Confirmation of Significant Iron Formations During “Boring Billion” in Altyn Region, China: A Case Study of the Dimunalike Iron Deposit

Abstract

It is generally believed that the ancient oceans during the “boring billion” (1.85–0.8 Ga) lacked the capacity to form large-scale iron formations (IFs), though localized small-scale IFs deposition persisted. The Altyn region of China hosts abundant IFs, with the Dimunalike IFs being the largest and most representative, characterized by typical banded iron–silica layers. Detailed fieldwork identified a tuff layer conformably contacting the IFs at the roof rocks of IFs and a ferruginous mudstone layer at the floor rocks of IFs in drill core ZK4312. Geochemical and zircon U-Pb-Hf isotopic analyses were performed. The tuff has a typical tuff structure, mostly made of quartz, and contains a significant amount of natural sulfur. It also has high SiO₂ content (77.90%–80.49%) and sulfur content (0.78%–3.06%). The ferruginous mudstone has a volcanic clastic structure and is mainly composed of quartz and chlorite, with abundant coeval pyrite. It shows lower SiO₂ content (53.83%–60.32%) and higher TFe₂O₃ content (10.29%–16.24%). Both layers share similar rare earth element (REE) distribution patterns and trace element compositions, with light REE enrichment and negative Eu, Nb, and Ti anomalies, consistent with arc volcanic geochemistry. Zircon U-Pb ages indicate crystallization of the tuff at 1102 ± 13 Ma and maximum deposition of the mudstone at 1110 ± 41 Ma. These data suggest formation during different stages of the same volcanic–sedimentary process. The ε_Hf(t) values (3.60–12.35 for tuff, 2.92–8.19 for mudstone) resemble those of Algoma-type IF host rocks, implying derivation from re-melted new crust. The Dimunalike IFs likely formed in a submarine volcanic–sedimentary environment. In conclusion, although the Mesoproterozoic ocean was generally in a low-oxygen state, which was not conducive to large-scale IF deposition, localized submarine volcanic–hydrothermal activity could still lead to IF formation.

1. Introduction

Iron formations (IFs) are iron-rich chemical sedimentary rocks formed by specific geological process, typically characterized by iron content ≥ 15% and exhibiting a layered or slaty structure. The main mineral constituents are magnetite and hematite, with quartz as a secondary component, along with minor amounts of carbonates and silicates [1,2,3]. Based on the ore-forming environment, IFs can be classified into two types: Algoma-type and Superior-type [4,5]. Algoma-type IFs are notably associated with submarine volcanic–hydrothermal sources. Superior-type IFs are deposited in stable shelf shallow marine environments. They alternate rhythmically with normal marine carbonates and are accompanied by shales, conglomerates, and cherts, which are non-volcanic clastic rocks. Volcanic rocks are present in minimal amounts [4]. Although this classification system is widely accepted, some IFs exhibit characteristics of both types [5,6].

The formation and disappearance of IFs are often linked to significant environmental events. IFs were primarily formed between 3.8 Ga and 1.8 Ga, with peak periods around ~2.7 Ga, ~2.5 Ga, and ~1.9 Ga [5,7]. After ~1.8 Ga, dramatic changes in the marine environment and climate led to the near cessation of IF formation [3,5,7,8]. However, during ~0.8 to 0.6 Ga, IFs reappeared on a large scale due to the “Snowball Earth” event [9,10]. Although the period between 1.8 Ga and 0.8 Ga is considered the “boring billion” for IFs [11,12,13,14], global studies suggest that the ocean still had the potential to form localized IFs during this time. Environments such as submarine volcanic activity, hydrothermal systems, and rift basins provided the conditions for small-scale IF deposition [7,15]. Notable examples include (1) IFs associated with volcanogenic massive sulfide (VMS) deposits, such as layered iron oxides interlayered with sulfides in the Mesoproterozoic Pb-Zn-Ag deposits of Broken Hill, Australia, which suggest a contribution from hydrothermal deep-sea deposits [15,16,17]; (2) IFs in non-typical redox environments, such as the 1.72 Ga IFs in the Pecos Greenstone Belt of New Mexico, formed in a suboxic deep-sea environment, independent of sulfide conditions [18]; (3) rift-related IFs, such as the 1.7 Ga A-type rhyolitic volcanic clastic rock-associated IFs in the Olary region of South Australia [19] and the 1.1 Ga Jingtieshan IFs in the northern Qilian Mountains of China, whose sedimentary exhalative (SEDEX) type genesis may be related to rift basin hydrothermal activities [20,21,22]. Notably, Davies et al. [23] confirmed through in situ hematite U-Pb dating that the Hamersley Province of Western Australia, the world’s largest IF metallogenic camp, was formed between 1.4 and 1.1 Ga, with its formation closely related to supercontinent rifting and convergence, as well as submarine volcanic–hydrothermal activity. This provides strong evidence that large-scale IFs could form during the “boring billion”.

In recent years, significant progress has been made in iron ore exploration in the Altyn and North Qilian regions of China, with numerous IF iron ore deposits discovered. For example, new IFs have been exposed in the Quanji block on the northern edge of the Qaidam Basin [24], and the Baijianshan IFs, associated with the Snowball Earth event, have been identified in the eastern Altyn region, formed around 737 Ma [25]. These discoveries, together with the Jingtieshan IFs in the northern Qilian Mountains and the Dimunalike iron ore belt in Altyn, form an IF iron ore belt extending over 900 km (Figure 1a) [22,26]. Regional geological studies show that the IFs in this belt are often closely associated with volcanic activity and exhibit a clustered spatial distribution, revealing the potential for IF metallogenesis in the ancient oceanic environment of the Altyn and North Qilian regions.

The Dimunalike iron metallogenic belt in Altyn is part of the Tarim metallogenic province, where several iron ore deposits are located, including the Yuling iron deposit, Chengyi iron deposit, and Dimunalike iron deposit. Its prospective resource is estimated to reach 1 billion tons [27]. Two main viewpoints currently exist regarding its genetic type and ore-forming age: the early viewpoint holds that it is a marine volcanic–sedimentary type deposit formed between 623 and 500 Ma, related to the evolution of the South Altyn Ocean [28,29]; more recent studies by Li et al. [30] suggest that this ore deposit is a typical IF, with the age of basalts determined as 745 ± 1.6 Ma, providing an indirect constraint on the ore-forming age of the deposit. This study, based on detailed field geological surveys and related element-isotope geochemical and zircon U-Pb-Hf testing, precisely defines the ore-forming age of the Dimunalike IFs, revealing the tectonic environment and further improving the ore-forming theory of Mesoproterozoic IFs.

2. Regional Geology

The Altyn Orogenic Belt is an important component of the Central Orogenic Belt, extending in a northeast–southwest direction and connecting the Tarim Basin with the Tibetan Plateau. The region is divided into four tectonic units based on major regional faults, from north to south: the North Altyn Block, the North Altyn (Hongliugou–Lapeiquan) subduction mélange zone, the Central Altyn (Milanhe–Jinyanshan) Block, and the South Altyn (Apa–Mangya) subduction mélange belt [31,32]. These tectonic units are separated by the North Altyn Fault, Jinyanshan Fault, and Aihmaiti Kashi–Kangtugai Fault. The Altyn Orogenic Belt is a complex orogenic belt that has undergone multiple periods of tectonic activity and evolution [33]. It experienced Archaean crustal formation activities that formed the continental core and crystalline basement, Mesoproterozoic continental margin sedimentation, Neoproterozoic to Early Paleozoic plate expansion, and Caledonian plate subduction–collision events [33].

The Altyn region is extensively developed with iron metallogenesis, exhibiting multiple periods of mineralization (Figure 1a) [22,26]. The Dimunalike iron ore belt is controlled by the South Altyn Tagh Fault and is hosted within the flysch slices of the Changshagou tectonic ophiolitic complex belt [29]. The iron ores at Shaliangxi and Tatrek are found within the Mesoproterozoic Bashkurgan Group, a metamorphosed sedimentary–volcanic rock sequence [34]. The Jiefanggou iron ore is closely associated with the metamorphic basement of the Paleoproterozoic Dakendaban Group [24].

The iron-bearing strata of the Dimunalike iron metallogenic belt in the Altyn region are primarily composed of tectonically mixed flysch slices [29]. A widespread phenomenon of interleaved and overlaid rocks bodies is observed in this region, which has undergone significant metamorphic deformation. Previous studies have collectively referred to these mélange bodies as the “Changshagou tectonic serpentinite mélange” [35]. The geological survey report of the Gurgah map sheet (J45C003003) [36] at a scale of 1:250,000 classifies this formation as the Qingbaikou system Solkurli Group Pingwaogou Formation, while other scholars have attributed it to the Ordovician Qimantag Group of the Eastern Kunlun [37]. The differences in age classification reflect the complex tectonic history of the rock sequence and suggest that it may have undergone multiple overlapping phases of tectonic and thermal events. Therefore, accurately determining the formation age of the flysch slices is of significant importance for understanding the tectonic and ore-forming evolution of the Altyn Orogenic Belt. Ding et al. [27], through detailed geological mapping at a scale of 1:10,000, found that the region does not widely expose intermediate-basic volcanic rocks, but is primarily composed of flysch slices. After undergoing metamorphism at the greenschist facies, the lithology of these flysch slices has changed, with the primary lithologies being sericite phyllite, carbonaceous phyllite, quartzite, carbonaceous sericite phyllite, and marble (Figure 1b) [27]. Based on this, we tentatively name this unit the “Dimunalike flysch slices” (Figure 1a) [22,26].

3. Local Geology

3.1. Stratigraphy and Structure of the Local

The Dimunalike iron ore belt is controlled by the southern Altyn fault and exhibits a NEE-trending distribution, with the ore-hosting phyllite strata extending nearly east–west (Figure 1a,b) [22,26,27]. Six iron ore deposits have been identified within the belt, including two large-scale deposits and three medium- to small-scale deposits (Figure 1a,b) [22,26,27], with a total proven iron reserve of approximately 450 million tons [27].

The strata in the Dimunalike IF mining area generally trend SE and exhibit flysch formation characteristics. The exposed lithologies include carbonaceous sericite phyllite, sericite–quartz phyllite, carbonaceous marl, sericite phyllite, marbleized carbonaceous marl, quartzite, and magnetite quartzite (Figure 2) [29]. The sericite phyllite serves as the ore-hosting stratum, with local occurrences of tuffaceous phyllite, suggesting a volcanic–sedimentary environment. The mining area is primarily characterized by the NNE-trending F1 and F2 main controlling faults and NNW-trending secondary faults.

3.2. Ore Body Characteristics

The Dimunalike IFs extend discontinuously for approximately 2.6 km and can be divided into two mining segments, east and west (Figure 2) [29]. The eastern segment hosts the Fe36 and Fe37 main ore bodies, while the western segment contains the Fe1-Fe2-Fe3-Fe39 secondary ore bodies. The Fe36 ore body exhibits typical stratiform features, with its spatial distribution aligning with the layering of the local strata. The ore body generally trends NW with dip angles ranging from 32° to 57°, extending 1180 m, and has a total iron grade of 28.73%–30.5% (average 29.45%). This ore body has undergone plastic deformation under tectonic stress, forming undulating structures, with some areas displaying overturned folds. The adjacent Fe37 ore body extends 1726 m, with grades ranging from 25.14% to 30.85% (average 28.86%), and similarly trends NW with dip angles between 29° and 63°. In addition to plastic deformation, branching and complex features are also observed.

The roof rocks of the main Fe36 ore body are a 50 m-thick alteration zone, characterized by a bluish-gray color and mineral alterations primarily to goethite, followed by the formation of jarosite and other sulfate minerals [29]. At the bottom of the alteration zone, a volcanic rock layer was identified, in conformable contact with the Fe36 ore body (Figure 3), with a dip direction of 21° and a dip angle of 52°. Through our detailed petrographic examination, the analytical results demonstrate that the volcanic rock is tuffaceous in nature. The floor rocks of the Fe36 ore body are sericite phyllite, with single-layer thicknesses ranging from 0.2 to 0.6 m. The rocks are dark gray-green with a foliated texture and clearly visible remnant layering.

Drill core ZK4312 is located approximately 100 m northeast of the Fe36 main ore body (Figure 2) [29]. The drill log shows that at depths of 108.5–99 m, the roof rocks are green chlorite–sericite phyllite, gray-green in color and containing minor pyrite, in conformable contact with the ore body. The iron ore layer (depth 132–108.5 m) corresponds to the Fe36 main ore body, with some samples containing rhodochrosite. The floor rocks of the ore body are strongly chloritized green mudstone (depth 146–132 m), followed by a continuous succession of 35 m thick, alternating layers of carbonaceous dolomitic phyllite and green chlorite–sericite phyllite. Beneath the carbonaceous dolomitic phyllite, at depths below 181 m, iron-rich mudstone is found (Figure 4).

The Dimunalike ore body primarily displays a banded structure (Figure 5a–c), with relatively wide bands of silica and iron, approximately 3–5 cm thick. The silica and iron bands are gray-white and gray-black, respectively (Figure 5b). The ore minerals are primarily magnetite, exhibiting a subhedral-to-anhedral granular structure (Figure 5c,d). The magnetite has a steel-gray metallic luster (Figure 5c,d). The gangue minerals are quartz and chlorite. The quartz is transparent and colorless, occurring as anhedral, dense distributions within the silica bands (Figure 5c,d). In the iron bands, only a small amount of quartz is scattered among the chlorite (Figure 5d). The chlorite is light green, occurring in subhedral, acicular, and prismatic forms, densely and interwoven in the gaps between magnetite (Figure 5d).

The ore of the Dimunalike iron deposit is characterized by high iron content (Fe₂O₃ ranging from 27.15% to 50.8%, with an average of approximately 35.7%) and high silica content (SiO₂ ranging from 17.94% to 61.62%) [29]. It is also notably low in alkali elements, with Na₂O content ranging from 0.18% to 2.92% and K₂O content ranging from 0.09% to 0.54%. The contents of Al₂O₃ and TiO₂ are generally low, ranging from 2.33% to 4.34% and 0.17% to 0.38%, respectively, indicating the relative absence of impurity minerals. Some samples exhibit slightly elevated P₂O₅ content, with a maximum value of 0.754%. The levels of CaO and MgO are generally low, suggesting a limited presence of carbonate minerals such as calcite and dolomite.

3.3. Sample Characteristics

The samples reported in this study were collected from volcanic rock layers between the Fe36 main ore body and the alteration zone in the measured section (Figure 3) and from a mudstone layer at a depth of 181 m in drill core ZK4312 (Figure 4), corresponding to tuff and ferruginous mudstone samples, respectively.

The tuff lies above the iron ore body and is in conformable contact with it. It is light gray to light yellowish-gray in color (Figure 6a), with a tuffaceous structure and vesicular texture, containing a high amount of cryptocrystalline felsic minerals (Figure 6b,c), which are typical characteristics of volcanic rocks. The mineral grains are fine, less than 20 μm in size, primarily composed of quartz (about 65%), followed by cryptocrystalline feldspar, including potassium feldspar (about 20%) and sodium feldspar (about 8%), and native sulfur (about 3%) (Figure 6d). The native sulfur may be a composite product of primary native sulfur from volcanic eruptions and post-alteration of pyrite. These vesicular features and high sulfur content suggest that the tuff likely formed during the late stage of volcanic eruption.

The ferruginous mudstone lies below the iron ore layer in drill core ZK4312, is gray-green in color (Figure 6e), with a volcanic clastic structure (Figure 6f,g), and has undergone green schist facies metamorphism. The volcanic material constitutes approximately 85%, mainly composed of quartz and cryptocrystalline chlorite, with about 15% scattered anhedral pyrite (Figure 6g,h), likely formed during the rapid deposition of volcanic clasts. The lithic material accounts for about 10%, mainly consisting of foliated chlorite with a small amount of quartz (Figure 6f,g). The ferruginous mudstone likely originated from the metamorphism of sed-volcanic pyroclastic rocks formed during the early volcanic–sedimentary process.

4. Analytical Techniques

The whole-rock major and trace element analyses of the tuff and ferruginous mudstone were completed at Wuhan Sample Solution Analytical Technology Co., Ltd., Wuhan, China. The sample pretreatment of whole rock for major element analysis was made by the melting method. The flux is a mixture of lithium tetraborate, lithium metaborate, and lithium fluoride (45:10:5). Ammonium nitrate and lithium bromide were used as oxidants and release agents, respectively. The melting temperature was 1050 °C, and the melting time was 15 min. Zsx Primus II wavelength dispersive X-ray fluorescence spectrometer (XRF) (RIGAKU, Tokyo, Japan) was used for the analysis of major elements in the whole rock. The X-ray tube is a 4.0 Kw end window Rh target, the test conditions are voltage: 50 kV, current: 60 mA, and all major element analysis lines are kα. The standard curve uses the national standard material: rock standard sample GBW07101-14 [38]. The data were corrected by the theoretical α coefficient method. The relative standard deviation (RSD) is less than 2%. Trace element analysis of whole rocks was conducted on an Agilent 7700e ICP-MS at the Wuhan Sample Solution Analytical Technology Co., Ltd., Wuhan, China. The detailed sample-digesting procedure was as follows: (1) sample powder (200 mesh) was placed in an oven at 105 °C for drying of 12 h; (2) 50 mg sample powder was accurately weighed and placed in a Teflon bomb; (3) 1 mL HNO₃ and 1 mL HF were slowly added into the Teflon bomb; (4) Teflon bomb was put in a stainless steel pressure jacket and heated to 190 °C in an oven for >24 h; (5) after cooling, the Teflon bomb was opened, placed on a hotplate at 140 °C, and evaporated to incipient dryness, and then 1 mL HNO₃ was added and evaporated to dryness again; (6) 1 mL of HNO₃, 1 mL of MQ water and 1 mL internal standard solution of 1 ppm In were added, and the Teflon bomb was resealed and placed in the oven at 190 °C for >12 h; (7) the final solution was transferred to a polyethylene bottle and diluted to 100 g by the addition of 2% HNO₃.

In this study, zircon samples were selected from the tuff and ferruginous mudstone, with the zircon selection, preparation, CL imaging, and other tasks conducted by Langfang Tuoxuan Rock and Mineral Testing Service Co., Ltd., Langfang, China. The zircon U-Pb dating was performed at Nanjing Hongchuang Geological Exploration Technology Service Co., Ltd., Nanjing, China. The laser ablation (LA) system used was the Resolution SE model, a 193 nm deep ultraviolet laser ablation sample introduction system (Applied Spectra, Fremont, CA, USA). The mass spectrometer (ICPMS) used was an Agilent 7900 inductively coupled plasma mass spectrometer (Agilent, Santa Clara, CA, USA). During testing, the laser beam spot size was 32 μm, with a frequency of 6 Hz, using helium as the carrier gas and argon as the makeup gas. The instrument was optimized using the National Institute of Standards and Technology (NIST) SRM610 standard reference material [39], which was also used as the external standard for trace element analysis. The standard zircon 91500 was used as the external standard for age dating, and the Qinghu zircon standard was used as the monitoring sample. During the sample analysis, two standard zircon 91500 samples were measured for every five sample points. The first 20 s of each sample were used for background signal collection, and the sample signal collection time was 50 s. After testing, the data were processed using the ICPMSDataCal software (version 12.2, China University of Geosciences, Beijing, China) [40], and age calculation and concordia diagrams were generated using Isoplot 3.0.

The zircon Hf isotope in situ analysis was performed at the Institute of Geology and Geophysics, Chinese Academy of Sciences, in the multi-collector inductively coupled plasma mass spectrometry (LA-MC-ICP-MS) laboratory. A Geolas-193nm excimer laser sampling system (MicroLas, Göttingen, Germany) was used with a laser frequency of 8 Hz and a spot diameter of 60 μm, with a signal collection time of 26 s. Subsequently, the Thermo Fisher Neptune multi-collector inductively coupled plasma mass spectrometer (LA-MC-ICP-MS) was used for testing. The zircon PLE and Mud Tank standards were used as external standards for the experimental tests, and the analysis procedures are detailed in Wu et al. [41].

5. Analysis Results

5.1. Whole Rock Geochemical Characteristics

5.1.1. Major Elements

The major element analysis results [42] (

Table 1. Geochemical composition of the and ferruginous mudstone from the Dimunalike iron deposit.

Sample NO.	DIP3-1-1	DIP3-1-2	DIP3-1-4	DIP3-1-5	DIP3-1-6	ZK4312-181-1	ZK4312-181-2	ZK4312-181-3	ZK4312-181-4
Major elements (%)
SiO2	80.49	78.03	79.79	77.90	80.09	57.41	56.89	53.83	60.32
Al2O3	8.67	9.17	10.54	8.70	8.27	14.52	14.87	12.46	14.12
TiO2	1.07	1.16	1.24	1.05	1.13	1.28	1.27	1.15	1.21
TFe2O3	0.88	1.05	1.31	1.48	1.45	12.41	11.41	16.24	10.29
CaO	0.07	0.06	0.03	0.05	0.10	0.25	0.21	0.25	0.30
MgO	0.81	0.60	1.04	0.57	0.68	2.93	3.41	2.55	2.91
K2O	2.90	2.55	4.08	2.61	2.93	4.14	4.21	3.59	4.02
Na2O	0.95	1.88	0.23	1.41	0.60	0.11	0.11	0.10	0.11
MnO	0.00	0.00	0.00	0.00	0.00	0.02	0.02	0.02	0.02
P2O5	0.02	0.01	0.01	0.01	0.03	0.12	0.13	0.11	0.12
TS	2.02	2.68	0.78	3.06	2.30
LOI	4.12	5.05	1.84	6.05	4.63	6.62	7.28	9.54	6.43
Total	99.96	99.56	100.10	99.83	99.92	99.82	99.81	99.83	99.83
Rare earth and trace elements (ppm)
Li	19.89	17.14	27.21	16.40	18.66	39.1	43.0	36.2	39.2
Be	2.16	2.15	2.60	1.92	2.00	2.58	2.34	2.14	2.45
Sc	23.08	23.26	26.48	21.96	23.17	20.7	26.7	18.1	23.7
V	148.00	146.75	178.94	131.51	136.69	178.00	169.00	164.00	174.00
Cr	103.69	99.96	152.22	89.94	96.98	178.00	168.00	174.00	219.00
Co	0.29	0.17	0.16	0.22	0.40	27.6	23.9	29.3	25.1
Ni	2.70	11.57	2.24	2.71	6.77	96.3	76.9	101.00	80.0
Cu	4.89	5.36	5.33	4.96	5.55	101.00	69.0	71.0	74.6
Zn	21.32	21.98	25.77	19.68	21.73	78.2	79.4	58.1	67.7
Ga	13.39	13.62	16.83	12.52	12.62	22.9	21.6	20.1	22.9
Rb	108.31	92.05	140.75	98.84	105.80	160.00	152.00	139.00	158.00
Sr	20.70	31.13	12.66	22.01	14.89	16.6	16.8	15.8	18.0
Zr	208.43	236.43	225.64	214.34	249.83	223.00	216.00	194.00	200.00
Nb	20.55	22.00	20.89	19.58	21.82	25.3	23.2	20.9	22.0
Mo	0.84	1.30	0.40	0.74	0.49	0.56	0.74	0.64	0.94
In	0.10	0.08	0.12	0.08	0.08	0.15	0.14	0.11	0.13
Cs	2.74	2.45	3.86	2.84	3.14	3.54	3.77	3.17	3.48
Ba	486.23	444.17	697.82	419.67	440.32	648.00	593.00	551.00	628.00
Hf	7.08	7.91	7.49	7.33	8.52	7.31	7.07	6.83	6.80
Ta	1.60	1.72	1.58	1.55	1.73	1.69	1.61	1.51	1.47
W	3.30	4.17	2.68	3.58	4.29	0.91	1.46	0.46	1.09
Tl	0.61	0.52	0.81	0.55	0.60	1.07	1.04	0.99	1.04
Pb	24.68	24.56	12.69	28.42	25.53	56.7	51.2	77.9	50.5
Bi	0.56	0.93	0.21	1.10	1.64	7.82	1.70	1.48	1.73
Th	4.98	5.74	7.29	5.27	6.32	12.3	11.3	8.47	10.8
U	1.54	1.70	1.53	1.48	1.75	2.73	2.54	2.23	2.38
Y	25.80	31.20	41.07	34.52	34.04	30.3	30.6	26.5	29.9
La	30.84	29.64	31.04	38.12	39.65	35.7	32.4	48.9	68.9
Ce	64.47	59.77	61.29	76.07	83.61	72.4	68.3	96.9	136.00
Pr	7.62	7.02	7.48	9.12	9.64	9.28	8.83	12.6	17.5
Nd	29.51	27.03	28.90	36.08	37.05	34.9	33.3	47.8	66.4
Sm	5.61	5.40	5.76	6.60	6.88	7.05	6.49	9.12	12.4
Eu	0.92	0.98	1.03	1.14	1.12	1.28	0.89	1.18	1.61
Gd	4.91	5.28	6.04	6.08	5.94	6.45	5.67	7.36	9.70
Tb	0.84	0.96	1.17	1.09	1.09	1.12	0.97	1.04	1.30
Dy	5.08	5.89	7.45	6.57	6.95	6.20	5.54	5.07	5.90
Ho	1.00	1.20	1.52	1.31	1.38	1.18	1.08	0.94	1.05
Er	2.72	3.24	4.01	3.39	3.55	3.25	3.07	2.78	2.96
Tm	0.37	0.46	0.54	0.47	0.47	0.53	0.50	0.46	0.46
Yb	2.36	2.64	2.96	2.69	2.86	3.57	3.40	3.09	3.16
Lu	0.36	0.38	0.42	0.38	0.40	0.50	0.48	0.44	0.44
∑REE	182.4	181.08	200.68	223.63	234.62	213.71	201.52	264.18	357.68
δEu *	0.52	0.55	0.53	0.54	0.52	0.89	0.69	0.68	0.69
(La/Yb)N *	9.38	8.07	7.51	10.18	9.95	0.74	0.70	1.17	1.61

* δEu = Eu_N/(0.5 Sm_N + 0.5 Gd_N); (La/Yb)_N = La_N/Yb_N, where N indicates chondrite-normalized values (adapte from [42]).

) show that the tuff is rich in silicon and aluminum but poor in alkali, calcium, magnesium, and iron. Its SiO₂ content ranges from 77.90% to 80.49%, with an average of 79.26%; K₂O content ranges from 2.55% to 4.08%, with an average of 3.01%; Na₂O content ranges from 0.23% to 1.88%, with an average of 1.01%; TiO₂ content ranges from 1.05% to 1.24%, with an average of 1.13%; MgO content ranges from 0.57% to 1.04%, with an average of 0.74%; TFe₂O₃ content ranges from 0.88% to 1.45%, with an average of 1.23%; Al₂O₃ content ranges from 8.27% to 10.54%, with an average of 9.07%; S content ranges from 0.78% to 3.06%, with an average of 2.17%.

The major element analysis results (Table 1) show that the mudstone is rich in silicon, aluminum, iron, and potassium, but poor in calcium, sodium, and magnesium. Its SiO₂ content ranges from 53.83% to 60.32%, with an average of 57.11%; K₂O content ranges from 3.59% to 4.21%, with an average of 3.99%; Na₂O content ranges from 0.10% to 0.11%, with an average of 0.11%; TiO₂ content ranges from 1.15% to 1.28%, with an average of 1.23%; MgO content ranges from 2.55% to 3.41%, with an average of 2.95%; TFe₂O₃ content ranges from 10.29% to 16.24%, with an average of 12.59%; Al₂O₃ content ranges from 12.46% to 14.87%, with an average of 13.99%; loss on ignition (LOI) ranges from 6.43% to 9.54%, with an average of 7.47%.

5.1.2. Rare Earth and Trace Elements

The total rare earth element (REE) content of the tuff ranges from 181.08 × 10⁻⁶ to 234.62 × 10⁻⁶, with an average of 204.48 × 10⁻⁶. The (La/Yb)_N ratio ranges from 7.51 to 10.18, with an average of 9.02, indicating clear differentiation between light and heavy rare earth elements. The rare earth element distribution curve shows a right-tilted pattern with relative enrichment of light rare earth elements (LREE) and relative depletion of heavy rare earth elements (HREE) (Figure 7a). The δEu value ranges from 0.52 to 0.55, with an average of 0.53, reflecting a negative Eu anomaly.

The total REE content of the ferruginous mudstone ranges from 170.92 × 10⁻⁶ to 327.78 × 10⁻⁶, with an average of 229.95 × 10⁻⁶. The (La/Yb)_N ratio ranges from 6.84 to 15.64, with an average of 10.25, indicating clear differentiation between light and heavy rare earth elements. The rare earth element distribution curve also shows a right-tilted pattern with relative enrichment of LREE and relative depletion of HREE (Figure 7b). The δEu value ranges from 0.43 to 0.57, with an average of 0.47, reflecting a negative Eu anomaly.

In the primitive mantle normalized trace element spider diagrams [42] (Figure 7c,d), the five tuff samples and four ferruginous mudstone samples exhibit similar distribution patterns, characterized by enrichment in high field strength elements (HFSE), with Nb and Ti showing negative anomalies, similar to arc volcanic rocks. Nd shows a positive anomaly. Large ion lithophile elements (LILE) are enriched, with Rb, Ba, K, and Pb showing positive anomalies, while Sr shows a negative anomaly. P also shows a negative anomaly.

5.2. Results of Zircon U–Pb Dating

5.2.1. Tuff

The zircon grains in the tuff sample D1P3-1-6 range in length from 70 to 100 μm, with length to width ratios of 1:1 to 1:2. Most grains are prismatic with blunt conical terminations, while a few are elliptical. All grains are gray-brown in color (Figure 8). The CL images show that most zircons display typical oscillatory zoning (Figure 8), with significant variations in CL intensity, suggesting that they may be magmatic zircons.

LA-ICP-MS U-Pb isotopic dating and CL images of 28 zircons reveal three distinct zircon groups in the tuff (Figure 8). The first group consists of four scattered older zircons, with U and Th contents ranging from 601 to 1014 ppm and 461 to 699 ppm, respectively, and a Th/U ratio between 0.49 and 0.77, with an average value of 0.65. These zircons are likely to represent captured older zircons. The ²⁰⁷Pb/²⁰⁶Pb surface ages range from 1966 to 1672 Ma (Figure 8,

Table 2. Summary of U-Pb isotopic dating results.

No.	Isotopic Ratios						Age (Ma)						Concentration (ppm)
	207Pb/206Pb	±1σ (%)	207Pb/235U	±1σ (%)	206Pb/238U	±1σ (%)	207Pb/206Pb	±1σ (%)	207Pb/235U	±1σ (%)	206Pb/238U	±1σ (%)	Th	U	Th/U
D1P3-1-6, tuff analyzed by LA–ICP–MS
1	0.0752	0.0018	1.7917	0.0454	0.1723	0.0021	1072	43.5	1042	16.5	1025	23.0	1068	975	1.10
2	0.0893	0.0017	2.8991	0.0569	0.2349	0.0033	1411	39.8	1382	14.9	1360	17.4	478	1085	0.44
3	0.0890	0.0017	2.7877	0.0583	0.2265	0.0029	1406	37.0	1352	15.7	1316	15.1	467	595	0.79
4	0.0847	0.0014	2.2727	0.0498	0.1940	0.0033	1309	31.0	1204	15.5	1143	25.5	918	2369	0.39
5	0.0754	0.0020	1.7069	0.0468	0.1637	0.0022	1080	49.1	1011	17.6	978	21.4	350	1052	0.33
6	0.0879	0.0015	2.6073	0.0502	0.2148	0.0034	1389	32.9	1303	14.2	1255	18.3	1051	1680	0.63
7	0.0754	0.0013	1.7445	0.0388	0.1674	0.0028	1080	39.0	1025	14.4	998	21.7	115	2506	0.05
8	0.0768	0.0016	2.0002	0.0459	0.1886	0.0028	1117	40.7	1116	15.5	1114	24.7	428	878	0.49
9	0.0878	0.0019	2.7042	0.0634	0.2228	0.0035	1389	40.7	1330	17.4	1297	18.5	430	633	0.68
10	0.0762	0.0015	1.7952	0.0441	0.1706	0.0034	1102	38.9	1044	16.0	1015	22.8	561	1683	0.33
11	0.0849	0.0017	2.2151	0.0496	0.1886	0.0026	1315	38.9	1186	15.7	1114	25.0	888	1585	0.56
12	0.0767	0.0015	1.8251	0.0472	0.1719	0.0031	1122	37.8	1055	17.0	1023	22.8	381	1743	0.22
13	0.0817	0.0015	2.3940	0.0483	0.2118	0.0028	1239	36.6	1241	14.5	1239	14.9	162	689	0.24
14	0.0765	0.0019	1.8910	0.0445	0.1790	0.0021	1107	54.6	1078	15.6	1062	23.6	273	508	0.54
15	0.0773	0.0016	1.9676	0.0445	0.1841	0.0029	1129	41.2	1104	15.2	1089	24.7	447	940	0.48
16	0.0772	0.0016	1.9517	0.0473	0.1827	0.0030	1128	40.7	1099	16.3	1082	24.6	410	1278	0.32
17	0.0756	0.0022	1.9213	0.0548	0.1841	0.0031	1085	58.5	1089	19.1	1089	24.8	141	260	0.54
18	0.1207	0.0023	5.9543	0.1265	0.3565	0.0052	1966	33.6	1969	18.5	1966	24.8	461	601	0.77
19	0.0812	0.0018	2.3211	0.0554	0.2069	0.0035	1228	44.4	1219	17.0	1212	18.7	452	501	0.90
20	0.0859	0.0017	2.4851	0.0518	0.2092	0.0026	1400	37.5	1268	15.1	1225	13.8	567	712	0.80
21	0.1026	0.0017	4.0520	0.0864	0.2851	0.0039	1672	31.2	1645	17.4	1617	19.8	544	912	0.60
22	0.0880	0.0017	2.6312	0.0558	0.2161	0.0025	1383	36.0	1309	15.6	1261	13.5	383	1005	0.38
23	0.0867	0.0021	2.7877	0.0676	0.2331	0.0038	1355	52.3	1352	18.1	1351	20.0	167	276	0.61
24	0.1060	0.0017	4.4422	0.0784	0.3031	0.0036	1732	29.9	1720	14.7	1707	17.9	492	1014	0.49
25	0.0816	0.0017	2.2086	0.0462	0.1961	0.0026	1235	40.7	1184	14.6	1154	25.8	336	751	0.45
26	0.0764	0.0019	1.9776	0.0516	0.1877	0.0027	1106	50.0	1108	17.6	1109	24.7	200	464	0.43
27	0.1107	0.0021	4.4933	0.0991	0.2938	0.0045	1811	39.4	1730	18.4	1661	22.3	699	916	0.76
28	0.0748	0.0022	1.8454	0.0544	0.1789	0.0025	1062	54.6	1062	19.4	1061	23.3	344	407	0.84
ZK4312-181, ferruginous mudstone analyzed by LA–ICP–MS
1	0.0891	0.0021	3.1215	0.0827	0.2540	0.0060	1407	45.2	1438	20.4	1459	30.6	130	73	0.56
2	0.0723	0.0025	1.7616	0.0622	0.1767	0.0043	994	67.4	1031	22.9	1049	23.3	56	44	0.79
3	0.0808	0.0020	2.1969	0.0601	0.1973	0.0046	1215	48.5	1180	19.1	1161	24.9	187	87	0.46
4	0.0802	0.0019	1.9260	0.0501	0.1742	0.0041	1202	45.7	1090	17.4	1035	22.3	572	77	0.13
5	0.1129	0.0025	4.9929	0.1247	0.3208	0.0075	1846	39.9	1818	21.1	1794	36.5	123	60	0.49
6	0.0806	0.0026	2.1088	0.0703	0.1898	0.0046	1211	61.5	1152	23.0	1120	24.7	103	45	0.43
7	0.1197	0.0028	5.1663	0.1337	0.3130	0.0073	1952	41.4	1847	22.0	1756	36.0	222	146	0.66
8	0.0750	0.0019	1.9208	0.0516	0.1858	0.0043	1068	48.7	1088	18.0	1099	23.5	151	54	0.36
9	0.1043	0.0025	4.0816	0.1058	0.2840	0.0066	1701	42.8	1651	21.1	1611	33.3	95	71	0.75
10	0.0947	0.0024	3.5303	0.0980	0.2704	0.0064	1522	47.7	1534	22.0	1543	32.2	69	56	0.82
11	0.0764	0.0018	1.9694	0.0512	0.1869	0.0043	1106	46.5	1105	17.5	1105	23.6	191	61	0.32
12	0.0963	0.0023	3.4118	0.0895	0.2571	0.0060	1553	44.3	1507	20.6	1475	30.7	104	124	1.19
13	0.0787	0.0025	2.3448	0.0785	0.2161	0.0052	1164	62.2	1226	23.8	1261	27.4	49	31	0.64
14	0.0792	0.0017	1.8595	0.0444	0.1703	0.0039	1178	41.4	1067	15.8	1014	21.5	551	48	0.09
15	0.1286	0.0028	6.5428	0.1573	0.3691	0.0085	2079	37.3	2052	21.2	2025	40.1	133	162	1.22

), which may indicate the presence of ancient Proterozoic basement rocks in the Altyn region. The second group consists of 12 older zircons, with U and Th contents ranging from 276 to 2369 ppm and 162 to 1051 ppm, respectively, and Th/U ratios between 0.24 and 0.90, with an average of 0.57. These zircons are likely to be captured magmatic zircons. The surface ages are highly scattered, ranging from 1411 to 1228 Ma (Figure 8, Table 2). The third group consists of 12 younger zircons, 11 of which have U and Th contents ranging from 260 to 1743 ppm and 141 to 1068 ppm, respectively, with Th/U ratios ranging from 0.22 to 1.10, with a large variation and an average of 0.51 (Figure 8, Table 2). One zircon has U and Th contents of 2506 ppm and 115 ppm, with a Th/U ratio of 0.05, which may have undergone metamorphic alteration, although this has limited impact on its surface age. The ²⁰⁷Pb/²⁰⁶Pb surface ages of the 12 points range from 1129 to 1062 Ma, with a weighted average of 1102 ± 13 Ma (MSWD = 1.0, Figure 8), which likely represents the crystallization age of the tuff.

5.2.2. Ferruginous Mudstone

The zircon grains in the ferruginous mudstone sample ZK4312-181 range in length from 40 to 100 μm, with an aspect ratio of 1:1 to 1:1.5. Most of the grains are elliptical in shape, with a few prismatic grains exhibiting blunt conical terminations. All the grains are gray-brown in color (Figure 9). The CL images show that most zircons exhibit typical oscillatory zoning (Figure 9), with significant variations in CL intensity, suggesting that they are likely magmatic zircons.

LA-ICP-MS U-Pb isotopic dating, CL images, and trace element characteristics of 15 zircons indicate the presence of two zircon groups in the ferruginous mudstone (Figure 9). The first group consists of seven scattered older zircons, with U and Th contents ranging from 69 to 222 ppm and 46 to 162 ppm, respectively. The Th/U ratio ranges from 0.33 to 1.22, with an average value of 0.78. These zircons are likely to represent captured older zircons. The ²⁰⁷Pb/²⁰⁶Pb surface ages range from 2079 to 1407 Ma (Figure 9, Table 2), potentially indicating the presence of older Proterozoic basement rocks in the Altyn region. The second group consists of eight younger zircons, seven of which have U and Th contents ranging from 49 to 572 ppm and 31 to 87 ppm, respectively, with Th/U ratios ranging from 0.13 to 0.79, with an average of 0.45 (Figure 9, Table 2). One zircon has U and Th contents of 551 ppm and 48 ppm, with a Th/U ratio of 0.09, which may have undergone metamorphic alteration, although this has limited impact on its surface age. The probability distribution of the 15 zircons shows that the zircon formation ages are mainly concentrated between 1100 and 1200 Ma. The ²⁰⁷Pb/²⁰⁶Pb surface ages of the eight points range from 1033 to 1187 Ma, with a weighted average of 1110 ± 41 Ma (MSWD = 1.3, Figure 9), which likely represents the maximum depositional age of the ferruginous mudstone.

5.3. Results of Zircon Hf Isotopes

This study conducted in situ zircon Lu-Hf isotopic testing on selected zircons representing the crystallization age of the tuff and the maximum depositional age of the ferruginous mudstone. In the tuff samples, the ¹⁷⁶Hf/¹⁷⁷Hf ratio of nine points ranges from 0.282199 to 0.282433 (

Table 3. Zircon Hf isotopic results of the tuff and ferruginous mudstone from the Dimunalike iron deposit.

NO.	t (Ma)	2σ	176Yb/177Hf	176Lu/177Hf *	176Hf/177Hf *	2σ	fLu/Hf	εHf (0)	εHf (t)	2σ	TDM1 (Hf)	TDM2 (Hf)
01	1102	30	0.0484	0.0017	0.282439	0.000029	−0.95	−11.78	11.41	1.02	1170	1188
02	1102	30	0.0154	0.0006	0.282443	0.000070	−0.98	−11.65	12.35	2.47	1132	1141
03	1102	30	0.0198	0.0007	0.282332	0.000052	−0.98	−15.55	8.33	1.84	1290	1345
04	1102	30	0.0717	0.0025	0.282263	0.000053	−0.93	−17.98	4.62	1.86	1450	1532
05	1102	30	0.0510	0.0018	0.282271	0.000025	−0.95	−17.72	5.39	0.90	1413	1493
06	1102	30	0.0303	0.0012	0.282263	0.000021	−0.96	−18.01	5.54	0.74	1402	1486
07	1102	30	0.0243	0.0010	0.282391	0.000026	−0.97	−13.47	10.25	0.91	1215	1247
08	1102	30	0.0408	0.0014	0.282345	0.000019	−0.96	−15.10	8.29	0.68	1294	1346
09	1102	30	0.0227	0.0008	0.282199	0.000020	−0.98	−20.25	3.60	0.70	1475	1584
10	1110	30	0.0226	0.0008	0.282207	0.000017	−0.98	−19.98	4.03	0.61	1465	1569
11	1110	30	0.0178	0.0006	0.282267	0.000015	−0.98	−17.86	6.27	0.54	1376	1455
12	1110	30	0.0271	0.0010	0.282330	0.000013	−0.97	−15.65	8.19	0.46	1304	1358
13	1110	30	0.0283	0.0010	0.282181	0.000016	−0.97	−20.91	2.92	0.58	1511	1625
14	1110	30	0.0514	0.0018	0.282305	0.000015	−0.95	−16.52	6.77	0.52	1365	1430

* ¹⁷⁶Lu/¹⁷⁷Hf = 0.0384; ¹⁷⁶Hf/¹⁷⁷Hf = 0.28325; λ = 1.867 × 10⁻¹¹ year⁻¹ [43].

), with an average value of 0.282327, showing a narrow variation range. In the ferruginous mudstone samples, the ¹⁷⁶Hf/¹⁷⁷Hf ratio of five points ranges from 0.282181 to 0.282330 (Table 3), with an average value of 0.282258, again showing a narrow variation range, indicating uniform Hf isotopic distribution in both volcanic rock zircons, suggesting a single magma source region. The ¹⁷⁶Yb/¹⁷⁷Hf ratios for the tuff range from 0.0154 to 0.0717 (Table 3), and for the ferruginous mudstone, they range from 0.0178 to 0.0514 (Table 3). The majority of zircons have a ¹⁷⁶Lu/¹⁷⁷Hf ratio less than 0.002, indicating that Hf formed by Lu decay is minimal, so the ¹⁷⁶Hf/¹⁷⁷Hf measurements can reflect the Hf composition of zircons at the time of crystallization [43]. The zircon ε_Hf(t) values and two-stage model ages (T_DM2) were calculated using the weighted average ²⁰⁷Pb/²⁰⁶Pb ages. For the tuff, using an age of 1102 Ma, ε_Hf(t) ranges from 2.81 to 11.55 (average 6.98), and T_DM2 ranges from 1584 to 1141 Ma (Table 3); for the ferruginous mudstone, using an age of 1110 Ma, ε_Hf(t) ranges from 2.92 to 8.19 (average 5.63), and T_DM2 ranges from 1625 to 1358 Ma (Table 3), representing the time when new crust differentiated from the depleted mantle. Zircons from both the tuff and ferruginous mudstone show consistently positive ε_Hf(t) values, suggesting a homogeneous source with magma mainly derived from juvenile crustal remelting.

6. Discussions

6.1. Ore-Forming Age of the Dimunalike IFs

There are currently three main views regarding the ore-forming age of the Dimunalike IFs. Hu et al. [37] suggested that the iron deposit belongs to the Ordovician Qimantagh Group; Yang et al. [28], based on a zircon age of 622.6 ± 1.4 Ma from a volcanic breccia, combined with regional tectonic events, proposed an Ediacaran–Cambrian (500–623 Ma) metallogenic epoch; while Li et al. [30], based on a zircon age of 745.2 ± 1.6 Ma from a basic metamorphosed volcanic rock, supported a Tonian metallogenic epoch. The sampling stratigraphy in the aforementioned studies may have some flaws. For instance, the basaltic volcanic breccia selected by Yang et al. was taken from a core (ZK4005) that did not contain iron ore, and the relationship between this breccia and the iron ore layer is unclear [28]. Additionally, the altered basaltic volcanic rocks collected by Li et al. were not sampled from cores or measured profiles, thus lacking a clear relationship with iron mineralization [30]. Therefore, it is essential to conduct chronological studies on strata that are more closely associated with the iron ore to precisely constrain the ore-forming age of the Dimunalike IFs.

As a typical IF, the ore-forming age of the Dimunalike IFs is often constrained by its spatial and temporal relationship with coexisting volcanic rocks [3]. The Dimunalike tuff developed above the Fe36 main ore body and formed at 1102 ± 13 Ma, while the ferruginous mudstone developed beneath the Fe36 main ore body and was deposited no later than 1110 ± 41 Ma. These layers show significant differences in SiO₂ and TFe₂O₃ content, but similar REE distribution patterns, suggesting that they may have formed at different stages of the same volcanic activity. The IF ore layer likely formed during an intervolcanic period of volcanic activity, deposited around 1.1 Ga.

Based on the above findings, we propose that the Dimunalike iron metallogenic belt is likely not part of the southern Qimantagh Group nor classified under the Solkuli Group, but rather is more likely a part of the Altyn Crustal Rocks within the Altyn Rock Group. Hao [26] conducted U-Pb dating on detrital zircons from feldspathic quartz schist, quartz schist, mica schist, and mica quartz schist within the Altyn Rock Group, and the youngest detrital zircons have ages ranging from 1181 to 1084 Ma. The ore-forming age of the Dimunalike iron deposit is very close to the formation epoch of the Altyn Crustal Rocks, and spatially, the Dimunalike flysch slices and the Altyn Crustal Rocks are in a tectonic mosaic relationship. This suggests a close relationship between the two, and the Dimunalike IFs are more likely a part of the Altyn Crustal Rocks, although it was incorporated into the tectonic mélange under the influence of the southern fault zone.

6.2. Provenance of the Dimunalike IFs

As previously mentioned, the Dimunalike flysch slices is closely related to the Altyn Crustal Rocks within the Altyn Rock Group. Therefore, it is crucial to clarify the properties of the Altyn Rock Group. Recent research has revealed that the Altyn Rock Group are not, as previously believed, part of the Archaean–Proterozoic crystalline basement of the Tarim Craton, but rather constitute a tectonic mélange composed of rocks of different ages and attributes. These include metamorphosed crustal rocks, various types of intrusive rocks, and exogenous rocks that have undergone high-pressure to ultra-high-pressure metamorphism [31,32]. Additionally, they have been significantly affected by Early Paleozoic deep subduction metamorphism [31,44]. The ε_Hf(t) values of the Altyn Crustal Rocks range from −2.89 to 7.04 [26], which are similar to the ε_Hf(t) values of the Dimunalike tuff and ferruginous mudstone, indicating that the Dimunalike IFs may be related to the evolution of the South Altyn Ocean.

Based on Hf isotopic studies of typical IF host rocks (Supplementary Table S1) [45,46,47,48,49,50,51,52], the ε_Hf(t) values of Algoma-type and Superior-type IF host rocks show significant differences: the ε_Hf(t) values of Algoma-type IF host rocks are mostly positive, indicating a mantle-dominated source (Figure 10) [43]; while the ε_Hf(t) values of Superior-type IF host rocks are mostly concentrated between −2.22 and +1.10, indicating a mixed crust-mantle source (Figure 10) [43]. This difference aligns with their formation environments. Algoma-type IFs form in arc-related setting, where magma is predominantly sourced from the mantle, with a small amount of upper crustal sedimentary debris incorporated, leading to some zircons having relatively low positive or even negative ε_Hf(t) values. Superior-type IFs, on the other hand, develop at stable cratonic margins, where the metallogenic material is primarily derived from upper crustal sedimentary debris, resulting in zircon ε_Hf(t) values fluctuating near zero [53]. The ε_Hf(t) values of the top and bottom host rocks of Dimunalike IFs range from 3.60 to 12.35 and 2.92 to 8.19, respectively, which are similar to the ε_Hf(t) values of Algoma-type IF host rocks, suggesting that they may also have formed in a volcanic context with the re-melting of new crustal material into the lithospheric system [5,54].

6.3. Implications of Confirmation the Mesoproterozoic IFs in the Altyn Region

IFs are highly sensitive to the redox state of the ocean. Their global enrichment is primarily concentrated in the Archaean to Paleoproterozoic (~3.8–1.8 Ga) and late Neoproterozoic (~0.8–0.6 Ga), with a significant decrease in IFs during the “boring billion” (1.8–0.8 Ga). Three competitive hypotheses have been proposed to explain this phenomenon: (1) the classical oxidation model suggests that the ocean was generally oxygen-rich during the Mesoproterozoic, inhibiting the dissolution and precipitation of Fe²⁺ [55,56]; (2) the sulfidic ocean model emphasizes that sulfide-rich water bodies reacted with Fe²⁺ to form pyrite, thereby hindering IF formation [57]; (3) the sulfide wedge hypothesis proposes that local sulfide zones along continental margins selectively consumed Fe²⁺, while open ocean basins could still preserve Fe²⁺ activity [58]. Recent studies have revealed that the redox state of the Mesoproterozoic ocean was spatially and temporally heterogeneous: while the atmosphere–ocean system overall exhibited low oxygen fugacity [59], with transient oxygenation pulses [60], deep-sea regions maintained anoxic to hypoxic conditions that were favorable for iron precipitation over long periods [59,61], providing a potential environment for localized IF formation.

Although the overall oceanic conditions of the Mesoproterozoic were not conducive to large-scale IF deposition, specific tectonic–hydrothermal contexts still provided local ore-forming potential. Several examples of “boring billion” IFs have been identified globally, with their ore-forming mechanisms closely related to unique environments, for instance, IFs associated with volcanic hydrothermal activity in Broken Hill, Australia, and Yavapai, USA [15,16,17,62]; IFs associated with rift basins in Olary, South Australia, and the Qilian Mountains [19,20,22]. Notably, Davies et al. [23] confirmed through in situ hematite U-Pb dating that the Hamersley Province in Western Australia, the world’s largest IF ore-forming domain, formed between 1.4 and 1.1 Ga, rather than during the previously assumed 2.2–2.0 Ga period, which was believed to be linked to the Great Oxidation Event [63,64,65,66,67,68,69].

The Dimunalike IFs and the Hamersley Province IFs not only share similar ore-forming ages but also have similar ore-forming environments [30]. The 1.1 Ga IF ore-forming event in the Dimunalike region coincides with the rifting of the Columbia supercontinent [70]. During the Mesoproterozoic, the Altyn Block was positioned between the Tarim Plate and the Australian Plate, and before 1.1 Ga, it was likely connected to the Western Australian Plate [70,71]. Thus, the Dimunalike IFs and the Hamersley Province IFs may have similar tectonic backgrounds, and the Dimunalike IF metallogenic belt may have an affinity with the IFs of the Pilbara Craton, showing significant ore-forming potential. This spatial–temporal coupling offers a potential direction for searching for large IF deposits from the same period in the Altyn region.

Although many Mesoproterozoic IFs have been reported globally, most of these deposits, except for the Jingtieshan IFs in the Qilian Mountains, China (with reserves exceeding 600 million tons), are relatively small or occur as by-products of VMS/SEDEX deposits. In contrast, the Dimunalike iron ore belt also has over 450 million tons of reserves, forming a “boring billion” IF metallogenic belt along the northern margin of the Tibetan Plateau, together with the Jingtieshan IFs. The discovery of this metallogenic belt not only confirms the potential for large-scale IF formation in this region but also provides important theoretical support for regional exploration. Importantly, the Dimunalike IFs (1.1 Ga), Jingtieshan IFs, and Hamersley Province IFs collectively demonstrate that low-oxygen oceans during the Mesoproterozoic could host large-scale IFs, likely associated with extensive submarine hydrothermal activity [21,23]. This groundbreaking discovery not only enriches the theory of IF metallogenesis but also provides a classic example for studying the ore-forming ages of global IFs.

7. Conclusions

(1) By U-Pb dating of volcanic zircons from the roof and floor rocks of the iron ore layer, this study has clearly determined that the Dimunalike IFs formed during the Mesoproterozoic (1151–1089 Ma), resolving previous controversies regarding the ore-forming age of the Dimunalike IFs.

(2) Zircon Hf isotopes (with predominantly positive ε_Hf(t) values) and the geochemical characteristics of the host rocks indicate that the metallogenic material of the Dimunalike IFs primarily originates from the re-melting of new crust. The Dimunalike IFs have similar ε_Hf(t) values to those of Algoma-type IFs, suggesting a connection with submarine volcanic activity.

(3) The Dimunalike IF metallogenic belt (with reserves exceeding 450 million tons), along with the Jingtieshan IFs in the Qilian Mountains, forms the Mesoproterozoic IF metallogenic belt along the northern edge of the Tibetan Plateau. This finding confirms that during the “boring billion,” localized low-oxygen submarine volcanic environments could still host large-scale IFs. This discovery not only enriches the metallogenic theory of Mesoproterozoic IFs in the Altyn region but also provides crucial geological evidence for exploring contemporaneous IF deposits in the region.