29.7
QLIII-2-13	47.4	1277.6	0.0334	0.0002	0.3130	0.0079	0.0680	0.0017	0.0101	0.0002	0.7106	211.5	1.1	276.5	7.0	869.8	50.3
QLIII-2-14	61.4	1692.9	0.0335	0.0002	0.2322	0.0070	0.0503	0.0015	0.0090	0.0001	0.7399	212.3	1.2	212.0	6.4	208.3	67.9
QLIII-2-15	40.2	1131.6	0.0339	0.0002	0.2352	0.0051	0.0502	0.0010	0.0112	0.0002	0.4780	215.2	1.1	214.4	4.6	206.0	47.0
QLIII-2-16	47.4	1185.4	0.0333	0.0002	0.4467	0.0131	0.0972	0.0027	0.0123	0.0003	0.7782	211.3	1.1	375.0	11.0	1572.0	52.0
QLIII-2-17	42.4	1160.7	0.0333	0.0002	0.3058	0.0064	0.0665	0.0014	0.0103	0.0001	0.6330	211.4	1.1	270.9	5.7	823.2	42.6
QLIII-2-18	47.3	1264.6	0.0334	0.0002	0.3745	0.0088	0.0814	0.0018	0.0105	0.0002	0.6719	211.5	1.1	323.0	7.6	1231.6	43.3
QLIII-2-19	38.6	1026.5	0.0336	0.0002	0.2325	0.0035	0.0502	0.0007	0.0104	0.0001	0.7750	212.8	1.1	212.3	3.2	206.5	34.1
QLIII-2-20	65.0	1763.3	0.0333	0.0002	0.2311	0.0047	0.0503	0.0008	0.0081	0.0001	0.9222	211.2	1.3	211.1	4.3	210.3	36.2
QLIII-2-21	30.1	852.2	0.0332	0.0002	0.2306	0.0048	0.0504	0.0010	0.0076	0.0000	0.7688	210.5	1.1	210.7	4.4	213.1	46.7
QLIII-2-22	36.5	957.5	0.0334	0.0002	0.3638	0.0101	0.0791	0.0021	0.0101	0.0002	0.7870	211.6	1.1	315.0	8.8	1174.2	52.0
QLIII-2-23	46.4	1251.4	0.0328	0.0002	0.2308	0.0049	0.0510	0.0011	0.0071	0.0000	1.1645	208.0	1.1	210.8	4.5	242.2	48.4
QLIII-2-24	54.3	1407.6	0.0334	0.0002	0.4036	0.0160	0.0877	0.0028	0.0096	0.0002	0.8521	211.6	1.3	344.3	13.7	1376.9	61.8
QLIII-2-25	48.4	1276.9	0.0333	0.0003	0.3960	0.0207	0.0862	0.0031	0.0103	0.0004	0.7355	211.2	1.6	338.7	17.7	1343.7	69.1
QLIII-2-26	30.1	769.7	0.0334	0.0002	0.4791	0.0224	0.1040	0.0040	0.0131	0.0006	0.6304	211.8	1.5	397.5	18.6	1697.5	71.3
QLIII-2-27	49.5	1263.0	0.0333	0.0002	0.4129	0.0194	0.0899	0.0034	0.0101	0.0003	0.8786	211.3	1.4	350.9	16.5	1422.4	71.8
QLIII-2-28	55.6	1443.4	0.0332	0.0002	0.2306	0.0083	0.0504	0.0018	0.0076	0.0000	1.2598	210.4	1.1	210.7	7.6	213.9	82.8
QLIII-2-29	73.6	267.5	0.2202	0.0013	7.1301	0.0895	0.2349	0.0030	0.0731	0.0014	0.6190	1282.8	7.5	2127.8	26.7	3085.5	20.5
QLIII-2-30	42.0	1123.2	0.0334	0.0003	0.3947	0.0238	0.0858	0.0036	0.0105	0.0005	0.6552	211.5	1.8	337.8	20.4	1334.3	82.1
QLKQ-2-1	120.9	227.7	0.0339	0.0001	0.2368	0.0016	0.0507	0.0003	0.0042	0.0002	0.7663	214.6	0.9	215.8	1.3	227.8	13.0
QLKQ-2-2	310.8	406.7	0.0337	0.0001	0.2344	0.0017	0.0504	0.0003	0.0033	0.0002	1.3756	213.8	0.9	213.9	1.4	216.7	8.3
QLKQ-2-3	155.8	369.8	0.0339	0.0002	0.2359	0.0020	0.0505	0.0003	0.0031	0.0002	0.7694	214.6	1.0	215.0	1.6	220.4	12.0
QLKQ-2-4	168.9	292.7	0.0338	0.0001	0.2355	0.0017	0.0506	0.0003	0.0032	0.0002	1.0025	214.2	0.8	214.8	1.4	220.4	13.0
QLKQ-2-5	357.1	459.9	0.0339	0.0001	0.2382	0.0012	0.0509	0.0002	0.0036	0.0002	1.1971	215.0	0.8	216.9	1.0	239.0	13.9
QLKQ-2-6	121.2	127.0	0.0572	0.0003	0.4251	0.0028	0.0540	0.0003	0.0060	0.0004	0.8862	358.3	1.6	359.7	2.0	368.6	8.3
QLKQ-2-7	230.5	341.2	0.0339	0.0001	0.2398	0.0014	0.0513	0.0002	0.0036	0.0002	1.0306	214.9	0.9	218.3	1.1	253.8	4.6
QLKQ-2-8	193.1	320.7	0.0340	0.0001	0.2486	0.0016	0.0531	0.0003	0.0039	0.0002	0.8331	215.5	0.9	225.5	1.3	331.5	13.0
QLKQ-2-9	156.2	363.0	0.0340	0.0001	0.2397	0.0015	0.0512	0.0002	0.0036	0.0002	0.6174	215.3	0.9	218.2	1.2	250.1	11.1
QLKQ-2-10	194.2	312.3	0.0349	0.0002	0.2434	0.0019	0.0506	0.0003	0.0036	0.0002	0.8714	220.8	1.2	221.2	1.6	233.4	13.0
QLKQ-2-11	95.7	243.3	0.0340	0.0002	0.2390	0.0019	0.0510	0.0003	0.0039	0.0003	0.5306	215.7	1.1	217.6	1.6	239.0	10.2
QLKQ-2-12	43.1	142.4	0.0338	0.0003	0.2367	0.0043	0.0508	0.0009	0.0053	0.0008	0.4598	214.3	1.8	215.7	3.5	231.6	40.7
QLKQ-2-13	139.6	321.6	0.0335	0.0001	0.2325	0.0016	0.0504	0.0003	0.0029	0.0003	0.7127	212.1	0.8	212.2	1.3	213.0	11.1
QLKQ-2-14	135.8	314.1	0.0336	0.0001	0.2336	0.0015	0.0505	0.0003	0.0029	0.0003	0.7356	213.0	0.9	213.2	1.2	216.7	8.3
QLKQ-2-15	387.1	492.1	0.0339	0.0002	0.2367	0.0016	0.0507	0.0002	0.0028	0.0003	1.3633	214.7	0.9	215.7	1.3	227.8	11.1
QLKQ-2-16	301.9	595.5	0.0340	0.0001	0.2367	0.0013	0.0505	0.0002	0.0026	0.0004	0.9106	215.5	0.8	215.7	1.1	216.7	12.0
QLKQ-2-17	40.3	231.0	0.0336	0.0002	0.2337	0.0025	0.0505	0.0005	0.0024	0.0005	0.4910	212.9	1.3	213.2	2.1	216.7	24.1
QLKQ-2-18	311.3	579.8	0.0340	0.0001	0.2378	0.0014	0.0507	0.0003	0.0023	0.0004	1.0586	215.8	0.8	216.6	1.2	233.4	13.0
QLIV-7-1	80.2	2257.5	0.0325	0.0002	0.2239	0.0034	0.0500	0.0008	0.0118	0.0001	0.5814	206.1	1.1	205.2	3.1	194.4	35.7
QLIV-7-2	84.0	2448.8	0.0324	0.0002	0.2221	0.0033	0.0498	0.0007	0.0100	0.0001	0.5592	205.3	1.1	203.7	3.1	184.8	34.6
QLIV-7-3	99.8	2903.5	0.0324	0.0002	0.2238	0.0031	0.0501	0.0006	0.0087	0.0000	0.6501	205.4	1.1	205.0	2.8	201.2	29.4
QLIV-7-4	79.7	2388.4	0.0325	0.0002	0.2228	0.0035	0.0496	0.0008	0.0072	0.0000	0.6060	206.5	1.1	204.2	3.2	178.5	35.6
QLIV-7-5	73.2	2243.6	0.0322	0.0002	0.2222	0.0040	0.0501	0.0009	0.0063	0.0000	0.6060	204.2	1.0	203.7	3.6	198.7	40.6
QLIV-7-6	98.2	3057.2	0.0320	0.0002	0.2225	0.0026	0.0505	0.0006	0.0054	0.0000	0.6438	203.0	1.1	204.0	2.4	215.9	26.8
QLIV-7-7	51.6	1628.4	0.0321	0.0002	0.2241	0.0042	0.0507	0.0009	0.0043	0.0000	0.6730	203.4	1.0	205.3	3.8	226.7	42.0
QLIV-7-8	47.5	1598.7	0.0310	0.0002	0.2139	0.0037	0.0501	0.0008	0.0035	0.0000	0.5140	196.5	1.0	196.9	3.4	200.8	39.0
QLIV-7-9	77.2	2452.8	0.0320	0.0002	0.2237	0.0053	0.0507	0.0012	0.0046	0.0001	0.5878	203.1	1.1	205.0	4.9	226.7	53.4
QLIV-7-10	83.7	2621.7	0.0322	0.0002	0.2243	0.0052	0.0506	0.0012	0.0054	0.0001	0.5697	204.1	1.1	205.5	4.8	221.5	53.7
QLIV-7-11	69.3	2043.7	0.0326	0.0002	0.2242	0.0056	0.0498	0.0012	0.0057	0.0001	0.8520	206.9	1.1	205.4	5.2	187.6	55.0
QLIV-7-12	73.5	2192.7	0.0325	0.0002	0.2244	0.0060	0.0501	0.0013	0.0072	0.0001	0.6289	206.3	1.1	205.6	5.5	197.6	59.1
QLIV-7-13	71.1	2123.5	0.0324	0.0002	0.2246	0.0036	0.0503	0.0008	0.0065	0.0001	0.7127	205.3	1.1	205.7	3.3	211.0	36.8
QLIV-7-14	60.3	1857.6	0.0321	0.0002	0.2224	0.0044	0.0502	0.0010	0.0077	0.0001	0.4829	203.8	1.1	203.9	4.0	204.4	45.0
QLIV-7-15	65.4	1923.7	0.0318	0.0002	0.2197	0.0049	0.0500	0.0010	0.0075	0.0001	0.7794	202.0	1.0	201.6	4.5	197.3	48.1
QLIV-7-16	64.0	1897.7	0.0321	0.0002	0.2232	0.0042	0.0504	0.0010	0.0083	0.0001	0.6197	203.9	1.0	204.6	3.9	212.7	43.9
QLIV-7-17	39.2	1174.2	0.0324	0.0002	0.2238	0.0033	0.0500	0.0007	0.0068	0.0000	0.6608	205.8	1.1	205.1	3.1	196.5	34.3
QLIV-7-18	78.9	2259.4	0.0324	0.0002	0.2251	0.0040	0.0504	0.0009	0.0065	0.0001	0.9644	205.4	1.0	206.2	3.6	215.6	39.9
QLIV-7-19	139.7	4211.6	0.0321	0.0002	0.2246	0.0033	0.0508	0.0007	0.0054	0.0000	0.8432	203.6	1.0	205.8	3.1	231.0	33.8
QLIV-7-20	66.4	1989.3	0.0322	0.0002	0.2228	0.0056	0.0502	0.0011	0.0050	0.0001	0.9459	204.2	1.1	204.2	5.2	204.6	52.0
QLIV-7-21	66.0	1990.4	0.0323	0.0002	0.2243	0.0038	0.0503	0.0008	0.0045	0.0000	0.9488	205.2	1.0	205.5	3.5	208.1	38.6
QLIV-7-22	63.3	1974.1	0.0323	0.0002	0.2235	0.0029	0.0502	0.0007	0.0037	0.0000	0.8506	204.7	1.1	204.8	2.6	206.2	30.0
QLIV-7-23	55.2	1769.7	0.0321	0.0002	0.2220	0.0044	0.0502	0.0010	0.0036	0.0001	0.6399	203.5	1.1	203.6	4.0	204.2	45.1
QLIV-7-24	86.5	2748.8	0.0322	0.0002	0.2246	0.0027	0.0505	0.0006	0.0025	0.0000	0.9992	204.5	1.1	205.7	2.5	219.5	28.2
QLIV-7-25	74.9	2092.7	0.0322	0.0002	0.2226	0.0044	0.0502	0.0010	0.0068	0.0001	1.0918	204.2	1.1	204.1	4.0	202.6	45.1
QLIV-7-26	60.7	1700.7	0.0323	0.0002	0.2245	0.0046	0.0503	0.0010	0.0127	0.0002	0.5672	205.2	1.0	205.6	4.2	210.9	46.3
QLIV-7-27	65.2	1895.2	0.0322	0.0002	0.2231	0.0040	0.0503	0.0009	0.0121	0.0002	0.4920	204.0	1.1	204.5	3.6	210.3	39.5
QLIV-7-28	59.6	1578.2	0.0323	0.0002	0.2229	0.0076	0.0500	0.0015	0.0194	0.0007	0.4942	205.0	1.1	204.3	7.0	196.6	69.2
QLIV-7-29	80.5	1886.7	0.0325	0.0002	0.2234	0.0062	0.0499	0.0012	0.0196	0.0004	0.7680	206.0	1.1	204.8	5.7	190.9	57.2
QLIV-7-30	65.1	1748.6	0.0322	0.0002	0.2219	0.0046	0.0500	0.0010	0.0216	0.0002	0.4211	204.4	1.0	203.5	4.2	192.8	47.1

and

Table 3. Lu–Hf Isotope contents of samples from the Qieliekeqi deposit.

Sample Number	t (Ma)	176Yb/177Hf(corr)	176Lu/177Hf(corr)	176Hf/177Hf(corr)	176Hf/177Hf	2σm	eHf(0)	eHf(t)	2σ	TDM1(Hf)	TDM2(Hf)	fLu/Hf
QLIII-1.1	212	0.0438	0.0010	0.2825	0.2825	0.00002	−10.5465	−6.0390	0.8098	1102	2222	−0.9697
QLIII-1.2	212	0.0455	0.0014	0.2825	0.2825	0.00003	−8.7630	−4.3032	1.1686	1040	2066	−0.9593
QLIII-1.3	215	0.1303	0.0030	0.2825	0.2825	0.00004	−8.9125	−4.6207	1.2722	1094	2096	−0.9103
QLIII-1.4	215	0.0856	0.0020	0.2826	0.2825	0.00003	−8.4374	−4.0062	1.1184	1045	2041	−0.9397
QLIII-1.5	215	0.0307	0.0007	0.2825	0.2825	0.00003	−10.1774	−5.5673	1.0911	1079	2182	−0.9779
QLIII-1.6	216	0.0543	0.0012	0.2825	0.2825	0.00003	−9.6285	−5.0699	1.0332	1072	2138	−0.9624
QLIII-1.7	214	0.0613	0.0014	0.2825	0.2824	0.00003	−11.4843	−6.9858	1.0027	1150	2308	−0.9588
QLIII-1.8	213	0.0368	0.0010	0.2826	0.2826	0.00002	−6.9251	−2.3941	0.7719	958	1895	−0.9698
QLIII-1.9	213	0.0429	0.0010	0.2826	0.2826	0.00002	−7.3147	−2.7772	0.8466	972	1930	−0.9712
QLIII-1.10	214	0.0439	0.0012	0.2826	0.2826	0.00002	−6.2498	−1.7216	0.6746	935	1835	−0.9646
QLIII-1.11	212	0.0482	0.0016	0.2825	0.2825	0.00002	−8.6944	−4.2670	0.6899	1044	2063	−0.9523
QLIII-1.12	213	0.0562	0.0017	0.2826	0.2826	0.00002	−7.5630	−3.1349	0.7275	1002	1962	−0.9478
QLIII-1.13	212	0.0660	0.0019	0.2826	0.2826	0.00002	−7.4717	−3.0935	0.7261	1004	1957	−0.9416
QLIII-1.14	213	0.0409	0.0010	0.2826	0.2825	0.00002	−8.3082	−3.7809	0.6309	1013	2020	−0.9691
QLIII-1.15	213	0.0571	0.0014	0.2826	0.2826	0.00002	−6.3234	−1.8412	0.7341	942	1845	−0.9593

. The corresponding Concordia diagrams of zircon U–Pb ages are shown in Figure 6.

Sample QLIII-1 is a gneissic biotite granite. The zircon crystals are prismatic, ranging from 150 to 250 μm in length, with aspect ratios mostly around 2:1 and occasionally up to 3:1. Cathodoluminescence (CL) images show well-developed oscillatory zoning typical of magmatic zircons, with rare occurrences of core–rim structures (Figure 7a). Of the 30 analytical spots, 29 yielded 206Pb/238U ages ranging from 210.87 to 215.49 Ma. The weighted mean age is 213.04 ± 0.45 Ma (n = 29, MSWD = 0.68), and the average Th/U ratio is 0.34, consistent with a magmatic origin [26,27]. One inherited zircon grain yielded a significantly older age of 434.06 Ma, with a Th/U ratio of 0.63, representing a xenocrystic component derived from older crustal material. 15 Lu–Hf isotope analyses were conducted on the same sample. The εHf(t) values range from −6.99 to −1.72, with an average of −3.94. The corresponding single-stage model ages (TDM1) range from 935 to 1150 Ma (average 1030 Ma), and the two-stage model ages (TDM2) range from 1835 to 2308 Ma (average 2037 Ma). These results suggest that the magma was likely derived from partial melting of late Paleoproterozoic to Mesoproterozoic crustal material.

Sample QLIII-2 is also a gneissic biotite granite. The zircon grains are prismatic, with grain sizes ranging from 80 to 200 μm and aspect ratios generally around 2:1. CL images show uneven internal luminescence, suggesting a relatively complex crystallization history (Figure 7b). 30 analytical spots were selected for U–Pb isotopic analysis. Among them, 29 spots yielded 206Pb/238U ages between 208.05 and 215.22 Ma, with a weighted mean age of 211.83 ± 0.52 Ma (n = 29, MSWD = 0.82). The Th/U ratios range from 0.48 to 1.26, averaging 0.74, which is characteristic of magmatic zircon. One spot yielded a significantly older age of 1282.81 Ma, with a Th/U ratio of 0.62, and is interpreted as an inherited zircon derived from older crustal material.

Sample QLKQ-2 is a banded plagioclase amphibolite. The zircon grains are prismatic, ranging from approximately 100 to 250 μm in length, with aspect ratios around 2:1. The CL images show weak luminescence, indicating low internal zonation intensity (Figure 7c). A total of 18 spots were analyzed for U–Pb isotopes, and one spot was excluded due to significant discordance. The remaining 16 spots yielded 206Pb/238U ages ranging from 212.12 to 215.82 Ma. The weighted mean age is 214.54 ± 0.59 Ma (n = 16, MSWD = 0.75). Th/U ratios range from 0.46 to 1.38, with an average of 0.87, consistent with a magmatic origin. One additional spot yielded an older age of 358.31 Ma, with a Th/U ratio of 0.89, and is interpreted as an inherited zircon derived from pre-existing crustal material.

Sample QLIV-7 is a biotite–quartz gneiss. The zircon grains are predominantly prismatic, with sizes ranging from 100 to 250 μm and aspect ratios mostly greater than 2:1. CL images reveal heterogeneous luminescence within the grains (Figure 7d). 30 spots were analyzed for U–Pb isotopes, and one discordant spot was excluded from the dataset. The remaining 29 spots yielded 206Pb/238U ages ranging from 201.99 to 206.93 Ma, with a weighted mean age of 204.63 ± 0.44 Ma (n = 29, MSWD = 0.65). The Th/U ratios range from 0.42 to 1.09, with an average of 0.70, indicating a magmatic origin. This age represents a relatively younger magmatic intrusion event in the study area.

Figure 6. Concordia diagram of zircon U–Pb ages for intrusive rocks from the Qieliekeqi deposit.

Figure 6. Concordia diagram of zircon U–Pb ages for intrusive rocks from the Qieliekeqi deposit.

Figure 7. Representative cathodoluminescence (CL) images of zircons from the intrusive rocks and their U–Pb ages (Ma) from the Qieliekeqi deoposit. (a) QLIII-1, (b) QLIII-2, (c) QLKQ-2, (d) QLIV-7. White circles indicate LA-ICP-MS analytical spots for U–Pb–Hf isotope analyses.

Figure 7. Representative cathodoluminescence (CL) images of zircons from the intrusive rocks and their U–Pb ages (Ma) from the Qieliekeqi deoposit. (a) QLIII-1, (b) QLIII-2, (c) QLKQ-2, (d) QLIV-7. White circles indicate LA-ICP-MS analytical spots for U–Pb–Hf isotope analyses.

5.3. C–O Isotopes

Carbon and oxygen isotope analyses were performed on four siderite ore samples, including three coarse-grained samples (QLIII-8, QLIV-2, QLIV-5) and one representative banded sample (QLKQ-8). As a carbonate mineral, siderite preserves δ¹³C and δ¹⁸O signatures that can constrain the CO₂ source in the ore-forming fluids and shed light on the evolution of the hydrothermal system. The analytical results are presented in

Table 4. C-O isotope contents of samples from the Qieliekeqi deposit.

Sample Number	QLIII-8	QLIV-2	QLIV-5	QLKQ-8
Type	Coarse-Grained	Coarse-Grained	Coarse-Grained	Banded
δ13CPDB(‰)	−6.0	−5.3	−4.6	−5.2
δ18OPDB(‰)	−14.1	−14.4	−13.6	−14.2
δ18OSMOW(‰)	16.4	16.0	16.9	16.2

.

The analytical results show that the δ¹³C_PDB values of the four siderite samples fall within a narrow range of −6.0‰ to −4.6‰, with an average of −5.3‰. The δ¹⁸O_PDB values are also relatively consistent, ranging from −13.6‰ to −14.4‰, with an average of −14.1‰. When converted to the SMOW scale, the δ¹⁸O values range from 16.0‰ to 16.9‰, with an average of 16.4‰.

6. Discussion

6.1. Timing of Mineralization and Magmatic Activity

The age of the ore-hosting strata in the Qieliekeqi siderite deposit has long been a subject of debate, primarily due to the lack of direct geochronological constraints on the marine carbonate sequence in which the deposit is hosted. Earlier studies attempted to infer the stratigraphic age based on regional tectonic context or indirect evidence, resulting in varying interpretations and classifications [9,10,11,12,13].

In this study, zircon U–Pb and Lu–Hf isotopic analyses were performed on intrusive rocks from the eastern margin of the Qiukutai pluton, which crosscuts the ore-bearing strata, to constrain the timing and source characteristics of magmatism potentially associated with siderite mineralization. Among the four samples dated by U–Pb, sample QLIII-1 was also analyzed for Lu–Hf isotopes. It is lithologically representative and shares similar petrographic and field features with the other intrusions, making it a suitable proxy for characterizing the magmatic source. All four zircon samples exhibit features consistent with a magmatic origin, including oscillatory zoning and high Th/U ratios, as previously described. These results confirm that the intrusive rocks crystallized during the Late Triassic Indosinian period, marking a significant regional magmatic event. Although this magmatism postdates the primary siderite mineralization, it may have triggered late-stage hydrothermal overprinting and local remobilization. The Lu–Hf isotopic signatures, characterized by negative εHf(t) values and Paleoproterozoic model ages, suggest derivation from partial melting of ancient lower crustal material. These geochemical features suggest that Indosinian magmatism may have played a role in the late-stage thermal or hydrothermal evolution of the Qieliekeqi deposit.

Inherited zircon ages were also identified in three of the analyzed samples: QLIII-1, QLIII-2, and QLKQ-2, yielding ages of 434.06 Ma, 1282.81 Ma, and 358.31 Ma, respectively. These inherited zircons exhibit well-defined prismatic morphologies, core–rim structures, and clear oscillatory zoning under CL imaging (Figure 7), all indicative of a magmatic origin. The 1282.81 Ma age corresponds to the timing of the breakup of the Columbia supercontinent during the Mesoproterozoic, suggesting that the Tashkurgan block, where the Qieliekeqi deposit is located, may have been involved in global-scale lithospheric extension and rifting during this period. The 434.06 Ma and 358.31 Ma ages are likely related to the closure and subduction processes of the Proto-Tethys Ocean in the region. Recent zircon U–Pb dating conducted by Zhu et al. [22] on siderite from the Qieliekeqi deposit yielded a crystallization age of 488 Ma and a later metamorphic overprint age of 434 Ma. The 434 Ma metamorphic age coincides with the inherited zircon age found in sample QLIII-1 in this study, indicating that a regional magmatic or hydrothermal event occurred during the Early Silurian (~434 Ma). Meanwhile, the 488 Ma zircon crystallization age provides direct geochronological evidence that the primary formation of siderite at Qieliekeqi occurred during the Late Cambrian.

Based on the available zircon geochronology and Lu–Hf isotopic data, a tentative model of crustal evolution in the study area may be proposed as follows: During the Paleoproterozoic (~2037 Ma), the crust likely formed through differentiation from a depleted mantle source. Between the Mesoproterozoic and Neoproterozoic (~1030 Ma), the crust experienced modification and recycling. By the Late Cambrian (~500–488 Ma), a marine sedimentary environment prevailed, during which the initial siderite-bearing strata were deposited. In the Late Triassic (~221–204 Ma), under a tectonic setting associated with the subduction of the Tethys Ocean, the crust underwent partial melting again, generating granitic magmas that intruded into the ore-bearing strata. These intrusions were likely accompanied by hydrothermal fluids during cooling, which may have locally remobilized pre-existing siderite mineralization. While this model is based on limited isotopic data, it provides a useful first-order framework for interpreting the crustal and metallogenic evolution of the Qieliekeqi deposit.

6.2. Ore-Forming Environment

The siderite samples are generally characterized by low TiO₂ and Al₂O₃ contents. Compared to the banded siderite, the coarse-grained samples show lower average TiO₂ (0.07% vs. 0.15%) and Al₂O₃ (1.86% vs. 2.56%) contents, indicating differences in the input of terrigenous detritus during deposition. Although the banded siderite samples contain higher absolute concentrations of both Al₂O₃ and TiO₂, their Al₂O₃/TiO₂ ratios (average 17.07) are lower than those of the coarse-grained samples (average 29.14). This discrepancy suggests that the high Al₂O₃/TiO₂ ratios in the coarse-grained ores may reflect stronger Ti depletion rather than increased terrigenous input. This interpretation is further supported by the behavior of other major elements. The coarse-grained samples exhibit significantly higher SiO₂/TiO₂ (average 260.5) and K₂O/TiO₂ (average 9.38) ratios than the banded ores (84.7 and 5.49, respectively). However, the absolute SiO₂ contents are comparable between the two groups (13.36 wt% in coarse-grained vs. 12.58 wt% in banded), and the average K₂O content is only slightly higher in the coarse-grained samples (0.61 wt% vs. 0.82 wt%). These data indicate that the elevated ratios are primarily the result of lower TiO₂ concentrations in the coarse-grained ores rather than a substantial increase in detrital components such as quartz, feldspar, or clay minerals. Therefore, the systematic enrichment in element-to-Ti ratios is best explained by selective Ti leaching during diagenesis or hydrothermal alteration rather than enhanced terrigenous input. The extremely low TiO₂ contents in the coarse-grained siderite samples further support a predominant hydrothermal sedimentary origin, where Ti was largely excluded due to its low solubility and immobility in hydrothermal fluids. In contrast, the relatively higher TiO₂ and Al₂O₃ contents and lower element-to-Ti ratios in the banded siderite ores may reflect a greater contribution of volcanic or clastic detritus, which introduced additional Ti- and Al-bearing phases into the depositional system [28,29]. This suggests that the banded and coarse-grained siderite types likely represent different degrees of hydrothermal versus detrital influence during ore formation. In addition, the Al/(Al + Fe + Mn) ratio is commonly used as an indicator of the extent of hydrothermal fluid input [30,31]. Typical seafloor hydrothermal sediments have ratios around 0.01, while the average for shales is approximately 0.62. In the Qieliekeqi deposit, the Al/(Al + Fe + Mn) ratios of coarse-grained and banded siderite samples are 0.02 and 0.03, respectively. These relatively low values suggest a significant influence from hydrothermal fluids in both ore types. However, since the values are close to the threshold distinguishing hydrothermal sedimentation from hydrothermal replacement, this index alone may not fully resolve the genetic mechanism. Further mineralogical evidence would be needed to confidently distinguish between these two processes. Based on the differential chemical mobility of Al, Fe, Mn, Na, and Mg in various depositional environments, Nicholson proposed the use of Al–Fe–Mn ternary diagrams and Na–Mg binary plots for genetic discrimination [8,32]. As shown in Figure 8a, both siderite types from the study area plot within the hydrothermal field, consistent with hydrothermal sedimentary siderite from the Bayan Obo region [33], but clearly distinct from lacustrine terrigenous siderite from the Anya area in the Ordos Basin [34]. This further supports a hydrothermal origin for the Qieliekeqi siderite. The Na–Mg binary diagram (Figure 8b) reveals differences in the depositional water depth between the two ore types. While most samples formed in a shallow marine setting, the coarse-grained siderite appears to have precipitated in relatively deeper waters. Additionally, the MnO/TiO₂ ratio in the coarse-grained ores (average 29.52) is notably higher than that in the banded ores (average 11.75). This difference may suggest variations in depositional conditions, possibly linked to water depth, because Mn tends to accumulate under reducing conditions that are more typical of deeper marine environments. However, Ti is primarily derived from terrigenous input and remains largely immobile under low-temperature hydrothermal conditions. This introduces uncertainty when using the MnO/TiO₂ ratio as a direct proxy for water depth. Therefore, this ratio is considered a qualitative indicator rather than a definitive one. In addition, the coarse-grained siderite ores typically occur in stratigraphically deeper layers, which may provide further support for a relatively deeper and more reducing depositional environment compared to the banded ores. Collectively, these geochemical indicators point to a seafloor depositional setting with substantial hydrothermal fluid input.

The enrichment factor (EF) is a key geochemical indicator used to characterize the enrichment behavior of elements in ore-forming fluids and to trace their material sources [36]. It provides insights into the metal-carrying capacity of hydrothermal fluids and the characteristics of the depositional environment. Enrichment factors (EF) were calculated by normalizing the trace element concentrations to the global average composition of marine carbonate rocks, as reported by Tribovillard et al. [37] (Figure 9), Mn (EF = 11.1 to 13.1) and Co (with a maximum EF of 97) are significantly enriched. This enrichment indicates that the ore-forming fluids were highly reducing and operated under low to moderate temperatures, which is consistent with the formation of Fe-carbonates and disseminated pyrite observed in the ore. The high EF value of Co may reflect focused hydrothermal inputs and/or a lithological control by mafic to intermediate source rocks, which are common in the regional tectonomagmatic framework [38,39]. Ba (EF = 1.7 to 10.8) and U (EF = 0.6 to 2.3) are also enriched. This suggests that hydrothermal fluids were capable of transporting Ba and U as sulfate- or carbonate-complexed species. The presence of barite and siderite in the ores supports this interpretation [40,41]. Although U is generally immobile under reducing conditions, its enrichment implies the presence of redox interfaces or partial re-oxidation zones where U⁶⁺ species could form complexes and precipitate. These interpretations are reinforced by the occurrence of pyrite, which is indicative of a reducing environment [42,43]. In contrast, Zn (EF = 0.10 to 0.20) and Sr (EF = 0.01 to 0.06) are consistently depleted across all samples, indicating a generally low metal load in the ore-forming fluid and the absence of typical sulfide mineralization characteristics. The marked depletion of Sr may be attributed to early-stage precipitation of carbonate minerals, particularly siderite, which can incorporate Sr²⁺ into its crystal lattice due to the similar ionic radius and charge between Sr²⁺ and Ca²⁺/Fe²⁺ [44]. In addition, prolonged water–rock interaction may have further reduced Sr availability by facilitating adsorption onto clay minerals or promoting its removal via fluid–mineral equilibria under reducing conditions [45]. Overall, the deposit exhibits the geochemical signature of a classic low- to moderate-temperature, strongly reducing hydrothermal sedimentary system, with minor input of terrestrial detritus and clear evidence of selective elemental migration and precipitation during mineralization.

The Fe–Mn–(Cu+Co+Ni) ternary diagram [46,47] further classifies the samples from the study area within the hydrothermal field (Figure 10a), consistent with marine hydrothermal siderite deposits such as those at Bayan Obo [33]. In addition, the Cr–Zr binary plot [48] supports this interpretation, with all sample points falling near the trend line associated with modern hydrothermal sedimentation (Figure 10b).

The PAAS-normalized rare earth element (REE) patterns of the siderite samples from the study area exhibit distinct features, including low total REE concentrations (ΣREE), depletion in light REEs (LREEs), enrichment in heavy REEs (HREEs), and pronounced positive Eu anomalies. These features are markedly different from high-temperature hydrothermal precipitates such as mid-ocean ridge Fe-carbonates, which typically show ΣREE > 100 ppm with strong LREE enrichment and no significant HREE dominance. Instead, the observed patterns resemble those of low- to moderate-temperature seafloor hydrothermal sediments formed under oxidized bottom water conditions [36,49,50,51]. The enrichment in HREEs may be partly attributed to ligand complexation effects. In seawater, carbonate and sulfate ligands preferentially form more stable complexes with HREEs, enhancing their mobility and incorporation into minerals during fluid mixing [39]. The positive Eu anomalies, while indicative of high-temperature reducing fluid input, are also influenced by other factors such as the host mineral phase and fluid residence time. In particular, Eu²⁺ is more readily incorporated into carbonates than into silicates due to favorable ionic radius matching and charge balance [52]. A comparison of Eu/Eu* values between banded and coarse-grained samples suggests subtle differences in redox conditions or fluid-rock interaction intensity, with higher Eu/Eu* in the latter supporting proximity to hydrothermal sources. The pronounced negative Ce anomalies in both coarse-grained and banded siderite samples (δCe = 0.83 and 0.84, respectively) further indicate formation under conditions influenced by oxidizing bottom waters, consistent with a hydrothermal–seawater interaction environment [48]. The coarse-grained ores may have formed closer to hydrothermal vents where mixing with oxidized seawater was limited, while the banded ores likely experienced more extensive seawater dilution. Notably, sample QLIII-11 displays an anomalously high ΣREE value (313.03 ppm), with strong LREE enrichment and an intensified positive Eu anomaly. This suggests that the sample was deposited near a hydrothermal vent center, where hydrothermal fluid contributions were dominant and seawater dilution was minimal. Its REE pattern closely matches that of typical hydrothermal carbonates [48].

Overall, the major and trace element ratios, together with the REE distribution patterns of the siderite ores in the study area, suggest that their formation was controlled by seafloor hydrothermal activity combined with mixing with oxidized seawater. This depositional model is consistent with modern hydrothermal carbonate systems. The coarse-grained and banded siderite ores display distinct geochemical compositions. These differences most likely reflect variations in terrigenous input or sedimentary environment rather than fundamental differences in the fluid source. In support of this interpretation, the δ¹³C and δ¹⁸O values of the two ore types are broadly similar, indicating precipitation from a common hydrothermal fluid. The geochemical and isotopic evidence together point to a unified ore-forming system influenced by localized depositional conditions.

6.3. Source and Evolution of Ore-Forming Fluids

The oxygen isotope composition of carbonate rocks is relatively sensitive to alteration. When there is no significant correlation between δ¹³C_PDB and δ¹⁸O_SMOW values, it typically indicates that the carbonate has preserved its original isotopic signature [53]. In this study, the δ¹³C_PDB and δ¹⁸O_SMOW values of siderite samples show no evident correlation. In addition, the average Mn/Sr ratio is 0.13, which is significantly lower than the alteration threshold of Mn/Sr = 10 proposed by Kaufman et al. [54]. Although the direct application of this Mn/Sr threshold to hydrothermal siderite warrants caution due to the influence of fluid composition, pH, and redox state [55], the following lines of evidence support the preservation of primary isotopic signals: First, the PAAS-normalized REE patterns retain systematic HREE enrichment and distinct Eu anomalies without flattening or LREE enrichment that would indicate diagenetic redistribution [50,51]. Second, major element ratios fall within the typical ranges of hydrothermal carbonate precipitation [48]. Based on these integrated lines of evidence, we infer that the δ¹³C and δ¹⁸O values in the present study largely reflect primary fluid compositions during siderite formation. Because the two types of siderite ores show similar isotopic ranges and means, they are discussed together in the following sections.

The stable isotopic compositions of carbon and oxygen in crustal fluids are widely used to trace the sources and evolutionary processes of ore-forming fluids. Previous studies have shown that the carbon and oxygen in such fluids may originate from multiple sources, including volatile components released from sedimentary organic matter, thermal decomposition and recycling of marine carbonates, and degassing of magmatic volatiles enriched in light isotopes [29,56,57,58,59]. These fluid sources exhibit distinct δ¹³C_PDB and δ¹⁸O_SMOW signatures. For example, magmatic fluids typically show δ¹³C_PDB values between −10‰ and +4‰, whereas marine carbonate-derived fluids tend to have more positive values. Fluids derived from sedimentary organic matter are significantly more depleted in ¹³C, with δ¹³C_PDB values ranging from −30‰ to −15‰. In terms of oxygen isotopes, magmatic fluids generally have δ¹⁸O_SMOW values between 6‰ and 12‰, while fluids that have undergone extensive water–rock interaction in skarn systems can reach values of 16‰ or higher [60,61,62,63].

The carbon and oxygen isotope results for siderite samples from the Qieliekeqi deposit show δ¹³C_PDB values ranging from −6.0‰ to −4.6‰ and δ¹⁸O_SMOW values between 16.0‰ and 16.9‰. These values do not correspond exactly to those of typical sedimentary carbonates (δ¹³C_PDB > 0‰, δ¹⁸O_SMOW < 20‰), organic matter-derived hydrothermal fluids (δ¹³C_PDB < −20‰), or magmatic fluids (δ¹⁸O_SMOW = 5.5‰ to 10‰) [63,64,65], but fall within the broader range of mixed or evolved fluid systems. The relatively negative δ¹³C values may reflect the input of deep-seated inorganic carbon sources or carbon-bearing sedimentary components, while the high δ¹⁸O values are consistent with isotope exchange between hydrothermal fluids and host rocks during prolonged fluid–rock interaction under low-temperature conditions. Taken together, the isotopic data suggest that the ore-forming fluids likely evolved through a combination of magmatic, sedimentary, and possibly seawater-derived components. A binary or ternary fluid mixing model, or an isotopic evolution pathway resulting from open-system water–rock interaction, may better explain the observed values.

On the δ¹³C_PDB vs. δ¹⁸O_SMOW plot (Figure 11a), the Qieliekeqi samples plot between the magmatic and marine carbonate fields near the fluid mixing trend line. Their positions are comparable to hydrothermal sedimentary siderite from the Rudňany deposit in Slovakia and the CYZ-XXJ (C-X) deposit in Guizhou, China [66,67], but are clearly different from siderite formed in brackish water environments such as at Yujingshan, Yunnan [29], and freshwater sedimentary siderite from the Huopu area, Guizhou [28]. This pattern is also evident in the fluid source discrimination diagram (Figure 11b), where the Qieliekeqi samples plot within the hydrothermal–skarn field, showing a concentrated distribution. These isotopic characteristics collectively suggest that the formation of siderite in the study area involved significant water–rock interaction with hydrothermal fluids.

To quantitatively estimate the carbon source composition of the Qieliekeqi siderite, we applied the Bayesian stable isotope mixing model MixSIAR [68] using δ¹³C_PDB and δ¹⁸O_SMOW data from the siderite samples. The model incorporated three potential carbon end-members with the following isotopic values: (1) magmatic CO₂ (δ¹³C_PDB = −6.0 ± 0.3‰, δ¹⁸O_SMOW = 8.5 ± 0.5‰) [63,69]; (2) sedimentary organic matter (δ¹³C_PDB = −25.0 ± 1.0‰, δ¹⁸O_SMOW = 16.5 ± 0.6‰) [61,70]; and (3) marine carbonate (δ¹³C_PDB = +5.15 ± 0.4‰, δ¹⁸O_SMOW = 18.0 ± 0.4‰) [71,72]. Theδ¹³C_PDB value for the carbonate end-member was adjusted by +3.15‰ to account for the SPICE (Steptoean Positive Carbon Isotope Excursion) event during the Cambrian period [73]. Since organic matter generally lacks structural oxygen, its δ¹⁸O value cannot be directly measured. However, for model input purposes, the background δ¹⁸O value of hydrothermal fluids in water–rock systems were used in the end-member (2). In addition, a single-isotope model using only δ¹³C_PDB was also run for comparison. The model was based on the joint posterior distribution of δ¹³C_PDB and δ¹⁸O_SMOW and conducted using 1000 simulations to estimate the relative contributions of each carbon source. The results (

Table 5. Comparison of posterior distributions of joint δ¹³C_PDB-δ¹⁸O_SMOW model vs. δ¹³C_PDB-only model.

Source Endmenber	Joint Model (δ13C + δ18O)		δ13C-only Model
	Median	Credible Interval	Median	Credible Interval
Magmatic CO2	19.70%	8.6%–33.4%	16.50%	5.6%–32.4%
Organic Carbon	30.60%	18.6%–45.1%	29.30%	15.7%–45.6%
Marine Carbonate	49.70%	33.4%–63.6%	53.70%	32.5%–70.1%

) show that marine carbonate was the predominant carbon source, with a median contribution of 49.7% (credible interval: 33.4% to 63.6%), followed by sedimentary organic carbon at approximately 30.6% (18.6% to 45.1%), and magmatic CO₂ at 19.7% (8.6% to 33.4%) (Figure 12a). In contrast, the single-isotope model (δ¹³C-only) yielded similar estimates but with significantly wider credible intervals (Figure 12b). These results reflect a mixed carbon system in the hydrothermal environment involving deep-sourced magmatic gases, shallow-derived thermally degraded organic matter, and seawater-derived carbonate. The model also provides a quantitative basis for reconstructing carbon cycling in Cambrian oceanic crustal settings.

The δ¹⁸O_PDB values of carbonate minerals are commonly used to reconstruct paleotemperatures during mineral deposition [28,29,74]. In this study, the oxygen isotope paleotemperature was calculated using the equation proposed by Shackleton and Kennett [75]:

T = 16.9 − 4.38(δ_C − δ_W) + 0.10(δ_C − δ_W)²

(1)

where δ_C is the δ¹⁸O_PDB value of the siderite sample, and δ_W is the δ¹⁸O_SMOW value of coeval seawater. A Late Cambrian global average seawater δ¹⁸O_SMOW value of –6.5‰ was used in the calculation [76,77]. The results show that the precipitation temperatures of siderite in the study area range from 53.03 °C to 57.74 °C, with an average of 55.82 °C. Although reported precipitation temperatures of hydrothermal siderite cover a wide range (13.5 °C to 250 °C) [29,55,57,78,79,80,81,82], many studies have shown that low-temperature hydrothermal siderite commonly forms below 100 °C. For example, clumped isotope analyses of siderite in the Gunflint Iron Formation suggest formation temperatures between 40 °C and 132 °C, with common values in the 50–60 °C range under microbially influenced, early diagenetic conditions [55]. In southwestern China, siderite associated with tuffaceous hydrothermal deposits shows precipitation temperatures between 19.47 °C and 47.12 °C, reflecting interaction between hydrothermal fluids and volcanic–sedimentary material [29]. Additionally, although not formed in a hydrothermal setting, siderite precipitated at ≤60 °C has been documented in stratified, reducing lacustrine environments [57], further demonstrating the plausibility of siderite formation at similar low temperatures under suitable redox and geochemical conditions. Therefore, the measured temperatures in this study fall well within the expected range for low-temperature hydrothermal systems and support a hydrothermal sedimentary origin for the deposit.

To further reconstruct the oxygen isotopic composition of the ore-forming fluid (δ¹⁸O_fluid), we applied the empirical fractionation equation for siderite–water systems proposed by Zheng [83]:

1000 × ln α = (4.23 × 10⁶)/T² − (4.58 × 10³)/T + 1.73

(2)

where T is the precipitation temperature in Kelvin, and α is the fractionation factor between siderite and water. According to classical isotope fractionation theory [48,56], the fractionation between a mineral and a fluid can be expressed as

Δ¹⁸O = δ¹⁸O_mineral − δ¹⁸O_fluid ≈ 10³ ln α

(3)

Thus, the δ¹⁸O value of the fluid can be calculated as

δ¹⁸O_fluid = δ¹⁸O_minera − 10³ ln α

(4)

Using the average precipitation temperature of 55.82 °C (328.97 K) determined in this study, the calculated fractionation factor is approximately α ≈ 1.028, corresponding to a fractionation of Δ¹⁸O ≈ +27.0‰. Based on the measured δ¹⁸O_SMOW values of siderite samples (16.0‰ to 16.9‰), the inferred δ¹⁸O_fluid values range from −10.89‰ to −9.99‰, with an average of −10.52‰. This value is slightly lower than that of typical Cambrian seawater, suggesting that the ore-forming fluid was not unmodified shallow seawater. Instead, it may have experienced deep circulation and isotopic exchange with surrounding rocks in an open system. In the study area, potential rock types contributing to this exchange include Neoproterozoic to Cambrian sedimentary units such as mudstone, shale, and dolostone, which are widespread in the regional stratigraphy [5,16,22]. These rocks are known to have relatively high δ¹⁸O values and can significantly modify the isotopic composition of infiltrating fluids during water–rock interaction. Alternatively, isotopic dilution may have also occurred through mixing with early meteoric water or formation water. These processes are consistent with a shallow marine to continental shelf depositional environment.

The carbon and oxygen isotope data, together with the modeling results, indicate that the ore-forming fluid at Qieliekeqi was derived from a mixed system. Marine carbonate was the predominant carbon source, with additional input from sedimentary organic matter and magmatic CO₂, reflecting a coupling between shallow depositional settings and deep magmatic degassing. The reconstructed δ¹⁸O_fluid values suggest that the mineralizing fluid formed through low-temperature mixing of meteoric water, seawater, and deep hydrothermal fluids, followed by extensive water–rock interaction in an open system. This highlights a hydrothermal system driven by deep magmatic heat and modified by surficial inputs, providing strong evidence for the coupling between carbon cycling and mineralization in an Early Paleozoic tectonic context.

6.4. Metallogenic Model

The metallogenic process of the Qieliekeqi siderite deposit is characterized by multi-stage evolution and the coupling of multiple fluid sources. It can be broadly divided into two main stages: an early phase of hydrothermal sedimentation and a later stage of magmatic–hydrothermal overprinting.

Zircon Lu–Hf isotopic data from intrusive rocks in the Qieliekeqi area reveal Paleoproterozoic single-stage model ages averaging ~2037 Ma and Mesoproterozoic–Neoproterozoic two-stage ages around ~1030 Ma, indicating the presence of an evolved, multi-stage continental basement. Although this Precambrian crust predates ore formation by over a billion years, its long-lived evolution and compositional heterogeneity may have influenced the localization of syn-depositional faults and fluid migration during Cambrian extension. Inherited basement structures could have acted as preferential pathways for ascending Fe-rich hydrothermal fluids, shaping the spatial configuration of the mineralization system. This tectonic framework is supported by isotopic evidence and provides a plausible link between crustal inheritance and early-stage ore formation.

In the Late Cambrian of the Early Paleozoic (~500 to 488 Ma), the study area was situated in an initial rifting and spreading setting of the Proto-Tethys Ocean. Deep lithospheric thermal activity facilitated the rapid ascent of Fe-rich hydrothermal fluids along fault systems. These fluids are mixed with shallow, oxidized seawater, triggering the precipitation of siderite in a shallow marine to continental shelf environment. The resulting mineralization primarily formed coarse-grained and banded siderite (Figure 13a), reflecting rapid precipitation under relatively deep and shallow water conditions, respectively. Both ore types also show evidence of nearshore input of terrigenous detritus, as indicated by their major element geochemical characteristics. Elevated contents and characteristic ratios of Al₂O₃, SiO₂, and K₂O, in particular the high values of Al₂O₃/TiO₂ and SiO₂/TiO₂, are commonly recognized as reliable geochemical proxies for detrital silicate input in sedimentary and hydrothermal–sedimentary systems [84,85]. These signatures suggest the incorporation of fine-grained clastic material, likely transported from nearby continental or volcanic sources during deposition.

During the Late Triassic (221 to 204 Ma), the closure of the Paleo-Tethys Ocean in the West Kunlun–Altun region triggered intense tectonic activity. Post-collisional extensional collapse facilitated the emplacement of crust-derived intermediate to felsic granitoid magmas, including the Qiukutai pluton, along deep-seated faults cutting the ore-hosting strata. Zircon U–Pb dating indicates an emplacement age of 211.01 ± 0.82 Ma, and Hf isotopic data (average εHf(t) = −3.94) suggest a magma source dominated by ancient continental crust (Figure 13b). Hydrothermal fluids associated with the cooling of these intrusions led to minor overprinting of earlier siderite mineralization. Field evidence includes siderite- and quartz-siderite-filled fractures in both the orebody and host rocks [17]. These features likely reflect two concurrent processes: partial dissolution and remobilization of early siderite by hydrothermal fluids and direct Fe-rich fluid precipitation from the intrusions. Fluid inclusion studies on quartz veins related to this event reveal homogenization temperatures of 220–360 °C and salinities of 32.9–41.5 wt.% NaCl eqv., consistent with a moderate-temperature magmatic–hydrothermal regime distinct from the earlier low-temperature sedimentary system [3]. Such late-stage modification is spatially limited, and the primary sedimentary textures of the deposit remain well preserved.

Isotopic geochemical data further reveal the mixed-source nature of the ore-forming hydrothermal fluids. The MixSIAR model quantifies the carbon sources as predominantly marine carbonate (49.7%), with additional contributions from sedimentary organic matter (30.6%) and magmatic CO₂ (19.7%). The measured δ¹³C_PDB values (−6.0‰ to −4.6‰) fall between the typical ranges of carbonate and magmatic gases, supporting a multi-endmember fluid input model. The reconstructed δ¹⁸O_fluid value of the hydrothermal fluid averages –10.52‰, which is significantly lower than that of magmatic water. This indicates that the fluid was mainly derived from shallow seawater, which underwent substantial water–rock interaction with deeper thermal systems under open-system conditions. Combined with the δ¹⁸O-based temperature estimates (~55 °C), these results suggest that the Qieliekeqi siderite deposit formed in a low-salinity, low- to moderate-temperature hydrothermal environment within a shallow marine to continental shelf setting.

In summary, the formation of the Qieliekeqi siderite deposit was controlled by a two-stage metallogenic process involving Late Cambrian marine hydrothermal sedimentation and overprinting by Triassic magmatic–hydrothermal activity. This dual-stage mechanism is regionally representative. The metallogenic model proposed in this study provides important theoretical support and a comparative framework for understanding the genesis of similar hydrothermal sedimentary carbonate deposits in the West Kunlun region and adjacent areas.

7. Conclusions

Based on a comprehensive study of the geology, geochemistry, isotopic geochronology, and stable isotope characteristics of the Qieliekeqi siderite deposit, the following main conclusions can be drawn:

(1) The Qieliekeqi siderite deposit underwent multi-stage metallogenic evolution. The initial mineralization occurred during the Late Cambrian (ca. 500 to 488 Ma) within an intra-oceanic rift setting associated with the early rifting and spreading of the Proto-Tethys Ocean. Fe-rich hydrothermal fluids ascended along deep faults and mixed with oxidized seawater in a shallow marine environment, leading to the precipitation of siderite. The ores are mainly coarse-grained (average FeO: 47.44%, Al₂O₃/TiO₂: 29.14) and banded (average FeO: 48.18%, Al₂O₃/TiO₂: 17.07) in texture and exhibit geochemical signatures indicative of hydrothermal sedimentation, such as enrichment in Mn, Co, and Ba.

(2) Magmatic intrusion and hydrothermal overprinting during the Late Triassic (~211 Ma) modified the early-stage ore bodies. Zircon U–Pb dating indicates that the granitic intrusion has an age of 211.01 ± 0.82 Ma, and Hf isotopic data yield an average εHf(t) value of −3.94, suggesting that the magma was derived from partial melting of ancient continental crust. Degassing magmatic fluids locally reactivated the early siderite ores, resulting in metal remobilization and the formation of minor vein-type and disseminated siderite mineralization, as well as small-scale secondary hydrothermal precipitation within fractures.

(3) Stable isotope analyses reveal a mixed-source origin for the ore-forming fluids. MixSIAR carbon source modeling indicates that marine carbonate was the dominant contributor (49.7%), accompanied by sedimentary organic matter (30.6%) and magmatic CO₂ (19.7%). The reconstructed δ¹⁸O_fluid values average −10.52‰, significantly lower than those of magmatic water, suggesting that the mineralizing fluids were primarily derived from shallow, oxidized seawater. The more negative δ¹⁸O values observed in some samples may also reflect the involvement of meteoric or surface water in the hydrothermal system, resulting in mixed-source fluids heated by underlying magmatic activity. Combined with δ¹³C–δ¹⁸O-based temperature estimates, the mineralization occurred in a low- to moderate-temperature (53 to 58 °C) hydrothermal system in a shallow marine to continental shelf environment.

(4) The Qieliekeqi deposit represents a multi-stage metallogenic system jointly controlled by Cambrian seafloor hydrothermal sedimentation and Triassic magmatic–hydrothermal overprinting. The mineralization process reflects the coupled effects of deep-sourced thermal fluid input, shallow carbon sources, and the mixing of seawater with magmatic fluids. This metallogenic model provides a coherent explanation for the geochemical characteristics of the siderite ores in the study area, offering a robust theoretical framework for understanding the genesis of hydrothermal carbonate deposits in the West Kunlun region. It also holds important implications for regional mineral exploration and metallogenic prediction.