Mineralogical and Geochemical Characteristics of the Fe-Ti Mineralized Mafic-Ultramafic Intrusions at Wajilitag, Tarim Basin, China: With Special Emphasis on the Role of Apatite

Abstract

The Early Permian Tarim Large Igneous Province is a prominent magmatic-metallogenic province in China, hosting significant Fe-Ti mineralized mafic-ultramafic intrusions. Among them, the Wajilitag Fe-Ti oxide deposit stands out, which is hosted by olivine pyroxenite, clinopyroxenite, and gabbro. In the present study, we have examined the mineralogical and geochemical characteristics of apatite to elucidate a deeper understanding of the magmatic evolutionary processes and source characteristics of the mafic-ultramafic intrusions in the Wajilitag area. Petrographic analysis revealed three distinct types of apatite: (1) an inclusion phase within pyroxene and plagioclase, (2) an intergranular phase associated with Fe-Ti oxides, and (3) a late-stage phase found in association with biotite and/or amphibole. Geochemical analysis showed that the inclusion and intergranular apatites exhibited high fluoride (F) and low chlorine (Cl) concentrations, while the late-stage apatite displayed the reverse. A negative correlation between F and Cl was observed, suggesting different formation conditions for each apatite type. The high F/Cl ratios (>3) and enrichment of light rare earth elements (LREEs/HREEs = 12.8–29.5) in the apatite, in conjunction with Sr/Th-La/Sm diagrams, indicated that the parent magma originated from an enriched mantle source, influenced by ancient subduction-related fluids. Furthermore, low sulfur content (0.01%–0.16%) in apatite, along with estimated melt sulfur concentrations (19–54 ppm), points to a low sulfur fugacity environment. These findings collectively suggest that the Wajilitag deposit formed from magma derived from partial melting of an enriched mantle, followed by extensive magmatic differentiation, crystallization of Fe-Ti oxides, and low sulfur fugacity conditions.

1. Introduction

Apatite, a key accessory mineral in mafic-ultramafic intrusions, serves as an effective tracer for studying magmatic volatiles and mineralization processes due to its distinctive chemical composition [1]. This mineral is exceptionally resistant to weathering and preserves critical information about parental magma through its chemical signatures, making it highly sensitive to changes in the crystallization environment [2,3,4,5]. Recent studies have shown that apatite provides a more detailed geological picture than whole-rock geochemistry. The halogen content (F, Cl) provides insights into fluid sources and volatile content [4,6,7], whereas sulfur content indicates magmatic oxygen fugacity and sulfide saturation [4,8]. In addition, trace elements such as rare earth elements (REEs) and Sr are invaluable in tracing magmatic evolution [9]. These characteristics make apatite indispensable in studies of globally significant layered intrusions, including the Bushveld and the Sept Iles complexes [10,11,12].

The Early Permian Tarim Large Igneous Province is a significant magmatic-metallogenic province in China, hosting multiple mafic-ultramafic intrusions with magmatic Fe-Ti oxide mineralization. In comparison to typical deposits found in regions such as the Bushveld, Sept Îles, and Panzhihua, the Wajilitag complex exhibits distinctive mineralization characteristics. The Wajilitag complex is marked by indistinct rhythmic bedding and complex differentiation, yet it exhibits whole-rock mineralization, suggesting a unique magmatic evolution process. While previous studies have extensively examined the Tarim Large lgneous Province, the role ofapatite in magmatic evolution and mineralization remains poorly understood. This study aims to fill this gap by analyzing the mineralogical and geochemical features of apatite.

The present study examines the major and trace element composition of apatite from the Wajilitag deposit using petrographic observation, electron probe microanalysis (EMPA), and in situ laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS). This research addresses several key scientific questions: (1) the composition and evolutionary history of magmatic volatiles recorded in apatite; (2) the physicochemical conditions controlling magmatic Fe-Ti oxide mineralization; and (3) the relationship between magmatic differentiation in chambers and the resulting mineralization responses. Such information is crucial to provide new insights into the mineralization mechanisms within the Tarim Large Igneous Province, and offers valuable theoretical support for regional mineral exploration.

2. Geological Setting

The Tarim Craton is one of the three ancient cratons that form the primary tectonic framework of China. Located along the southern margin of the world’s largest Phanerozoic accretionary orogenic belt, the Central Asian Orogenic Belt (Figure 1a), the Tianshan Orogenic Belt borders the Tarim Craton to the north, the Kunlun Orogenic Belt to the southwest, and the Altyn Orogenic Belt to the southeast. The region displays an irregular rhombic shape, with a total area of 6.0 × 10⁵ km². The Tarim Craton is composed of two distinct components: a Precambrian metamorphic crystalline basement overlain by Phanerozoic sedimentary cover, forming a typical two-layer structure [13]. The Precambrian crystalline basement dates back to the Archean and Proterozoic eras, with major outcrops found in the Kuluktag and Aksu-Tiekele regions [14]. Thick Phanerozoic sedimentary strata, including shallow marine and terrestrial volcanic sediments, cover the Precambrian basement [15]. These strata primarily consist of Cambrian, Ordovician, Silurian, Devonian, Carboniferous, Permian, Cretaceous, and Paleogene-Neogene systems.

The Tarim Igneous Province is located along the western margin of the Tarim Craton (Figure 1b). It consists of extensive Early Permian continental flood basalts, mafic-ultramafic intrusions/dike swarms, alkali syenite and A-type granite bodies, along with minor picrite, andesite, rhyolite, tuff, tephrite, nephelinite, carbonatite, and calcite nephelinite. [17,18,19,20]. The formation of the Tarim Large Igneous Province represents the most intense and geographically extensive magmatic event in the geological history of the Tarim Craton. The Permian magmatic activity gave rise to three significant mafic-ultramafic intrusions, which host large-scale magmatic Fe-Ti oxide deposits: the Xiaohaizi (Mazartag, Qiawak South) Intrusion, the Wajilitag Intrusion, and the Puchang (Piquang) Intrusion.

The Wajilitag area is located in the Bachu region of southern Xinjiang, positioned along the western margin of the Tarim Craton. The area is characterized by various rock types, primarily mafic-ultramafic layered intrusions, calcite nephelinites, tephrite porphyry dikes, alkali syenites, alkali lamprophyre dikes, syenite porphyry dikes, nephelinites and carbonatites. Together, these constitute a complex porphyry system [21,22]. The complex displays a roughly north–south elongated, elliptical outline, covering an exposed area of approximately 12 km² and intruding into the Upper Devonian Yimugangtawu and Keziltege formations (Figure 2). The mafic-ultramafic layered intrusive body forms the primary component of the Wajilitag Complex. From a stratigraphic perspective, the formation is predominantly composed of gabbro, clinopyroxenite, and minor olivine pyroxenite, with a gradual transition from the upper to the lower strata [23]. An alkali syenite, occupying a very small area (<1 km²), is located at the summit of this body. Locally, minor amounts of nephelinite and extensive intrusive dikes are found within both the intrusive body and its host rock, including calcite-bearing nephelinite pipes and dikes, tephrite porphyry dikes, carbonate dikes, alkali lamprophyre dikes, and a series of intermediate-acidic dikes (diorite gabbro dikes and syenite dikes), generally trending northeast. Age dating of different crystalline rocks in the area can be summarized as follows, from oldest to youngest: Mafic-ultramafic layered intrusions (284–281 Ma), Calcite nephelinites (284–268 Ma), Tephrite porphyry dikes (283–272 Ma), Alkali syenites (282–277 Ma), Alkali lamprophyre dikes, Syenite porphyry dikes (275–274 Ma), Nephelinites (268 Ma), and Carbonatites (266 Ma) [18,19,21,24,25,26].

3. Petrology

3.1. Petrographic Characteristics

Olivine pyroxenite is primarily composed of clinopyroxene (50–60 vol.%), olivine (10%–15%), plagioclase feldspar (10%–20%), Fe-Ti oxides (5%–10%), and minor amounts of apatite (<1%). Olivine occurs as subautomorphic grains ranging in size from 0.3 to 2 mm, with some olivine grains displaying signs of alteration to serpentine. Monoclinic pyroxene also appears subautomorphic, with grain sizes varying from 0.5 to 3 mm. Plagioclase is present, filling irregular voids between olivine and clinopyroxene grains. Fine-grained Fe-Ti oxides are present in the interstices between olivine and clinopyroxene grains or encapsulated within olivine or clinopyroxene crystals.

Clinopyroxenite can be classified into two varieties: coarse-grained and fine-grained. Coarse-grained clinopyroxenites are primarily composed of clinopyroxene (75%–80%), Fe-Ti oxides (5%–10%), plagioclase feldspar (<5%), and olivine (<3%). The clinopyroxene exhibits an orthohedral-suborthohedral prismatic crystal form with a medium- to coarse-grained structure, and grain sizes range from 0.5 to 8 mm along the long axis. Ilmenite exsolution structures are commonly observed within the clinopyroxene. The distribution of orthohedral-suborthohedral Fe-Ti oxide inclusions within the clinopyroxene is a notable feature of the specimen.

Fine-grained clinopyroxenite consists predominantly of clinopyroxene (75%–85%) and Fe-Ti oxides (15%–20%). The rock displays an equigranular texture, with clinopyroxene grains measuring approximately 0.5 mm × 0.5 mm. Fe-Ti oxides fill the spaces between pyroxene grains. Brown amphibole reaction rims commonly form at the contact zones between clinopyroxene and Fe-Ti oxides. Additionally, the clinopyroxene grains exhibit sericite alteration, likely resulting from interaction with Fe-Ti oxides. These oxides, primarily magnetite and ilmenite, are intimately intergrown. Most ilmenite grains display relatively uncontaminated surfaces, with a few showing exsolution of hematite lamellae. Minor quantities of sulfides are also present, such as pyrite and pyrrhotite.

Gabbro is primarily composed of plagioclase feldspar (40%–50%), clinopyroxene (35%–45%), Fe-Ti oxides (<10%), amphibole (<5%), and apatite (<5%). The rock exhibits a gabbroic texture, characterized by tabular plagioclase crystals with widths ranging from 0.2 to 1 mm. Well-formed clinopyroxene crystals display amphibole reaction rims along their edges. Fe-Ti oxides, predominantly magnetite and ilmenite, fill irregular spaces between silicate minerals. Abundant euhedral apatite crystals are frequently engulfed by plagioclase, amphibole, or biotite. Olivine pyroxenite is primarily composed of clinopyroxene (50–60 vol.%), olivine (10%–15%), plagioclase feldspar (10%–20%), Fe-Ti oxides (5%–10%), and minor amounts of apatite (<1%).

3.2. Magmatic Evolution of Mafic–Ultramafic Intrusion in the Wajilitag Area

The Wajilitag complex displays a trend of decreasing basicity through crystallization differentiation, progressing from the lower olivine-pyroxene belt through the middle clinopyroxene belt to the upper gabbro belt. Microscopic analysis was conducted to confirm the sequence of mineral evolution associated with magmatic processes, which was found to follow this sequence: olivine + pyroxene + plagioclase ± Fe-Ti oxides → pyroxene ± plagioclase ± Fe-Ti oxides ± apatite → pyroxene + Fe-Ti oxides ± apatite. Three distinct stages of pyroxene crystallization are discernible under microscopic examination. In the initial phase of magmatic evolution, the parental magma exhibited minimal saturation with Fe-Ti oxides. As olivine and plagioclase crystallized, pyroxene and Fe-Ti oxides gradually enriched the residual magma. Exsolution microcrystals of Fe-Ti oxides became visible within the pyroxene at this stage. In the final phase, Fe-Ti oxides, further enriched in the residual magma, crystallized as separate pyroxene and Fe-Ti oxide phases, culminating in the formation of pyroxene-gabbroic Fe-Ti oxide deposits.

3.3. Apatite as a Useful Petrogenetic Indicator

The morphological characteristics of apatite can act as indicators of its genesis. This study systematically investigated the morphological features of apatite in the mafic-ultramafic intrusive rocks of Wajilitag. Based on the crystal morphology and occurrence, apatite can be classified into three types. The first type, referred to herein as the inclusion phase, is enclosed by pyroxene and plagioclase and exhibits a long prismatic shape with aspect ratios ranging from 3:1 to 10:1. The second type, termed the intergranular phase, grows intergranularly among silicate minerals, exhibiting an automorphic-semi-automorphic habit, and is frequently associated with Fe-Ti oxides. The third type, designated as the late-stage phase, displays a semi-automorphic to heteromorphic habit and is associated with late-stage biotite or amphibole.

4. Samples and Methods

This study selected 26 core samples from the Wajilitag intrusion mass for mineralogical and geochemical analysis. The lithologies of the samples include 6 gabbros, 9 clinopyroxenites, and 11 olivine pyroxenites. Among these, 14 samples were collected from the drilled core ZK4506, and 12 from the drilled core ZK3306. The locations for the drilled core of the Wajilitag intrusion are shown in Figure 2, and photographs of the samples are presented in Figure 3. The sampling depth is adjusted according to changes in lithology in the borehole to ensure lithological integrity, and the samples are immediately sealed on-site to prevent contamination.

The major element composition analysis of apatite was conducted using electron probe microanalysis (EPMA) at the Key Laboratory of Ore Genesis and Resource Evaluation, Ministry of Natural Resources, Institute of Mineral Resources, Chinese Academy of Geological Sciences. The analysis employed a JEOL JXA-8230 (JEOL Ltd., Tokyo, Japan) electron probe microanalyzer equipped with a four-channel spectrometer. Before testing, all sample surfaces were coated with a 20 nm thick carbon film to ensure conductivity. The specific analytical workflow followed the methodology outlined in [27]. The analytical conditions were set as follows: acceleration voltage of 15 kV, beam current of 20 nA, and beam spot diameter of 5 μm. Reference standards consisted of natural minerals and synthetic oxides, and the specific analytical conditions and parameters are provided in

Table A1. The analysis conditions and parameters of apatite by EMPA.

Element	Crystal	Spectral Line	Counting Time (s)		Standard Reference Material	Detection Limit (ppm)
			Peak	Background
Na	TAP	Kα	10	5	Jadeite	71–87
Mg	TAP	Kα	10	5	Forsterite	65–74
Al	TAP	Kα	10	5	Jadeite	77–81
F	TAP	Kα	10	5	Topaz	1098–1174
K	PETJ	Kα	10	5	K-feldspar	48–53
Ca	PETJ	Kα	10	5	Wollastonite	64–77
Si	PETJ	Kα	10	5	Jadeite	83–96
Fe	LIF	Kα	10	5	Hematite	87–99
Ti	LIF	Kα	10	5	Rutile	236–244
P	PETJ	Kα	10	5	Apatite	84–93
Cr	LIFH	Kα	10	5	Cr2O3	112–137
Mn	LIFH	Kα	10	5	MnO	104–113
Ni	LIF	Kα	10	5	NiO	136–154
V	LIFH	Kα	10	5	V2O5	104–192
Cl	PETH	Kα	10	5	NaCl	222–289

. All acquired data underwent ZAF correction to ensure the accuracy of quantitative results.

Micro-area elemental analysis of apatite was performed at the Key Laboratory of Ore Genesis and Resource Evaluation, Ministry of Natural Resources, Institute of Mineral Resources, Chinese Academy of Geological Sciences, using laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS). The analysis employed a RESOlution S-155 (Applied Spectra Inc., Sacramento, CA, USA) 193 nm excimer laser ablation system coupled with a Bruker M90 ICP-MS (Bruker Corporation., Billerica, MA, USA). During ablation, helium was used as the carrier gas, with argon introduced as a compensating gas to regulate sensitivity. Each time-resolved analysis consisted of approximately 15 s of blank signal and 45 s of sample signal. The detailed instrument operating conditions are summarized in

Table A2. ICP-MS and laser ablation system operational settings.

ICP-MS Instrument Parameters
Instrument Model	Bruker M90
Sampling Depth (mm)	6
Carrier Gas (Ar) Flow Rate (L/min)	1.08
Signal Acquisition Mode	Time Resolution
Laser Ablation System Parameters
Instrument Model	RESOlution S-155
Carrier Gas & Flow Rate (L/min)	He: 0.6
Energy Density (J/cm2)	6
Laser Spot Size (μm)	30/38
Repetition Rate (Hz)	5

. Quantitative calculations used multiple external standards, no internal standard method, employing USGS reference glasses (BCR-2G, BIR-1G, GSE-1G) as calibration standards, with recommended values for each element referenced from the GeoReM database (http://georem.mpch-mainz.gwdg.de/, accessed on 22 April 2024). Offline processing of all data—including sample/blank signal selection, instrument sensitivity drift correction, and elemental content calculation—was conducted using ICPMSDataCal software 9.1.

5. Results

5.1. Major Elements

The major element analysis in the form of oxides results for apatite are summarized in

Table A3. Apatite major oxide composition in the Wajilitag Intrusion (wt%).

Sample	F	Cl	MgO	Al2O3	Na2O	P2O5	CaO	SiO2	K2O	MnO	FeO	TiO	SO3	Total
Inclusion Phase
ZK4506
b1-2-1	1.98	0.29	0.06	0.03	0.04	41.65	54.77	0.57	0.02	0.00	0.25	0.04	0.07	99.77
b1-3-1	2.08	0.25	0.05	0.00	0.09	41.58	54.57	0.40	0.01	0.08	0.21	0.00	0.03	99.34
b1-3-2	2.25	0.31	0.01	0.00	0.13	41.60	54.24	0.39	0.00	0.02	0.19	0.00	-	99.13
b3-3-1	1.50	0.78	0.04	0.04	0.04	41.57	55.13	0.18	0.00	0.03	0.23	0.02	0.01	99.58
b3-4-1	1.60	0.32	0.03	0.01	0.04	41.40	56.50	0.29	0.00	0.08	0.10	0.00	0.01	100.39
b9-3-1	2.02	0.87	0.04	0.00	0.12	41.74	53.48	0.33	0.00	0.00	0.52	0.00	0.04	99.15
b9-3-2	1.85	0.97	0.03	0.00	0.10	41.39	53.65	0.37	0.00	0.04	0.70	0.00	0.04	99.12
b9-3-3	1.61	0.89	0.05	0.00	0.06	41.58	53.91	0.45	0.02	0.09	0.74	0.00	-	99.40
b13-4-1	1.73	0.81	0.00	0.01	0.03	41.54	54.83	0.17	0.01	0.06	0.17	0.00	-	99.37
b18-1-1	1.40	0.46	0.32	0.00	0.03	41.85	54.47	0.55	0.01	0.18	0.45	0.03	-	99.76
b18-1-2	2.02	0.49	0.05	0.02	0.07	41.42	54.48	0.36	0.02	0.08	0.39	0.04	0.05	99.49
b18-2-1	2.37	0.43	0.03	0.00	0.06	41.38	54.99	0.32	0.03	0.07	0.08	0.00	0.02	99.77
b18-2-2	2.03	0.41	0.05	0.00	0.08	41.68	54.43	0.36	0.05	0.06	0.13	0.01	0.03	99.32
b18-3-1	2.40	0.47	0.02	0.01	0.02	41.65	54.76	0.41	0.02	0.02	0.11	0.00	0.02	99.89
b18-3-2	1.99	0.45	0.00	0.00	0.05	41.46	55.03	0.63	0.00	0.11	0.23	0.00	0.07	100.03
b19-1-1	1.60	0.77	0.02	0.00	0.16	41.31	55.60	0.19	0.00	0.05	0.08	0.00	0.03	99.82
b19-2-1	2.02	0.69	0.02	0.00	0.13	41.60	54.33	0.33	0.00	0.04	0.14	0.00	0.04	99.33
b19-2-2	1.88	0.21	0.01	0.00	0.13	41.70	54.67	0.42	0.04	0.00	0.18	0.00	-	99.22
b19-2-3	1.94	0.65	0.01	0.02	0.10	41.55	54.79	0.31	0.02	0.00	0.14	0.01	0.01	99.54
b19-3-1	2.08	0.59	0.02	0.00	0.15	41.66	54.70	0.23	0.01	0.00	0.13	0.00	0.03	99.60
b19-4-1	1.40	0.81	0.05	0.00	0.26	41.41	54.93	0.22	0.01	0.03	0.07	0.00	0.02	99.19
b19-4-2	1.99	0.42	0.00	0.00	0.11	41.47	54.84	0.25	0.00	0.04	0.12	0.00	-	99.24
b20-1-1	1.57	0.13	0.09	0.02	0.04	42.16	55.77	0.26	0.02	0.08	0.34	0.08	0.01	100.56
b20-1-2	1.71	0.13	0.03	0.00	0.03	41.60	55.13	0.25	0.01	0.04	0.26	0.02	0.03	99.24
b22-1-1	1.96	0.65	0.04	0.02	0.08	41.53	54.55	0.31	0.00	0.08	0.22	0.00	0.02	99.46
b22-1-2	1.82	0.66	0.04	0.00	0.06	41.33	55.18	0.43	0.01	0.02	0.26	0.00	0.13	99.84
b22-1-3	1.94	0.53	0.02	0.01	0.10	41.60	54.74	0.19	0.00	0.03	0.19	0.02	0.11	99.48
b22-2-1	1.66	0.59	0.05	0.01	0.05	41.56	55.51	0.36	0.00	0.09	0.15	0.00	0.04	100.06
b22-2-2	1.62	0.55	0.05	0.02	0.07	41.73	54.69	0.20	0.00	0.11	0.21	0.00	0.02	99.27
b22-2-3	1.67	0.64	0.02	0.03	0.04	41.31	55.16	0.50	0.00	0.02	0.16	0.00	0.14	99.70
b22-3-1	1.66	0.64	0.03	0.00	0.06	41.42	55.42	0.26	0.00	0.06	0.12	0.00	0.09	99.76
b22-3-2	1.93	0.64	0.05	0.00	0.05	41.42	54.99	0.32	0.01	0.04	0.08	0.00	0.06	99.58
b22-3-3	1.44	0.54	0.08	0.00	0.07	41.51	53.35	0.24	0.00	0.07	0.10	0.00	0.08	97.48
intergranular phase
ZK3306
b3-1-01	1.76	0.71	0.09	0.01	0.03	41.63	55.22	0.22	0.01	0.11	0.31	0.02		100.11
b3-1-02	1.59	0.79	0.11	0.02	0.14	41.62	55.53	0.38	0.03	0.07	0.32	0.01		100.60
b5-1-01	1.74	0.34	0.05	0.00	0.06	41.47	55.31	0.14	0.00	0.04	0.42	0.04		99.60
b5-1-02	1.46	0.39	0.10	0.01	0.05	41.70	55.10	0.17	0.00	0.06	0.12	0.00		99.14
b10-1-01	1.47	0.60	0.00	0.00	0.01	41.79	54.97	0.13	0.01	0.00	0.31	0.00		99.28
b10-1-02	1.64	0.52	0.10	0.02	0.03	41.45	54.77	0.22	0.00	0.08	0.27	0.00		99.11
b10-1-03	1.34	0.59	0.25	0.04	0.08	41.38	55.30	0.30	0.04	0.00	0.48	0.05		99.86
b11-1-01	1.31	0.48	0.10	0.00	0.08	41.57	55.52	0.18	0.01	0.00	0.28	0.03	0.08	99.66
b11-2-01	1.73	0.31	0.08	0.01	0.06	41.31	55.42	0.15	0.01	0.00	0.32	0.00	0.06	99.45
b11-3-01	1.72	0.37	0.16	0.03	0.10	42.09	54.40	0.28	0.01	0.04	0.41	0.03	0.10	99.75
b11-4-01	1.54	0.39	0.10	0.01	0.02	41.63	55.00	0.25	0.00	0.01	0.42	0.01		99.38
b14-1-01	1.33	0.50	0.04	0.00	0.02	41.43	55.00	0.26	0.00	0.00	0.33	0.00		98.91
b14-1-02	2.66	0.28	0.11	0.00	0.01	41.72	55.55	0.25	0.00	0.06	0.22	0.03		100.89
b14-2-01	1.59	0.47	0.12	0.03	0.04	41.55	54.90	0.21	0.02	0.06	0.39	0.02		99.38
b14-3-01	1.23	0.53	0.10	0.03	0.03	41.42	55.42	0.38	0.01	0.03	0.25	0.00		99.43
b24-1-01	1.30	0.61	0.06	0.02	0.03	41.49	55.88	0.23	0.00	0.00	0.39	0.00		100.01
b24-2-01	2.24	0.34	0.73	0.11	0.07	41.43	52.84	0.91	0.02	0.02	1.54	0.14		100.38
b24-3-01	1.70	0.64	0.12	0.03	0.12	41.55	55.15	0.11	0.01	0.00	0.36	0.00		99.79
b34-1-01	1.83	0.38	0.48	0.07	0.30	41.81	54.26	0.20	0.07	0.05	0.35	0.02		99.81
b34-1-02	1.52	0.43	0.25	0.04	0.17	41.77	54.80	0.18	0.03	0.04	0.25	0.02		99.50
b34-2-01	1.85	0.43	0.29	0.06	0.10	41.50	53.16	0.33	0.04	0.00	0.39	0.00		98.13
b34-2-02	2.05	0.37	0.30	0.05	0.43	41.44	53.64	0.19	0.10	0.02	0.34	0.02		98.94
b40-2-01	1.70	0.38	0.10	0.00	0.02	42.06	55.08	0.12	0.01	0.00	0.16	0.00		99.63
b40-2-02	1.54	0.39	0.07	0.00	0.03	41.73	54.99	0.16	0.00	0.05	0.18	0.00		99.14
b40-3-01	1.24	0.40	0.19	0.10	0.11	41.70	54.56	0.36	0.02	0.00	0.28	0.08		99.04
b42-1-01	1.96	0.39	0.06	0.00	0.05	41.76	55.01	0.12	0.01	0.00	0.46	0.05		99.88
b42-1-02	1.43	0.34	0.89	0.13	0.00	41.78	52.54	1.51	0.00	0.02	0.51	0.02		99.17
b42-1-03	1.73	0.37	0.04	0.01	0.00	41.38	55.34	0.22	0.00	0.00	0.23	0.06		99.38
b42-2-01	1.51	0.33	0.03	0.00	0.03	41.49	55.31	0.12	0.00	0.00	0.23	0.06		99.11
b42-2-03	1.54	0.33	0.06	0.00	0.05	41.47	55.14	0.19	0.00	0.01	0.30	0.00		99.09
b42-3-01	2.18	0.35	0.04	0.00	0.10	42.03	55.48	0.14	0.01	0.03	0.31	0.00		100.67
b46-1-01	1.07	0.31	0.20	0.11	0.06	41.74	54.96	0.91	0.01	0.05	0.29	0.10		99.81
b46-2-01	1.90	0.19	0.05	0.00	0.09	41.62	54.58	0.29	0.00	0.08	0.37	0.02		99.21
b46-2-02	1.33	0.21	0.06	0.00	0.02	41.37	55.35	0.28	0.02	0.00	0.29	0.03		98.94
b46-2-03	1.37	0.20	0.03	0.10	0.08	41.52	55.57	0.54	0.00	0.05	0.07	0.00		99.51
b46-3-01	1.66	0.29	0.14	0.31	0.06	41.43	54.15	0.71	0.00	0.03	1.15	0.02		99.95
b46-3-02	1.56	0.34	0.04	0.03	0.13	41.56	55.03	0.19	0.03	0.00	0.10	0.03		99.05
b48-1-01	1.56	0.90	0.11	0.00	0.05	41.47	55.15	0.20	0.02	0.08	0.26	0.02		99.81
b48-2-01	1.45	0.51	0.30	0.01	0.07	41.52	54.39	0.54	0.01	0.08	0.29	0.02		99.19
b48-2-04	1.54	0.36	0.06	0.02	0.09	41.43	55.83	0.28	0.01	0.00	0.22	0.00		99.83
b48-3-01	1.54	0.63	0.10	0.00	0.05	41.66	54.98	0.21	0.00	0.01	0.23	0.00		99.43
b48-3-02	1.62	0.59	0.03	0.00	0.04	41.47	55.06	0.13	0.00	0.00	0.12	0.02		99.07
b54-1-02	1.39	0.67	0.08	0.03	0.09	41.43	55.05	0.26	0.00	0.12	0.31	0.04		99.45
b54-2-01	1.58	0.78	0.11	0.03	0.05	41.39	54.99	0.13	0.00	0.09	0.09	0.00		99.24
b54-2-02	1.32	0.67	0.09	0.00	0.07	41.76	54.87	0.07	0.01	0.04	0.15	0.02		99.07
b54-3-01	1.41	0.17	0.10	0.01	0.09	41.57	55.40	0.26	0.02	0.01	0.29	0.00		99.34
b54-3-02	1.70	0.30	0.07	0.01	0.13	41.63	55.05	0.09	0.01	0.06	0.29	0.05		99.40
b54-4-01	1.46	0.65	0.12	0.00	0.00	41.33	55.51	0.24	0.01	0.07	0.29	0.00		99.68
ZK4506
b3-1-1	1.89	0.54	0.03	0.01	0.01	41.69	54.77	0.19	0.02	0.00	0.14	0.00	0.02	99.30
b3-1-2	1.93	0.51	0.02	0.01	0.08	41.59	55.32	0.25	0.00	0.00	0.08	0.02	0.06	99.86
b3-1-3	2.03	0.52	0.06	0.00	0.10	41.79	54.69	0.09	0.00	0.05	0.04	0.06	0.07	99.50
b3-1-4	2.01	0.42	0.03	0.01	0.08	41.76	55.47	0.16	0.00	0.02	0.12	0.06	0.01	100.14
b5-1-1	1.56	0.44	0.06	0.00	0.03	41.61	54.98	0.25	0.01	0.01	0.29	0.00	0.02	99.23
b5-1-2	1.98	0.38	0.04	0.01	0.07	41.67	55.08	0.12	0.00	0.04	0.92	0.19	-	100.50
b5-2-1	1.94	0.33	0.01	0.00	0.01	41.40	55.19	0.17	0.00	0.02	0.17	0.00	-	99.25
b5-3-1	1.75	0.28	0.05	0.01	0.00	41.58	55.62	0.26	0.00	0.00	0.27	0.04	0.04	99.89
b5-3-2	1.99	0.22	0.03	0.00	0.01	41.77	54.85	0.08	0.02	0.00	0.19	0.01	0.02	99.17
b5-3-3	1.82	0.17	0.01	0.00	0.04	41.36	55.67	0.25	0.00	0.03	0.17	0.03	-	99.54
b5-4-1	1.88	0.22	0.06	0.00	0.06	41.61	56.48	0.23	0.00	0.02	0.14	0.01	0.03	100.73
b5-4-2	1.71	0.23	0.04	0.00	0.02	41.55	54.98	0.21	0.02	0.02	0.16	0.03	0.004	98.98
b5-4-3	1.61	0.22	0.05	0.00	0.00	41.79	55.43	0.17	0.00	0.05	0.14	0.04	-	99.50
b5-4-4	1.80	0.23	0.00	0.00	0.01	41.31	55.38	0.21	0.00	0.08	0.17	0.10	-	99.30
b6-1-1	2.00	0.25	0.05	0.02	0.07	41.42	55.33	0.24	0.00	0.00	0.19	0.00	0.05	99.62
b6-1-2	1.88	0.28	0.00	0.00	0.00	41.58	55.64	0.26	0.00	0.08	0.16	0.00	0.01	99.90
b6-1-3	1.25	0.30	0.03	0.02	0.01	41.72	55.57	0.24	0.00	0.06	0.39	0.07	-	99.66
b6-2-1	1.85	0.39	0.05	0.01	0.00	41.51	56.09	0.26	0.00	0.04	0.19	0.01	0.004	100.40
b6-3-1	1.70	0.64	0.04	0.00	0.02	41.40	54.78	0.19	0.02	0.07	0.06	0.00	0.01	98.92
b6-4-1	1.54	0.18	0.04	0.01	0.04	41.74	55.75	0.20	0.02	0.07	0.14	0.00	0.002	99.74
b7-3-1	1.54	0.55	0.12	0.02	0.02	41.43	54.98	0.47	0.01	0.01	0.38	0.06	-	99.61
b10-1-1	1.61	0.45	0.11	0.02	0.06	41.76	55.66	0.23	0.00	0.05	0.19	0.00	0.13	100.27
b10-1-2	1.55	0.44	0.10	0.00	0.03	41.71	54.79	0.31	0.00	0.08	0.26	0.00	0.04	99.30
b10-2-1	1.26	0.58	0.08	0.03	0.04	41.59	54.84	0.31	0.00	0.13	0.19	0.02	0.02	99.09
b10-2-2	1.37	0.52	0.07	0.03	0.05	41.59	55.13	0.20	0.00	0.00	0.14	0.02	0.02	99.12
b10-2-3	1.95	0.53	0.06	0.02	0.01	41.78	54.65	0.17	0.01	0.00	0.11	0.00	0.06	99.35
b12-3-1	1.60	0.47	0.11	0.00	0.08	41.42	55.75	0.19	0.00	0.10	0.28	0.00	0.04	100.04
b12-3-2	1.24	0.58	0.12	0.00	0.06	41.39	55.43	0.18	0.02	0.05	0.26	0.03	0.01	99.37
b25-2-1	1.93	0.24	0.08	0.00	0.06	41.72	55.17	0.15	0.00	0.01	0.12	0.00	0.04	99.53
b25-2-2	1.61	0.23	0.06	0.01	0.04	41.77	55.55	0.18	0.01	0.07	0.18	0.00	0.02	99.72
b25-4-1	2.07	0.24	0.03	0.00	0.09	41.76	54.85	0.25	0.03	0.00	0.06	0.00	0.02	99.39
b25-4-2	1.40	0.27	0.05	0.00	0.06	41.64	54.85	0.19	0.00	0.03	0.16	0.00	0.002	98.65
b25-4-3	1.67	0.25	0.05	0.00	0.06	41.77	55.16	0.21	0.01	0.00	0.14	0.00	0.03	99.33
b25-4-4	1.53	0.51	0.04	0.01	0.04	41.64	55.13	0.28	0.00	0.00	0.16	0.00	-	99.34
late-stage phase
ZK4506
b3-2-1	1.65	1.16	0.07	0.00	0.01	41.75	54.07	0.39	0.00	0.09	0.24	0.03	0.04	99.51
b7-1-1	1.56	1.12	0.02	0.00	0.09	41.41	54.54	0.20	0.00	0.07	0.25	0.00	0.05	99.31
b7-1-2	1.98	1.12	0.02	0.01	0.12	41.46	54.98	0.26	0.00	0.12	0.14	0.00	0.00	100.21
b7-1-3	1.50	1.26	0.04	0.00	0.10	41.45	54.24	0.25	0.00	0.00	0.23	0.02	0.03	99.12
b7-2-1	1.29	1.16	0.04	0.00	0.06	41.61	54.53	0.22	0.01	0.00	0.20	0.00	0.03	99.15
b9-1-1	1.63	1.02	0.03	0.00	0.09	41.70	54.49	0.54	0.03	0.00	0.10	0.02	0.04	99.68
b9-1-2	1.91	1.06	0.03	0.02	0.10	41.44	54.30	0.55	0.02	0.03	0.14	0.03	0.08	99.72
b9-2-1	1.51	1.22	0.04	0.01	0.15	41.77	55.03	0.58	0.00	0.00	0.12	0.00	0.12	100.66
b9-2-2	1.65	1.23	0.17	0.17	0.16	41.33	53.42	1.02	0.32	0.10	0.65	0.13	0.16	100.50
b10-3-1	1.47	0.91	0.10	0.01	0.09	41.55	54.77	0.14	0.00	0.05	0.19	0.01	0.06	99.35
b10-3-2	1.60	1.01	0.07	0.00	0.07	41.66	54.55	0.09	0.00	0.08	0.22	0.00	-	99.35
b12-1-1	1.02	2.33	0.01	0.02	0.11	41.35	55.07	0.16	0.00	0.00	0.07	0.00	0.05	100.18
b12-1-2	1.06	2.05	0.03	0.00	0.08	41.52	54.49	0.16	0.00	0.00	0.08	0.00	0.02	99.50
b12-1-3	0.65	2.39	0.05	0.00	0.07	41.32	55.64	0.13	0.01	0.00	0.05	0.00	-	100.30
b12-2-1	1.14	1.80	0.03	0.00	0.09	41.73	54.51	0.12	0.01	0.00	0.18	0.00	0.02	99.62
b12-2-2	1.30	1.84	0.06	0.00	0.06	41.59	54.09	0.25	0.02	0.00	0.15	0.03	0.04	99.38
b12-2-3	1.04	1.95	0.09	0.00	0.08	41.66	53.96	0.22	0.02	0.05	0.21	0.02	-	99.29
b13-1-1	1.34	0.91	0.03	0.03	0.08	41.90	55.00	0.21	0.01	0.07	0.15	0.00	0.09	99.81
b13-2-1	2.00	1.06	0.02	0.04	0.09	41.65	53.55	0.48	0.01	0.00	0.24	0.02	0.05	99.21
b13-3-1	1.21	1.02	0.03	0.02	0.09	41.76	54.15	0.72	0.04	0.11	0.19	0.00	0.09	99.43
b16-1-1	1.21	1.30	0.02	0.00	0.01	41.61	54.69	0.24	0.00	0.10	0.26	0.00	-	99.43
b16-2-1	0.65	2.96	0.04	0.00	0.07	41.62	53.47	0.24	0.01	0.04	0.17	0.00	-	99.27
b16-3-1	1.20	1.47	0.06	0.00	0.06	41.76	54.46	0.12	0.00	0.00	0.19	0.02	-	99.34
b16-3-2	1.34	1.62	0.03	0.00	0.06	41.62	54.31	0.12	0.00	0.11	0.23	0.00	0.01	99.45
b16-3-3	1.31	1.71	0.05	0.00	0.04	41.45	54.05	0.16	0.00	0.08	0.19	0.06	0.04	99.14
b20-2-1	1.54	1.06	0.03	0.00	0.09	41.66	54.70	0.46	0.06	0.11	0.38	0.01	-	100.11
b20-2-2	1.14	1.02	0.15	0.08	0.12	41.56	53.34	0.95	0.30	0.11	0.72	0.13	0.002	99.61
b20-3-1	1.58	1.12	0.01	0.02	0.10	41.50	53.97	0.42	0.06	0.05	0.25	0.04	0.002	99.13
b25-1-1	1.22	1.53	0.02	0.00	0.07	41.76	55.00	0.32	0.01	0.02	0.06	0.00	0.06	100.07
b25-1-2	1.67	1.49	0.03	0.01	0.06	41.77	54.38	0.22	0.02	0.01	0.10	0.00	-	99.75
b25-1-3	1.40	1.51	0.04	0.00	0.04	41.75	54.53	0.25	0.00	0.04	0.06	0.00	0.03	99.65
b25-3-1	1.44	1.35	0.04	0.01	0.06	41.57	54.46	0.19	0.00	0.00	0.14	0.01	-	99.26
b25-3-2	1.31	1.36	0.01	0.00	0.05	41.44	55.25	0.17	0.00	0.11	0.11	0.01	-	99.81

Note: Below the Detection Limit.

. The fluoride (F) content in the inclusion-phase, intergranular-phase, and late-stage apatite ranges from 1.40% to 2.40% (mean 1.84%), 1.07% to 2.66% (mean 1.65%), and 0.65% to 2.00% (mean 1.38%), respectively. Similarly, the chlorine (Cl) content varies between 0.13% and 0.97% (mean 0.55%), 0.17% to 0.90% (mean 0.42%), and 0.91% to 2.96% (mean 1.43%) in the respective apatite types. The contents of CaO, P₂O₅, SiO₂, and FeO (the measured Fe is a total iron content as ferrous) show little variation across the three apatite phases, indicating that late-stage hydrothermal alteration had minimal impact on these components. The Ca/P molar ratios for all apatites range from 2.06 to 2.23 (mean 2.16), a value close to the theoretical magmatic value of 1.67 [28], suggesting a magmatic origin. A progressive decrease in fluoride (F) content and an increase in chlorine (Cl) content were observed from the inclusion phase to the late-stage apatite. The inclusion-phase apatite displayed the highest fluoride content and relatively low chlorine levels. In contrast, the intergranular-phase apatite showed fluoride and chlorine levels comparable to the inclusion-phase, but with slightly lower fluoride and the lowest chlorine. The late-stage phase apatite had the lowest fluoride content and the highest chlorine content.

The major element analysis results for magnetite and ilmenite are presented in

Table A4. Magnetite-Ilmenite Main Element Composition in the Wajilitag Intrusion (wt%).

Sample	SiO2	TiO2	Al2O3	FeO	MnO	MgO	CaO	K2O	V2O3	Cr2O3	ZnO	NiO	CuO	CoO	Total
ZK4506
b1-Mt-1	0.02	1.11	0.28	92.29	0.05	0.13	0.00	0.00	0.25	0.12	0.06	0.00	0.05	0.17	94.51
b1-Ilm-1	0.00	47.81	0.03	50.40	1.26	0.27	0.00	0.01	0.47	0.00	0.03	0.00	0.00	0.07	100.34
b1-Mt-2	0.00	0.58	0.14	92.94	0.04	0.01	0.00	0.00	0.24	0.05	0.00	0.00	0.00	0.17	94.17
b1-Ilm-2	0.05	46.11	0.03	51.24	1.47	0.21	0.00	0.00	0.68	0.00	0.12	0.00	0.00	0.06	99.98
b5-Mt-1	0.02	1.59	0.48	92.16	0.10	0.48	0.00	0.02	0.41	0.01	0.00	0.00	0.00	0.15	95.40
b5-Ilm-1	0.03	52.68	0.00	41.24	0.81	4.58	0.11	0.01	0.62	0.06	0.02	0.00	0.00	0.09	100.23
b5-Mt-2	0.01	2.30	6.29	75.57	0.16	2.94	0.02	0.00	0.47	0.08	0.10	0.03	0.06	0.22	88.23
b5-Ilm-2	0.00	51.27	0.01	42.91	0.62	4.06	0.01	0.01	0.44	0.00	0.03	0.02	0.00	0.09	99.47
b6-Mt-1	0.01	5.65	1.58	82.21	0.10	1.03	0.00	0.00	0.44	0.20	0.05	0.00	0.08	0.22	91.57
b6-Ilm-1	0.00	49.95	0.00	41.27	0.59	3.60	0.00	0.00	0.61	0.00	0.04	0.00	0.00	0.11	96.16
b6-Mt-2	0.00	7.19	1.42	80.90	0.19	0.75	0.01	0.00	0.43	0.12	0.05	0.02	0.02	0.23	91.33
b6-Ilm-2	0.02	50.82	0.02	41.49	0.54	3.63	0.00	0.00	0.60	0.00	0.00	0.00	0.00	0.04	97.16
b9-Mt-1	0.00	1.01	0.24	91.63	0.00	0.05	0.06	0.00	0.38	0.08	0.10	0.00	0.14	0.19	93.87
b9-Ilm-1	0.01	48.17	0.02	47.19	0.99	0.93	0.12	0.01	0.55	0.06	0.11	0.00	0.00	0.13	98.28
b9-Mt-2	0.00	0.96	0.34	91.62	0.04	0.12	0.00	0.00	0.30	0.04	0.00	0.00	0.00	0.16	93.58
b9-Ilm-2	0.00	48.42	0.03	46.32	1.68	1.03	0.00	0.00	0.50	0.00	0.00	0.00	0.00	0.14	98.12
b10-Mt-1	0.00	1.60	2.00	87.47	0.04	0.84	0.00	0.01	0.43	0.21	0.13	0.04	0.04	0.16	92.98
b10-Ilm-1	0.02	51.29	0.03	44.00	0.95	2.46	0.00	0.00	0.54	0.08	0.15	0.00	0.00	0.10	99.62
b10-Mt-2	0.00	1.23	0.43	91.78	0.03	0.15	0.00	0.00	0.48	0.18	0.00	0.00	0.00	0.16	94.46
b10-Ilm-2	0.00	50.74	0.03	46.37	0.82	1.76	0.00	0.00	0.70	0.03	0.00	0.00	0.10	0.06	100.61
b22-Mt-1	0.02	1.34	0.54	92.06	0.10	0.15	0.00	0.01	0.54	0.00	0.01	0.00	0.00	0.17	94.93
b22-Ilm-1	0.00	49.40	0.02	46.15	1.63	1.47	0.02	0.01	0.54	0.00	0.05	0.00	0.00	0.12	99.42
b22-Mt-2	0.00	2.24	0.66	90.00	0.25	0.27	0.00	0.00	0.42	0.00	0.11	0.00	0.00	0.17	94.13
b22-Ilm-2	0.02	48.03	0.02	47.07	1.62	1.35	0.00	0.02	0.62	0.00	0.03	0.00	0.00	0.10	98.86
b25-Mt-1	0.07	3.25	0.43	90.22	0.22	0.23	0.00	0.00	0.50	0.14	0.06	0.00	0.00	0.19	95.30
b25-Ilm-1	0.03	49.35	0.04	46.19	1.32	1.20	0.00	0.00	0.41	0.04	0.00	0.00	0.04	0.11	98.73
b25-Mt-2	0.00	7.09	1.97	81.37	0.27	0.96	0.00	0.00	0.46	0.08	0.03	0.00	0.00	0.18	92.42
b25-Ilm-2	0.02	49.74	0.02	46.37	1.41	1.59	0.00	0.00	0.54	0.02	0.12	0.02	0.00	0.12	99.97
ZK3306
b34-Mt-1	0.03	10.45	1.93	75.99	0.30	1.83	0.00	0.00	0.43	0.29	0.00	0.00	0.06	0.19	91.50
b34-Ilm-1	0.03	50.95	0.04	39.47	0.46	4.38	0.00	0.00	0.41	0.03	0.14	0.00	0.01	0.08	96.01
b34-Mt-2	0.01	10.56	1.97	73.85	0.37	2.09	0.04	0.00	0.34	0.31	0.01	0.00	0.00	0.08	89.63
b34-Ilm-2	0.00	50.33	0.04	37.94	0.43	5.28	0.03	0.01	0.43	0.00	0.07	0.00	0.00	0.03	94.60
b40-Mt-1	0.01	10.08	1.75	78.37	0.25	1.85	0.00	0.01	0.52	0.32	0.00	0.08	0.00	0.17	93.42
b40-Ilm-1	0.00	52.44	0.02	38.97	0.49	4.92	0.00	0.00	0.63	0.00	0.06	0.00	0.00	0.02	97.54
b40-Mt-2	0.00	10.37	2.08	77.16	0.24	1.89	0.00	0.01	0.48	0.34	0.11	0.05	0.00	0.14	92.86
b40-Ilm-2	0.00	52.37	0.01	39.33	0.46	5.28	0.00	0.00	0.59	0.07	0.14	0.00	0.00	0.10	98.34
b48-Mt-1	0.02	8.59	2.11	76.21	0.34	1.30	0.00	0.00	0.36	0.41	0.03	0.00	0.00	0.17	89.53
b48-Ilm-1	0.01	47.62	0.01	40.68	0.46	3.70	0.03	0.00	0.47	0.04	0.00	0.00	0.00	0.01	93.03
b48-Mt-2	0.03	6.89	1.40	79.78	0.21	0.79	0.00	0.01	0.45	0.51	0.10	0.06	0.00	0.17	90.41
b48-Ilm-2	0.02	48.62	0.00	41.27	0.51	3.08	0.00	0.04	0.46	0.03	0.00	0.00	0.00	0.08	94.11

. The titanium dioxide (TiO₂) content in magnetite from the Wajilitag intrusive body shows considerable variation, with a maximum value of 10.56 wt%. The FeO^t content ranges from 73.85 to 92.29 wt%. Magnetite also contains 0.01–2.94 wt% MgO, 0.14–6.29 wt% Al₂O₃, 0.01–0.51 wt% Cr₂O₃, and 0.24–0.54 wt% V₂O₃. In comparison, ilmenite exhibits TiO₂ content ranging from 46.11 to 52.68 wt%, with MgO varying between 0.21 and 5.28 wt%. The FeO^t content in ilmenite ranges from 37.94 to 51.24 wt%, typically showing an inverse correlation with TiO₂. The MnO content ranges from 0.43 to 1.68.

5.2. Trace Elements

The trace element analysis of apatite was summarized in

Table A5. Apatite Trace Element Composition in the Wajilitag Intrusion (ppm).

Inclusion Phase
ZK4506
Sample		b1-2-1			b9-3-1					b9-3-2				b19-1-1				b19-2-1				b19-2-2				b19-2-3				b19-3-1				b19-4-1				b22-2-1			b22-2-2			b22-3-2		b22-3-3			b22-3-4
Rb		0.63			1.33					0.94				0.55				0.74				0.61				0.86				0.67				0.74				0.31			0.30			0.63		0.90			0.73
Ba		1.05			1.21					3.02				4.46				2.63				10.43				3.85				3.52				5.53				1.04			1.98			11.99		125.70			10.81
Th		35.34			267.30					261.43				21.33				118.94				154.62				87.55				49.99				23.43				9.89			16.43			11.79		19.52			21.10
U		-			42.72					36.26				9.42				35.46				38.41				35.18				21.47				7.45				2.81			4.98			2.15		3.94			3.71
Nb		0.13			0.24					-				0.10				0.08				0.08				-				-				0.09				0.02			0.09			0.06		0.09			-
La		1595.40			3063.52					2762.62				2678.24				2814.58				2921.97				2946.87				2488.59				2579.64				958.34			1149.06			1385.72		1180.69			1405.66
Ce		3672.17			6010.18					5405.43				4965.86				5521.65				5692.80				5667.50				4709.69				5149.95				2025.67			2562.05			2787.28		2341.14			2802.60
Pb		1.55			11.39					8.99				13.57				4.49				5.64				3.88				2.75				4.40				3.40			3.11			3.40		6.27			3.62
Pr		439.16			681.46					610.33				553.90				624.64				645.51				626.05				542.56				563.41				235.27			290.92			307.79		256.73			299.01
Sr		1466.44			1083.70					1143.52				998.66				1015.48				1097.86				1027.47				934.29				1106.42				1004.43			1864.66			1055.64		1316.90			859.78
Nd		1828.32			2716.78					2391.50				2184.25				2439.13				2533.72				2523.89				2152.31				2237.52				1047.07			1268.27			1296.09		1058.66			1248.02
Sm		325.09			493.35					446.01				370.46				407.04				437.69				417.98				377.68				381.56				199.24			240.10			233.03		189.59			221.02
Zr		11.10			15.82					9.21				3.19				1.62				-				-				-				9.94				9.98			13.37			13.35		10.61			13.74
Hf		0.00			0.08					-				-				-				-				-				-				0.05				-			0.12			-		-			-
Eu		69.32			101.09					94.21				58.89				70.78				77.70				74.61				63.28				63.82				46.08			55.72			56.66		43.53			50.39
Gd		266.07			414.39					362.11				284.50				312.52				328.74				314.60				294.15				293.21				172.39			195.70			189.91		154.32			182.86
Tb		29.73			48.54					42.08				32.78				36.16				38.38				36.26				35.49				35.71				18.08			22.30			20.50		16.48			20.65
Dy		154.04			240.82					216.31				165.67				189.44				198.66				194.82				182.04				178.91				91.63			113.01			107.92		85.90			99.36
Y		661.71			1045.70					932.12				768.78				845.99				874.00				839.31				811.40				787.08				392.21			470.63			464.69		376.66			442.06
Ho		24.23			39.32					34.72				27.85				30.20				32.84				31.24				30.85				28.63				14.80			18.08			17.31		13.71			17.15
Er		52.35			89.36					78.16				67.93				71.25				77.36				75.39				69.73				64.98				32.83			41.06			38.84		31.10			39.90
Tm		5.95			9.30					8.17				6.25				7.95				8.33				8.05				7.51				7.24				3.19			4.38			4.02		3.08			4.32
Yb		27.06			46.00					37.77				34.06				39.85				41.42				41.14				38.98				37.83				16.55			21.14			20.97		20.13			22.43
Lu		3.24			4.53					3.73				4.14				4.85				4.63				5.22				4.45				4.35				2.08			2.54			2.52		1.93			2.63
Intergranular Phase
ZK4506
Sample	b3-1-3			b5-1-1			b5-2-1				b5-3-1				b5-4-1				b6-1-1				b6-2-1				b6-3-1				b6-4-1				b10-1-1				b10-2-1			b12-3-1			b12-3-2			b25-2-1			b25-4-1
Rb	0.51			1.14			0.56				0.24				0.21				-				296.41				-				-				0.24				0.36			0.17			0.33			0.33			0.32
Ba	3.95			5.54			1.35				0.50				5.64				-				1682.84				6.43				4.35				8.18				3.98			4.79			3.05			1.43			5.93
Th	86.78			417.37			76.98				88.91				17.02				93.73				180.16				110.63				67.71				17.97				38.85			68.03			41.69			44.30			37.12
U	20.48			31.45			30.55				37.86				5.70				28.35				46.40				26.65				20.84				3.17				13.32			22.54			13.95			10.52			9.09
Nb	0.06			0.44			-				-				-				-				12.35				-				-				0.10				-			-			0.08			-			-
La	1263.26			670.20			797.88				669.52				801.87				950.76				1157.64				1128.88				877.56				1346.84				1203.61			1044.98			1094.55			1077.77			1028.37
Ce	2401.12			1335.26			1641.17				1279.03				1694.54				1746.72				2038.54				2007.75				1823.87				2749.28				2470.41			2024.57			2124.22			2197.85			2092.01
Pb	3.10			11.88			2.88				2.76				1.49				10.16				28.83				33.25				249.04				30.03				59.25			39.17			31.90			2.89			2.57
Pr	257.32			159.90			198.66				146.52				200.78				188.09				213.24				222.83				211.43				289.86				261.62			215.22			221.85			235.75			230.68
Sr	1365.74			911.70			1344.92				1576.06				1832.47				1545.63				1129.81				1104.40				1858.16				1344.80				1346.96			1288.36			1324.59			1840.68			2176.67
Nd	989.61			658.65			836.58				623.10				910.12				749.90				821.37				858.07				872.59				1151.25				1066.81			897.84			931.81			952.76			927.58
Sm	176.87			117.01			149.34				115.12				170.10				132.38				140.81				134.68				159.59				187.85				185.34			152.82			164.95			166.49			163.53
Zr	12.83			6.87			9.18				6.56				12.81				8.46				22.73				8.48				9.34				17.27				17.00			14.89			15.30			8.57			6.97
Hf	0.00			0.43			-				-				0.11				-				0.18				-				-				0.12				0.11			-			-			-			-
Eu	42.50			25.13			37.28				28.71				43.35				29.98				35.08				32.18				35.87				45.55				44.37			37.08			38.87			40.75			40.57
Gd	134.15			89.37			128.32				97.13				140.82				99.03				107.17				107.77				132.31				151.23				144.98			122.94			124.10			133.57			131.39
Tb	14.34			8.93			12.97				10.54				15.20				10.55				11.06				11.39				13.59				16.15				16.07			13.39			13.33			15.06			14.45
Dy	69.84			38.35			60.05				49.70				69.74				51.67				54.22				54.09				65.27				79.29				78.94			64.12			69.68			74.24			70.74
Y	301.68			163.33			263.98				209.92				302.95				212.40				205.52				233.86				278.58				328.56				329.35			270.40			287.88			324.59			317.85
Ho	10.63			6.03			9.84				8.26				11.10				7.90				7.89				8.22				10.20				12.46				12.40			10.26			10.07			12.17			12.10
Er	24.15			12.71			21.16				16.95				24.31				17.17				19.60				19.03				22.26				29.80				28.19			22.67			24.12			29.01			27.38
Tm	2.77			1.04			2.08				1.83				2.35				1.85				1.72				1.58				2.28				3.07				2.85			2.29			2.65			2.53			2.57
Yb	11.82			6.53			11.25				7.79				11.79				9.13				9.45				8.71				10.69				14.38				14.82			12.61			12.77			13.79			13.78
Lu	1.52			0.82			1.28				1.06				1.44				1.21				1.09				1.34				1.16				1.91				1.70			1.42			1.73			1.63			1.63
Intergranular Phase
	ZK4506							ZK3306
Sample	b25-4-2			b25-4-3				b10-1-1				b11-1-1				b11-3-1				b11-4-1				b24-2-1				b40-1-1				b40-1-2				b40-3-1			b42-2-1			b42-3-1			b46-1-1			b46-3-1			b46-3-2
Rb	0.31			1.11				0.24				0.25				-				0.36				0.42				0.17				0.47				0.19			0.25			2.56			0.84			2.42			0.69
Ba	5.12			9.08				5.73				2.48				2.16				3.05				1.24				2.58				11.97				2.11			1.48			3.94			4.33			8.78			8.71
Th	42.66			42.06				16.54				75.91				63.47				52.94				191.46				52.21				50.07				70.72			50.91			51.62			35.45			37.42			40.53
U	10.84			10.21				4.06				15.19				11.25				9.79				36.16				19.14				25.06				17.60			13.98			15.04			21.36			25.54			24.22
Nb	-			-				0.02				-				-				0.01				-				-				1.30				-			0.02			0.06			0.12			0.24			0.25
La	952.32			1029.62				867.02				664.95				646.99				724.93				743.16				669.02				607.75				681.89			746.81			841.43			1094.16			1275.09			1263.02
Ce	1945.88			2077.00				1691.86				1469.34				1408.71				1569.80				1523.55				1214.00				1091.44				1255.02			1491.08			1547.55			1625.83			2084.31			2061.48
Pb	4.42			5.99				7.29				15.21				6.91				7.99				17.83				10.22				8.58				12.76			6.57			20.10			3.81			4.16			4.55
Pr	214.69			232.62				183.19				165.88				160.70				180.28				172.27				133.35				120.16				134.45			161.46			166.88			151.99			205.09			200.54
Sr	2096.09			2129.59				1120.35				1356.03				1268.46				1257.35				1159.78				1118.75				1057.62				1076.32			1169.78			1078.10			2312.03			2278.94			2206.64
Nd	891.76			954.24				776.21				690.47				674.31				755.53				689.78				550.36				504.69				563.25			671.23			679.47			559.04			805.57			790.31
Sm	154.70			171.09				134.24				120.11				118.67				135.84				113.20				92.02				85.94				97.85			111.32			115.11			103.11			150.88			148.54
Zr	6.17			7.61				10.48				8.22				9.14				9.73				9.06				6.94				13.07				4.51			3.75			8.55			7.04			11.64			9.66
Hf	-			-				0.06				-				-				-				-				0.06				0.23				-			-			0.13			-			-			-
Eu	37.35			43.77				34.24				29.77				31.62				35.03				29.40				24.26				20.81				25.06			27.38			28.02			36.88			49.14			46.70
Gd	126.04			134.48				115.10				99.50				101.99				117.19				95.16				76.03				70.59				87.23			95.66			97.72			96.23			139.10			139.15
Tb	14.23			15.63				12.50				11.05				10.73				13.01				10.11				8.70				7.51				9.09			9.82			10.05			11.67			16.47			16.46
Dy	69.18			75.98				61.76				54.45				55.10				62.76				47.83				42.11				39.49				48.25			49.62			52.19			66.34			91.31			88.78
Y	291.02			337.42				260.17				232.89				230.36				284.38				205.41				182.92				163.44				198.78			213.74			215.78			324.69			427.14			436.52
Ho	10.96			12.31				9.65				8.34				9.36				11.04				8.09				6.57				5.94				7.71			8.30			7.85			11.49			15.27			15.46
Er	24.87			28.57				22.31				19.37				19.51				25.03				18.05				16.28				13.90				18.11			17.64			18.48			28.03			34.86			36.47
Tm	2.33			2.83				2.39				2.09				2.09				2.73				2.12				1.62				1.42				1.73			1.71			1.84			3.06			4.01			3.86
Yb	12.10			15.67				11.58				11.92				10.18				12.97				10.32				8.72				7.89				8.68			8.42			7.78			16.07			19.51			18.50
Lu	1.49			1.81				1.38				1.37				1.25				1.59				1.27				1.02				0.87				1.01			1.02			1.09			1.79			2.17			2.09
Intergranular Phase																					Late-Stage Phase
ZK3306																					ZK4506
Sample			b46-3-3			b48-2-1			b48-2-2				b48-3-1				b54-2-2				b12-1-1				b12-1-2				b12-2-1				b12-2-2				b12-2-3			b25-1-1			b25-1-2		b25-1-3		b25-3-1			b25-3-2
Rb			0.31			0.19			0.22				0.30				0.14				0.14				0.22				0.14				0.24				0.15			0.34			0.25		-		0.21			0.24
Ba			3.86			3.56			25.21				3.15				9.79				7.36				10.59				1.78				6.17				3.49			2.23			1.32		2.79		2.08			2.25
Th			51.61			60.53			60.78				98.52				91.78				105.70				72.49				77.10				94.08				91.64			54.09			40.84		85.94		41.45			59.12
U			18.92			16.81			20.24				19.78				26.72				27.81				20.97				21.81				24.73				27.22			20.49			19.15		26.54		14.72			16.02
Nb			0.13			0.04			-				-				-				-				-				-				0.03				0.04			0.05			0.05		-		-			0.04
La			1127.19			824.33			797.21				954.86				879.70				1043.78				936.84				1034.56				1068.35				952.58			1285.87			1183.96		1301.31		1193.19			1227.08
Ce			1815.77			1695.08			1620.34				1812.62				1643.87				2032.36				1771.27				1974.50				2003.96				1808.73			2509.57			2318.76		2499.91		2257.72			2323.88
Pb			4.22			13.85			12.13				20.44				14.87				51.68				43.07				105.32				88.33				58.72			2.41			1.86		3.88		1.52			1.84
Pr			175.57			186.35			179.69				191.00				165.78				215.57				193.44				205.84				208.08				185.51			261.95			241.17		259.71		233.51			229.55
Sr			2238.79			881.15			909.01				764.34				888.42				1646.18				1424.86				1450.65				1535.56				1267.72			2002.57			1883.72		1937.21		1782.84			1742.17
Nd			672.97			777.03			754.32				781.53				676.93				877.39				810.50				842.40				838.81				743.29			1029.06			963.89		996.98		932.12			889.11
Sm			127.38			141.54			136.22				133.62				121.60				149.98				139.15				133.58				141.33				123.70			170.33			161.22		156.20		156.56			146.87
Zr			5.75			7.55			7.38				8.81				5.04				3.65				2.99				2.49				2.60				3.56			3.44			3.34		1.73		4.70			3.29
Hf			0.04			-			-				0.05				-				-				-				-				-				-			-			-		-		-			-
Eu			43.19			34.17			31.33				32.15				29.16				33.13				31.56				32.56				31.64				30.78			42.22			38.81		38.93		36.57			36.95
Gd			125.28			122.81			116.41				118.68				103.25				120.64				115.95				106.46				101.73				99.33			134.60			129.01		118.60		123.61			111.87
Tb			14.90			12.36			11.92				12.54				11.01				11.12				11.01				11.15				10.22				10.27			14.24			14.42		12.37		13.56			12.19
Dy			82.60			61.88			53.29				59.12				54.03				56.28				53.99				51.73				44.61				48.60			70.70			68.85		61.63		66.97			59.84
Y			415.53			263.63			236.20				233.53				234.30				215.52				211.15				197.53				194.34				183.03			288.88			289.84		244.29		276.31			254.43
Ho			14.04			9.67			8.44				9.02				8.77				8.40				8.74				7.70				7.42				7.37			11.20			10.72		9.36		10.89			9.31
Er			33.26			20.89			17.78				20.66				19.10				18.13				17.62				17.17				15.54				15.18			25.02			23.74		21.27		24.42			20.79
Tm			3.59			2.15			1.93				2.03				2.00				1.79				1.89				1.70				1.72				1.50			2.61			2.57		2.03		2.39			2.15
Yb			17.20			11.16			9.87				9.29				9.95				8.38				9.38				8.23				8.05				7.53			11.97			11.72		10.40		10.75			9.99
Lu			1.86			1.48			1.25				1.24				1.20				0.94				1.11				1.00				0.79				0.88			1.40			1.20		1.06		1.43			1.25

Note: Below the Detection.

. The rare earth element (REE) concentrations in the inclusion, intergranular, and late-stage phases of apatite were notably high, with total REE (ΣREE) values of 9764.66 ppm, 4080.48 ppm, and 4800.14 ppm, respectively, showing considerable variation. The distribution of these elements demonstrated a clear enrichment in light rare earth elements (LREE), with LREEs/HREEs ratios ranging from 12.83 to 17.35, 13.54 to 29.46, and 17.67 to 22.58, respectively. This indicates a significant light-heavy REE fractionation. As shown in Figure 4, the chondrite-normalized REE patterns were consistently right-skewed, with weak Eu negative anomalies (mean Eu/Eu* values of 0.69, 0.85, and 0.82, respectively) and no discernible Ce anomalies (mean Ce/Ce* values of 1.04, 1.03, and 1.05, respectively). Additionally, apatite was enriched in high-field-strength elements (HFSE) such as Th and U while being notably depleted in large ion lithophile elements (LILE) like Rb, Ba, and Sr, as well as other HFSEs such as Pb, Nb, Zr, and Hf.

6. Discussion

6.1. Estimating Melt Volatile Compositions from Apatite: Insights into Magmatic Sources and Ore-Forming Environments

The halogen and hydroxyl (OH) components in apatite have been shown to be indicative of magmatic volatile composition [1]. Assuming that the anion sites in apatite are occupied solely by F, Cl, and OH, satisfying the equation X_F + X_Cl + X_OH = 1, the molar ratios of volatile components were calculated [30]. As shown in Figure 5, the endmember compositions of the three types of apatite from the Wagiritag cluster near the F–OH line, away from the Cl endmember, indicate crystallization in a silicate magmatic environment [31]. This compositional trend is analogous to apatite from the Skaergaard layered rocks in Greenland and the upper belt of the Bushveld Complex in South Africa, which host magnetite [11]. In late-stage apatite, a relative enrichment in chlorine (Cl) was observed. Yet, its composition differs significantly from the highly chlorine-rich apatite found in China’s Jinchuan copper-nickel sulfide deposits [32], as well as from chlorine-rich apatite associated with platinum group element (PGE) deposits in the lower belt of South Africa’s Bushveld Complex and the Stillwater Complex in Montana, USA [33]. This divergence may indicate an association between volatile composition and mineralization types. Notably, due to its comparatively smaller ionic radius and higher electronegativity than chlorine, fluorine (F) exhibits geochemical behavior more akin to oxygen (O). In contrast, chlorine (Cl) tends to behave more like sulfur (S). It is well established that chloride-enriched fluids are characteristically associated with copper-nickel sulfide mineralization, while fluoride-hydrogen-enriched melts preferentially favor Fe-Ti oxide mineralization. Conversely, Cl–F-enriched melts or Cl–OH fluids are often linked to platinum group element (PGE) mineralization [32].

The geochemical properties of fluorine (F) and chlorine (Cl) differ significantly. It is widely acknowledged that during magma differentiation, ascent, and degassing, F tends to remain in silicate melts, while Cl preferentially distributes into the fluid phase. As fractional crystallization progresses, the residual melt gradually becomes enriched in F. A comparative analysis of the elemental composition of Wajilitag apatite reveals a negative correlation between F and Cl content, while F shows a positive correlation with the F/Cl ratio (Figure 6). As chlorine (Cl) is predominantly enriched in mantle-derived fluids, while fluorine (F) is typically concentrated in crustal materials, and sedimentary components are generally characterized by elevated fluorine contents relative to chlorine, the incorporation of such sedimentary materials into magma sources followed by partial melting imparts their distinctive F/Cl signature to the derived magmas. This geochemical characteristic is subsequently preserved in crystallized magmatic minerals, particularly apatite. Consequently, elevated F/Cl ratios in apatite may indicate the incorporation of sedimentary material into the magmatic source region [34,35]. For example, studies of the Zijinshan porphyry copper deposit reveal that apatite from the ore-forming magmas exhibits chlorine-enriched characteristics, indicating an association with volatile-rich mineralizing fluids. Conversely, the incorporation of sedimentary materials into the magma can lead to an elevated F/Cl ratio in apatite [36]. Previous studies have demonstrated that the isotopic characteristics (variations in Sr-Nd and Mg-Zn isotopes) of the Early Permian mafic-ultramafic intrusive rocks in the Wajilitag area, combined with in situ major and trace element analyses of minerals such as olivine, collectively indicate the involvement of subducted sedimentary materials in the magma source [37,38]. The elevated F/Cl ratios in apatite further suggest the incorporation of subducted sediments into the magmatic source region of these bodies [39]. Consequently, the elevated F/Cl ratios observed in early-stage apatite from this study provide additional evidence supporting the hypothesis that subducted sedimentary contributions play a significant role in the magmatic source region.

The behavior of chlorine (Cl) in the fluid phase is analogous to that of incompatible elements, gradually enriching in evolving magmas prior to apatite crystallization. In magmas undergoing solely paragenetic crystallization of anhydrous minerals, the partition coefficient for fluorine (F) and chlorine (Cl) between the solid phase and melt approaches zero. Consequently, the F/Cl ratio is expected to remain stable in both the melt and apatite [1]. However, when fractional crystallization of hydrated minerals, such as amphibole and biotite, occurs, the stronger incompatibility of chlorine relative to fluorine leads to elevated chlorine/fluorine ratios in the residual melt and apatite. Alternatively, the dissolution and escape of volatile fluids from the melt can also influence halogen abundances. During crystallization, volatilization may be triggered by a reduction in confining pressure or an increase in volatile partial pressure, typically resulting in a decrease in Cl/F ratios in apatite [33]. In this study, the Cl/F ratio of the inclusion phase apatite ranged from 0.08 to 0.58 (mean 0.31), the intergranular phase from 0.10 to 0.58 (mean 0.26), and the late-stage phase from 0.53 to 4.60 (mean 1.20). The substantial increase in Cl/F ratios observed in late-stage apatite is likely related to the crystallization of hydrated minerals, including amphibole and biotite.

The halogen components in apatite can be used to estimate halide concentrations in the parent magma melt. Preliminary experimental studies suggest that fluoride (F) is more readily incorporated into the apatite lattice compared to chlorine (Cl) and hydroxide (OH). Although the F–Cl–OH ratio in apatite does not directly reflect the true proportions of these components in the melt during crystallization, elevated F content in apatite can still indicate relatively high F concentrations in the magma to some extent. Since Cl, F, and OH occupy equivalent lattice sites in apatite, their abundances are ultimately constrained by apatite stoichiometry. As a result, Cl and F concentrations in the melt cannot be directly estimated using Nernstian apatite/melt partition coefficients for Cl and F [7]. This study builds on previous research regarding diagenetic physicochemical conditions. It assumes an apatite crystallization temperature of 977 °C [40,41] and applies thermodynamic partition models developed by earlier researchers [34,42,43] to calculate halide concentrations in melts equilibrated with apatite. The results indicate that the average Cl content in magmas equilibrated with inclusion-phase, intergranular-phase, and late-phase apatite is 2619.61 ppm, 2400.11 ppm, and 6234.93 ppm, respectively, while the average F content is 1315.27 ppm, 1299.38 ppm, and 1319.65 ppm, respectively. Although the magma in equilibrium with late-stage apatite is enriched in chlorine, all three apatite types correspond to magmas with comparable and collectively high F and Cl concentrations, characterizing a volatile-enriched magmatic system.

Volatile-rich magmas typically exhibit lower viscosity and an extended crystallization path. During magmatic crystallization, early-formed silicate minerals such as olivine and pyroxene preferentially incorporate magnesium, thereby enriching the residual melt in iron. If the magma is volatile-rich and experiences prolonged crystallization, this fractionation process becomes more efficient. The presence of volatiles suppresses the early crystallization of magnetite, allowing iron to continuously accumulate in the residual magma and ultimately generating an iron-enriched melt. The low forsterite (Fo) contents (e.g., 67–76 mol%) observed in olivine from the Wajilitag intrusion provide direct evidence for the crystallization from such an iron-rich residual melt [38], as low-Fo olivine typically forms in highly differentiated, iron-rich magmas. As magmatic evolution progresses, iron and titanium eventually reach saturation under favorable physicochemical conditions, leading to the crystallization of large-scale Fe-Ti oxide ore deposits. The high concentrations of TiO₂ and Fe₂O₃ + FeO, along with the observed Fe-Ti oxide mineralization in the Wajilitag intrusion [38], collectively corroborate this genetic process.

The SO₃ content of apatite has been shown to serve as a reliable indicator of the redox state of silicate melts during crystallization [44] and of sulfur content in the melts themselves [45]. Two primary methods are currently used to estimate melt sulfur content based on apatite SO₃. The first method involves calculations utilizing the apatite saturation temperature (AST) [46], while the second assumes a direct correlation between apatite and melt SO₃ content, independent of temperature [6]. It is clear that neither of these methods provides highly precise calculations for the SO₃ content in the melt; rather, they offer only approximate estimates. This study employs the relationship between apatite and melt SO₃, as described by the second method, to calculate sulfur content in the magmatic-ultramafic melt at Wajilitag. The calculation formula is as follows:

SO_3ap (wt%) = 0.157 × lnSO_3melt (wt%) + 0.9834 (r² = 0.68)

(1)

The results indicate that the average sulfur content in magmas equilibrated with inclusion, intergranular, and late-stage apatite phases is 0.0025 wt%, 0.0023 wt%, and 0.0024 wt%, respectively. The sulfur concentrations in these three equilibrated magmas are consistently low and similar, suggesting that no significant sulfide melting occurred throughout the entire magmatic evolution. As a result, the sulfur content remained uniformly low across the system, whether in the bulk magma or local fluids. This consistently low sulfur characteristic is reflected in all apatites, regardless of their high-F or high-Cl classification. These findings provide compelling evidence that Fe-Ti oxide can accumulate iron without being reduced or competed for by sulfides.

6.2. Apatite Reveals Magma Oxygen Fugacity

Apatite exhibits increased resistance to weathering and alteration, making it a more effective indicator of magmatic oxygen fugacity than zircon. Previous studies have shown that the oxidation states of sulfur in apatite (S, S²⁻, S²⁺, S⁴⁺, S⁶⁺) are controlled by magmatic oxygen fugacity [47,48], thus recording the redox conditions of the primary magma. Given that only the partitioning behavior of S⁶⁺ has been extensively studied, the partition coefficients for other sulfur oxidation states remain poorly understood. As a result, it is not possible to directly calculate magmatic oxygen fugacity from the sulfur content in apatite.

The oxygen fugacity of the parental magma is also influenced by the abundance of other elements in apatite, including manganese (Mn), europium (Eu), and cerium (Ce). The concentration of manganese (Mn) in apatite can be utilized to estimate magmatic oxygen fugacity, which reflects the prevailing redox conditions [49]. This relationship is expressed by the following equation:

log f_O2 = −0.0022 (±0.0003) Mn (ppm) − 9.75 (±0.46)

(2)

The following calculations are derived from the previously mentioned formula. These calculations yield log f_O2 values for the inclusions, intergranular phases, and late-stage apatite in the Wajilitag magneferrous–ultramafic body as follows: the data sets range from −12.00 to −9.75 (average: −10.37) and from −11.79 to −9.75 (average: −10.55), respectively.

Nevertheless, the role of manganese (Mn) in apatite as an indicator of redox conditions remains a subject of considerable debate. Previous studies have shown that magmatic temperature and pressure exert a significant influence over the partitioning of manganese between apatite and silicate melts [50]. The mineralogical composition of apatite has been observed to show a positive correlation with the Mg content of whole-rock samples. However, this correlation seems to be independent of magmatic oxygen fugacity, as inferred from zircon-based analysis. Instead, the correlation is primarily influenced by melt composition and the degree of magmatic differentiation, rather than oxygen fugacity controls [51]. Consequently, variations in the content of a single element in apatite are inadequate to accurately determine changes in magmatic oxygen fugacity. As such, the validity of manganese (Mn) as a reliable oxygen fugacity indicator warrants further investigation.

Eu and Ce anomalies in apatite are widely recognized as key indicators of magmatic redox conditions [49,52]. Under oxidizing conditions, Ce³⁺ and Eu²⁺ are oxidized to Ce⁴⁺ and Eu³⁺, respectively. Higher oxygen fugacity facilitates the enrichment of Eu³⁺ in the melt, while at the same time, lower concentrations of Ce³⁺ and higher Eu³⁺ concentrations hinder Ce³⁺ incorporation into the apatite lattice. This leads to weak positive to negative Ce anomalies and moderate negative to positive Eu anomalies. In contrast, under reducing environments, higher Ce³⁺/Ce⁴⁺ and Eu²⁺/Eu³⁺ ratios, along with lower Eu³⁺ concentrations in the melt, produce positive Ce anomalies and strong negative Eu anomalies in apatite [40]. Based on these observations, the weak negative Eu anomaly observed in Wajilitag apatite suggests elevated Eu³⁺ concentrations, along with reduced Ce³⁺/Ce⁴⁺ and Eu³⁺/Eu²⁺ ratios. This implies that the formation of the mineral occurred under relatively high oxygen fugacity conditions.

The elemental composition of apatite during magmatic crystallization is influenced by a range of factors. For instance, plagioclase crystallization consumes Eu from the magma, and its fractional crystallization may lead to a reduction in Eu content. Simultaneously, changes in whole-rock composition, driven by variations in melt composition and fractional crystallization, can complicate the interpretation of plagioclase crystallization effects. This consequently reduces the reliability of apatite Eu anomalies as indicators of oxygen fugacity [53]. In contrast, the influence of Ce-bearing minerals and melt composition on apatite Ce anomalies is comparatively less significant. Furthermore, this study shows that Ce anomalies do not exhibit a strong correlation with whole-rock Ce anomalies or total rare earth element content [54]. Consequently, variations in the multivalent elements Eu and Ce can still reflect partial insights into magmatic oxygen fugacity. In the Wajilitag magmatic-ultramafic complex, Eu/Eu* values for inclusions, intergranular phases, and late-stage apatite were 0.56–0.82 (mean = 0.69), 0.75–1.13 (mean = 0.85), and 0.75–0.88 (mean = 0.82), respectively; Ce/Ce* values were 0.99–1.09 (mean = 1.04), 0.9–1.09 (mean = 1.03), and 1.02–1.07 (mean = 1.05), respectively. These results suggest relatively stable Ce/Ce* values within the intrusion. As shown in Figure 7, the Eu/Eu* values for all apatite samples exceed 0.5, indicating an environment with high oxygen fugacity [54]. Among these samples, the inclusional-phase apatite exhibits a pronounced negative anomaly, reflecting more reducing conditions during the early magmatic stages, while the intergranular-phase and late-stage apatite indicate more oxidizing conditions during the middle and late crystallization phases.

6.3. Apatite Geochemistry of Mafic–Ultramafic Intrusion in the Wajilitag Area: Constraints on Magma Source and Evolution

6.3.1. Tracing the Mantle Source: Evidence from Apatite Trace Elements

Apatite continuously saturates and precipitates during magma evolution, crystallizing at various stages and thus preserving a relatively complete record of the magma’s evolutionary history. As a result, magmatically derived apatite is highly effective in tracing magma sources and evolutionary processes [55]. The rare earth element holmium (Ho) as well as Yttrium (Y) have similar ionic radii; however, hydrothermal alteration can induce Y/Ho fractionation, with hydrothermally derived samples typically exhibiting higher Y/Ho ratios due to the preferential migration or selective precipitation of Ho by fluids. Consequently, the relationship between La/Sm and Y/Ho serves as a useful tool for distinguishing between magmatic and hydrothermally derived apatite, providing crucial insights into deposit genesis and fluid evolution. As illustrated in Figure 8a, the data point for this apatite sample falls within the magmatic zone, confirming its magmatic origin and supporting the earlier conclusion regarding the magmatic genesis of the apatite.

The La/Sm ratio in magma is only minimally affected by fractional crystallization or partial melting; it is primarily controlled by the degree to which subduction-related marine sediments have melted [56]. Primary magmatic apatite typically retains rare earth compositions that closely align with those of the parent magma, making apatite La/Sm ratios a reliable tool for tracing those of the magma [54]. Variations in the ratio can indicate magmatic mixing or assimilation of surrounding rocks, processes often associated with increased thorium (Th) content and decreased apatite strontium (Sr)/thorium (Th) ratios. On the other hand, the addition of fluids from the dehydration of subduction zone materials maintains relatively stable La/Sm ratios but widens the range of Sr/Th ratios in the melts. These two processes can be distinguished through the use of a two-component diagram of apatite Sr/Th versus La/Sm. As shown in Figure 8b, the apatite La/Sm ratio remains relatively stable, ranging between 5 and 9, while the Sr/Th ratio exhibits broader variability, ranging from 0 to 120. The combined variation in both ratios suggests the incorporation of dehydration fluids from the subduction zone into the Wajilitag magma.

Chondrite-normalized rare earth elements’ (REEs) distribution patterns are widely used as tracers for petrogenesis. The high REE concentrations in apatite allow for precise measurement, making it particularly valuable in determining the origin of its host rocks. In the apatite rare earth element ternary diagram (Figure 9) [57], the inclusion phase of the apatite plots in the mantle-crust mixture (MC) region. The intergranular phase is distributed in the mantle (M) and mantle-crust mixture (MC) regions, and the late-stage phase falls within the mantle (M) region. This indicates that the host rock’s source region is primarily mantle-derived with some crustal material incorporated. The REEs and trace element signatures in apatite suggest a different petrogenetic source than that indicated by previous whole-rock geochemical studies of the Wajilitag mafic–ultramafic intrusions. The Wajilitag intrusions exhibit LREE-enriched, right-sloping chondrite-normalized REEs patterns without significant Eu or Ce anomalies. Their primitive mantle-normalized trace element diagrams show distinct depletions in Th, U, Zr, and Hf, and enrichments in Ba, Ti, Nb, and Ta—characteristics resembling ocean island basalts (OIB), consistent with derivation from an enriched mantle and are associated with mantle plume activity [38]. The whole-rock geochemistry points to a dominant enriched mantle source, whereas the high trace-element sensitivity of apatite reveals the presence of recycled crustal components within this source. This indicates that the crustal signature in apatite is inherent, most probably derived from ancient subducted material in the mantle, rather than from crustal contamination during ascent.

The halogen and sulfur compositions of magmatic apatite serve as valuable tracers of magmatic sources [3,54]. During seafloor hydrothermal alteration, mafic oceanic crust, including the overlying sediments and basalts, absorbs sulfates, while its volatile components (such as Cl, S, and H₂O) become enriched through interactions with seawater. As a result, altered oceanic crust and marine sediments typically exhibit elevated Cl and S concentrations. During plate subduction, the dehydration of oceanic crust and sediments releases S- and Cl-rich fluids into the mantle wedge, leading to the formation of volatile-rich mafic melts derived from this process within the enriched lithospheric mantle [58,59]. The apatite compositional characteristics observed in this study (primarily high fluorine (F) and low chlorine (Cl), with some instances of high Cl and low F, and uniformly low sulfur (S)) deviate from the previously described scenario. The evolutionary process may involve the dehydration of subduction zones at depth, releasing fluids rich in F, Cl, H₂O, large-ion lithophile elements (LILEs, such as K, Ba, and Sr), and light rare earth elements (LREEs). These highly oxidizing fluids then infiltrate the overlying mantle wedge, replacing it in a process marked by the incorporation of fluorine (F) into mantle peridotite minerals (e.g., clinopyroxene) due to its compatibility. In contrast, chlorine (Cl), due to its high reactivity and fluid affinity, primarily remains in the fluid phase and can migrate along with metallic elements. Sulfur, often oxidized to sulfate ions (SO₄²⁻) in subduction fluids with high oxygen fugacity, behaves differently. As sulfates are significantly more soluble in silicate melts than sulfides, sulfur is largely transported away by fluids without being incorporated into the mantle. This results in relatively sulfur-poor but volatile-rich mantle zones. Ultimately, this process shapes a metasomatized mantle source region enriched in F, Cl, and H₂O, characterized by high oxygen fugacity (f_O2) and sulfur depletion. This enriched source region serves as the material and energy foundation for subsequent geological processes.

6.3.2. Apatite Reveals Crystallization History and Magmatic Evolution

The concentrations of rare earth elements (REE), strontium (Sr), and europium (Eu) in apatite can serve as indicators of the compositional evolution of the host magma [60]. Mass balance calculations suggest that apatite predominantly hosts most REEs, uranium (U), thorium (Th), Sr, and other elements. As a result, apatite plays a critical role in controlling the geochemical behavior of these elements. The concentrations of these elements in the melt can be calculated using the distribution coefficient formula:

C = C_ap/D

(2)

where C represents the element concentration in the melt, C_ap is the concentration of the element in apatite, and D is the distribution coefficient. The distribution coefficients used in these calculations were referenced from

Table A6. Apatite-melt partition coefficient.

Element	U	La	Ce	Pr	Sr	Nd	Sm	Eu	Gd	Tb	Dy	Y	Ho	Er	Tm	Yb	Lu
D	2.6	2.44	2.99	3.32	2.59	3.38	3.59	2.34	3.78	3.53	3.25	3.61	3.46	3.14	3.04	2.58	2.77

[61,62,63,64]. To recover the trace element composition of the parent magma in the intrusive complex, we divided the apatite content in the mineralized rock units of the Wajilitag intrusive complex by its distribution coefficient relative to the basaltic magma. As shown in Figure 10, standard diagrams of trace elements in the primitive mantle exhibit a right-sloping trend with light rare earth element (LREE) enrichment, strontium (Sr) depletion, and positive Eu anomalies ranging from 0.87 to 1.78, with an average of 1.27. No significant cerium (Ce) anomalies are observed, with values ranging from 0.98 to 1.09 and an average of 1.04. The average Eu/Eu* values for the inclusion phase, the intergranular phase, and the late-stage apatite equilibrium melt are 1.08, 1.34, and 1.30, respectively.

Strontium (Sr) has been observed to act as a compatible element in apatite and plagioclase but exhibits incompatibility in pyroxene, olivine, and Fe-Ti oxides. Consequently, variations in Sr content in residual magma are primarily controlled by the crystallization of apatite and plagioclase. Extensive crystallization of either mineral results in Sr depletion in the remaining melt. The Sr partition coefficient for apatite in layered intrusions is typically lower than that for plagioclase. Calculations based on trace element partition coefficients between apatite and the melt reveal a pronounced negative Sr anomaly in the melt, suggesting that extensive plagioclase crystallization preceded apatite crystallization–i.e., apatite crystallized after plagioclase. Theoretically, the fractional crystallization of plagioclase should lead to Eu depletion in the melt. However, a positive Eu anomaly was observed. Possibly due to the presence of Eu as Eu³⁺ under high oxygen fugacity conditions, the absorption capacity of plagioclase for Eu is significantly reduced, leading to the relative enrichment of Eu in the residual melt. Meanwhile, the crystallization of abundant Fe-Ti-rich oxides and pyroxenes, coupled with the dilution of rare earth elements and altered distribution patterns, serves to further obscure the residual negative Eu anomaly.

The mineral assemblage associated with apatite has been shown to significantly influence crystal/melt partition coefficients, thereby modulating the abundance of specific elements within apatite. Plagioclase crystallization notably affects the Sr and Eu content of the melt, with minimal impact on the Y content. This results in a marked decrease in the melt’s Sr/Y ratio. Given the susceptibility of whole-rock trace elements to post-depositional processes such as weathering and alteration, contrasted with the strong resistance of apatite to such alterations, the apatite Sr/Y ratio has been identified as an effective indicator for detecting plagioclase fractional crystallization. The apatite observed in this study shows a clear trend consistent with plagioclase fractional crystallization, as evidenced by enhanced Eu negative anomalies and decreased Sr/Y ratios (Figure 11). The sequence of plagioclase crystallization, from highest to lowest, is as follows: inclusion apatite, intergranular apatite, and late-stage apatite.

Furthermore, the ratios of Ce/Pb and Th/U in apatite have been identified as indicators of fluid mobility during magmatic processes [35]. Phosphate minerals from the Wajilitag intrusive complex exhibit high Ce/Pb ratios (7.32–2367.11, with an average of 497.91) and low Th/U ratios (1.47–13.27, with an average of 3.81), suggesting significant fluid activity during their crystallization.

Early separation crystallization of pyroxene/augite results in a positive correlation between apatite La/Yb and ΣREEs [4]. In contrast, separation crystallization of amphibole leads to elevated apatite Sr, La content, and La/Yb ratios within the crystallizing magmatic system [55]. The La/Yb ratio in the inclusion apatite correlates positively with both total rare earth element content (ΣREEs) and the Sm/Yb ratio. This indicates that early separation crystallization of monoclinic pyroxene leads to the enrichment of light rare earth elements (LREEs) relative to heavy rare earth elements (HREEs) in the residual magma. In intergranular and late-stage apatite phases, the La/Yb ratio exhibits no significant correlation with ΣREE; however, it exhibits a positive correlation with the Sm/Yb ratio. The Sr/Y ratio in apatite correlates positively with Sr content, indicating ongoing crystallization within the magma. Conversely, La/Yb demonstrates a negative correlation with Sr, indicating either the replenishment of the newly formed magma with high Sr content or crustal mixing (Figure 12).

6.4. Implications for Genesis of Fe-Ti Oxides

Oxygen fugacity (log f_O2) is recognized as a critical factor controlling the in situ crystallization of Fe-Ti oxides. The mineralogical composition of magnetite–ilmenite pairs has been demonstrated to be a reliable indicator of their respective equilibrium temperature and oxygen fugacity. During cooling below the Fe-Ti oxide subsolidus, magnetite and ilmenite undergo re-equilibration via Fe-Ti exchange, ultimately approaching the Fe₃O₄ and FeTiO₃ endpoints, respectively. This study estimates the temperature and oxygen fugacity during re-equilibration near the solidus for Fe–Ti oxides in the Wajilitag granite, based on magnetite–ilmenite mineral pair compositions. Temperatures range from 548.31 to 705.88 °C, with an average of 621.04 °C, and log f_O2 values range from −23.62 to −15.59, with an average of −19.19. In the magnetite–ilmenite temperature–oxygen fugacity diagram (Figure 13), the samples are distributed along the ferroolivine–magnetite–quartz (FMQ) buffer curve. These temperatures are consistent with previously estimated closure temperatures for magnetite–ilmenite pairs in this body (462–671 °C; 407.77–664.58 °C) [23,65], both of which are significantly lower than the crystallization temperatures of magnetite and ilmenite. This finding suggests that the magma chamber likely underwent a prolonged cooling process during the formation of the body.

The genesis of magmatic Fe-Ti oxide deposits in layered rock bodies remains a subject of ongoing debate, with two primary viewpoints currently prevailing. One viewpoint proposes that the separation crystallization of silicate minerals (e.g., olivine, clinopyroxene) within mafic magmas leads to the enrichment of Fe-Ti oxides [66,67,68,69,70]. During magmatic evolution, immiscibility or intergranular melt separation may occur, resulting in the formation of both Fe-Ti-rich and silica-rich melts [71,72,73]. These immiscible, Fe-Ti-rich melts are thought to ultimately evolve into Fe-Ti oxide deposits. Considerable evidence from previous studies now supports the concept of silicate melt immiscibility in layered geological formations. This includes the observation of immiscible two-phase melts within apatite melt inclusions and the identification of apatite with distinct rare-earth element (REE) signatures. These REE signatures are believed to crystallize from silica-rich and iron-rich melts, as determined by the analysis of apatite REE compositions [11,74]. However, the present study did not provide any such evidence, suggesting that further research is necessary to fully resolve this issue.

At present, the scientific community generally accepts two principal models that provide satisfactory explanations for the genesis of the Tarim magmatic province: the mantle plume model [17] and the slab subduction-lithosphere-plume interaction model [37,75]. This study demonstrates that the compositional characteristics of apatite offer crucial evidence for understanding Fe-Ti oxide mineralization within the Wajilitag mafic-ultramafic intrusive complex. Phosphate minerals from Wajilitag have been shown to exhibit high F/Cl ratios and low sulfur content. In conjunction with the locally occurring high-chlorine phosphates, these characteristics collectively support a composite genesis model driven by mantle plume heat and subduction-metasomatic enrichment of partially melted mantle source material:

(1)

In the pre-subduction stage, the subducting slab underwent deep dehydration, releasing fluids enriched in large ion lithophile elements (LILE), light rare earth elements (LREE), F, Cl, and H₂O, with elevated oxygen fugacity (f_O2). These fluids migrated upward, extensively metasomatizing the overlying lithospheric mantle. During this process, F, because of its compatibility, became incorporated into metasomatic minerals such as muscovite and amphibole. Cl predominantly remained in the fluid phase, while sulfur (S) was oxidized to sulfate (SO₄²⁻) and subsequently migrated out of the system with the fluid. This resulted in the formation of a metasomatized lithospheric mantle region characterized by elevated f_O2, abundant volatiles (F/Cl/H₂O), and sulfur-poor metasomatic rocks, establishing the material and geochemical foundation for mineralization.

(2)

In the triggering phase, ascending mantle plumes from the deep mantle interacted with the metasomatized lithospheric mantle. The primary role of these plumes was to supply substantial heat flux, rather than provide mineralizing materials. This thermal effect has been shown to induce thermal erosion and lithospheric thinning, significantly lowering solidus temperatures and triggering high-grade partial melting in the enriched mantle source region. Without the thermal input from the mantle plume, this enriched mantle would have undergone only minor melting and would not have been capable of generating substantial amounts of magma.

(3)

In the melting and evolution stage, the altered, enriched mantle underwent extensive partial melting, producing large volumes of sulfur-undersaturated parent magmas. These magmas inherited the source region’s high F/Cl ratio, high f_O2, and sulfur-poor characteristics. Following magmatic emplacement, the magma experienced significant crystallization and differentiation within the chamber. Volatiles such as F and Cl acted as fluxing agents, reducing viscosity and solidus temperatures while prolonging the duration of differentiation. The initial crystallization of mafic silicate minerals (e.g., olivine, clinopyroxene) resulted in the exceptional enrichment of incompatible elements such as Fe, Ti, V, and P in the residual melt.

(4)

During the mineralization process, ferrous ions (Fe²⁺) were oxidized to ferric ions (Fe³⁺) under conditions of elevated oxygen fugacity. Simultaneously, vanadium (V) was incorporated into the oxide lattice as a solid solution. Late-stage melts reached a state of supersaturation with respect to Fe-Ti oxide, which induced extensive crystallization, leading to the formation of ore bodies with layered or quasi-layered morphology. Late-stage exsolution or assimilation of chlorine-rich fluids from host rocks further activated and concentrated ore-forming elements (e.g., Fe, Ti, and V as chlorine complexes). This process is key to explaining the complex ore body structure and the presence of locally high-chlorine apatite.

7. Conclusions

The major and trace element composition of apatite in the Wajilitag magmatic Fe-Ti oxide deposit reveals several key characteristics:

(1)

A negative correlation between fluorine (F) and chlorine (Cl) in apatite is evident. Inclusion-phase and intergranular-phase apatites are F-rich and Cl-poor, while late-stage apatite is the opposite. All apatites are magmatic in origin. The elevated F/Cl ratio suggests that the parent magma originated from an F-rich, subducted metamorphic mantle source, with late-stage Cl enrichment likely linked to high volatile content or fluid dissolution in magma.

(2)

Estimates of magmatic volatilization and oxygen fugacity, based on apatite composition, indicate similar F, Cl, and sulfur contents across the three apatite types, with low sulfur content and high oxygen fugacity.

(3)

Apatite exhibits strong LREE-HREE fractionation, significant REEs’ enrichment, a weak negative Euanomaly, and distinct trace-element patterns (enriched in Th, U; depleted in Rb, Ba, Sr, Nb, Zr, Hf). This geochemistry indicates a predominantly mantle-derived magma with minor crustal contamination, further supported by Sr/Th and La/Sm ratios that reflect fluid input from subduction dehydration.

(4)

We hypothesize that ancient subduction fluids pre-altered the lithospheric mantle, creating a source region rich in volatiles (F, Cl, H₂O), high in oxygen fugacity, and sulfur-poor. This preconditioning provided the material basis and oxidizing environment for mineralization. As mantle plumes ascended, they induced large-scale partial melting, with magma undergoing fractional crystallization and ultimately forming Fe-Ti oxide deposits.