The dunites are situated along the south and north-western margins, in the central part and occasionally scattered as schlieren in the clinopyroxene peridotite of the western border and are divided into two types: light yellow dunites and yellow-green dunites (Figure 2). The light yellow dunites underwent silicification and carbonatation, were exposed to the surface and their color lightened (Figures 3A,D,E; Figures 4A–K). In the outcrop, dunites show small-scale layering features and typical cumulate textures (Figure 3F). These rocks primarily consist of olivine pseudomorphs and chromites with minor altered clinopyroxene relics (<5 vol%) and magnitites. Olivines in light yellow dunite are nearly all altered to antigorite-lizardite-magnetite assemblages exposing mesh textures due to serpentinization. The light yellow dunites comprise most of the chromite orebodies (Figures 4A–G).

Despite undergoing serpentinization, yellow-green dunites with unaltered olivine cores or whole grains are relatively fresh compared to light yellow dunites. Olivines in yellow-green dunites are homogenous, fresh, and free of opaque inclusions, and show crystalized textures of variable size, subeuhedral-euhedral crystals, and angular shapes with variable grain sizes (from 0.3 to 0.8 mm; Figures 5A,B). Dunites are composed of cumulus olivines with varying densities, disseminated chromian spinels, and subordinate Cpxs (<3 vol%). Serpentinite and carbonatite veins cut through a large number of olivine grains. Dunite schlierents in clinopyroxene peridotites are the exclusive metallogenic rocks of chromites. Chromian spinels in yellow-green dunite are generally sparsely disseminated or scattered texture (Figures 5J,L).

Clinopyroxene peridotites are widely exposed in the central and east areas and weathered red with nodular appearances (Figures 3G,H). In the southeast of the biotite monzogranite, clinopyroxene peridotites underwent serpentinization, carbonatization, silicification, and turned brown due to abundant iron oxide. The boundary between clinopyroxene peridotites and dunites normally curves and transitions gradually into clinopyroxene dunites. Occasionally, clinopyroxene peridotites alternate with dunite and pyroxenite layers and cut by dunite veins (Figure 3D), indicative of magma differentiation and evolution. Clinopyroxene peridotites consist of olivines, clinopyroxenes, amphiboles, and some altered minerals with olivines (80–85 vol%), Cpxs (5–10 vol%), Cr-spinels (<3 vol%), and opaques (<2 vol%). Olivines are subeuhedral columnar crystals that primarily morph into serpentines. Clinopyroxenes are mainly diopsides and sometimes replaced by neogenic pargasites, edenites, and actinolites at their edges (Figures 5C,G,I). Due to serpentinization and weathering, clinopyroxenes were oriented as pseudomorphs in the olivine matrix and were mostly substituted by serpentines, carbonate minerals, tremolites, and magnetites.

Gabbros are exposed on the north and southeast margins of the mafic-ultramafic complex (Figure 2). They have equigranular texture and are composed of clinopyroxenes (40–45 vol%), plagioclases (35–45 vol%), hornblendes (<5 vol%), and 3–5% opaque minerals (mostly magnetite). Several outcrops in the field show signs of rhythmic layering of clinopyroxenes and plagioclases (Figures 3B,C). Most clinopyroxenes morph into actinolites or tremolites, and plagioclases normally change into zoisites (Figures 5D,F,H). Gabbros on the north and southeast margins often show strong foliation with mylonitic textures and sharp tectonic contacts to the peridotites. S-C fabrics and rotational speckles characterize the dextral ductility shear (Huang and Jin, 2006b). Additionally, as a general feature of Alaskan-type complexes, several gabbroic dykes with magmatic layering intrude into dunites and clinopyroxene peridotites (Figure 3C). The dykes may form in later-stage residual melts after fractionation rather than the cumulate in ophiolite (Irvine, 1974; Snoke et al., 1981; Himmelberg and Loney, 1995). Unaltered gabbro samples from the southern margin were collected for zircon U–Pb dating and geochemistry analyses.

Chromites occur primarily in light yellow dunites, with subordinate Cpx-enriched dunites. Classified by shapes and spatial distributions of the chromite orebodies, the orebodies are labeled as: banded, lenticular, veined, thin-bedded, and brecciated texture (Figures 4A–G). Nearly 70–80% of modal chromites in Hongshishan show primary magmatic layering where individual seams involve layers of massive, spotted, schlieren, and banded-texture chromites. Banded chromites, occasionally with crossbedding, are primarily composed of massive and varying degrees of disseminated chromian-spinel grains that interbed alternately with inch-scale thickness of layers chromian spinels (Figures 4F,H–L). Orebodies of the banded subzone are tabular and more extensive laterally. Many bands show small-scale faulting and cataclasis.

Most chromites show weak deformations with polygonal or angular crystal appearances and pull-apart fractures, cataclastic textures (Figure 4D), and occasional ductile deformations (such as elongated worm-like shapes), and cluster to form chromite bands or disseminated varieties, though morphologically distinct from podiform chromites in supra subduction zone environments (Zhou et al., 2014). Skeletal textures also occurred in chromites and were interpreted as indicators for rapid crystallization, possibly due to supersaturation processes (Greenbaum, 1977). The chromite textures in Hongshishan showed that disseminated chromites crystallized earlier than the massive chromites.

The determination of major, trace, and rare-earth elements was conducted at the Beijing GeoAnalysis Technology Co., Ltd. using X-ray fluorescence (XRF-1800; SHIMADZU) on fused glasses. Inductively coupled plasma mass spectrometry (ICP-MS, 7500; Agilent) was conducted at Beijing Createch Testing Technology Co., Ltd. on samples after acid digestion in Teflon bombs. Loss on ignition was measured after heating to 1,000°C for 3 h in a muffle furnace. The precision of the XRF analyses was within ±2% for oxides with >0.5 wt% and within ±5% for oxides >0.1 wt%. Sample powders (∼40 mg) were placed in Teflon bombs and dissolved using a mixture of HF and HNO₃ for 48 h at 190°C. The solution was evaporated to dryness, re-dissolved using concentrated HNO_3, and evaporated at 150°C to dispel any fluorides. Samples were diluted to approximately 80 g for analysis after dissolution in 30% HNO₃ overnight. An internal standard solution containing Rh was used to monitor signal drift during the analyses. Results from USGS standards indicated the uncertainty for most elements was ±5%.

The chemistry of unaltered minerals (olivines, pyroxenes, and chrome spinel, etc.) in silicates and oxides was conducted by wavelength-dispersive X-ray analysis using a JEOL electron-probe micro analyzer (EPMA) JXA-8230 at the Institute of Mineral Resources, Chinese Academy of Geological Sciences. We used 15 kV for the acceleration voltage, 20 nA for the beam current, a 5 μm beam diameter, and the counting time was between 20 and 40 s for major elements and 40–60 s for minor elements. SPI mineral standards (USA) were used for calibration. The precision for all elements analyzed exceeded 98.5%. The Cr- and Mg-numbers (Cr# and Mg#) of the chromian spinel were the Cr/(Cr + Al) and Mg/(Mg + Fe²⁺) atomic ratios, respectively. We assumed all Fe in silicates was ferrous.

Zircon U-Pb dating was conducted using an LA-ICPMS at Beijing GeoAnalysis Co., Ltd. The Resolution SE model laser ablation system (Applied Spectra, United States) was equipped with an ATL (ATLEX 300) excimer laser and a Two-Volume S155 ablation cell. The laser ablation system was coupled to an Agilent 7900 ICPMS (Agilent, United States). Zircons were mounted in epoxy discs, polished to expose the grains, ultrasonically cleaned in ultrapure water, then cleaned again prior to the analysis using AR grade methanol. Pre-ablation was conducted for each spot analysis using five laser shots (∼0.3 μm in depth) to remove potential surface contamination. The analysis was performed using a 30 μm diameter spot at 5 Hz and a fluence of 2 J/cm².

Zircon 91500 and GJ-1 were used as primary and secondary reference materials, respectively. Zircon 91500 was analyzed twice and GJ-1 was analyzed once every 10–12 analyses. Typically, 35–40 s of the sample signals were acquired after 20 s of gas background measurement. NIST 610 and ⁹¹Zr were used to calibrate the trace element concentrations as the external reference material and the internal standard element, respectively. The ages of the reference materials in the batch are as follows: 91,500 (1061.5 ± 3.2 Ma, 2σ), GJ-1 (604 ± 6 Ma, 2σ), and agreed with the reference value within definite uncertainty.

Supplement 1 lists the analytical data of major, trace, and REE for samples from each rock unit of the Hongshishan complex. Disseminated chromian spinels-bearing dunites had high volatile levels (LOI. 9.43–14.83 wt%, except for D002-7 and D003-6). For the dunites, SiO₂ compositions varied from 35.23 to 38.82 wt%, Al₂O₃ from 0.39 to 4.45 wt%, and TiO₂ from 0.00 to 0.07 wt%. While the composition of clinopyroxene peridotites was SiO₂, 29.55–42.61 wt%, Al₂O₃, 0.43–7.97 wt%, TiO₂, 0.00–0.11 wt%, they were nevertheless distinguished from the dunites by much higher Ca concentrations (CaO 1.13–8.68%). Pyroxenite components had the following compositions, SiO₂ from 40.65 to 45.76 wt%, Al₂O₃ from 9.31 to 16.05 wt%, and TiO₂ from 0.07 to 0.14 wt%; gabbros SiO₂ from 42.01 to 49.58 wt%, Al₂O₃ from 12.81 to 22.25 wt% and TiO₂ from 0.02 to 0.30 wt%. These compositions reflect depleted mantle melts as the primary intrusion source.

MgO compositions decreased systematically from the dunites (23.53–43.02 wt%) and clinopyroxene peridotites (29.43–36.23 wt%), followed by pyroxenites (14.75–31.85 wt%) and gabbros (8.46–18.22 wt%). In Harker diagrams (Figure 6), these major compositions, e. g. Al₂O₃, CaO, and TiO₂ were lowest in dunites and peridotites cores, increased gradually in pyroxenite and maximized in the marginal gabbros. Most dunites have low CaO levels (<1 wt%), whereas clinopyroxene peridotites and pyroxenites had higher CaO concentrations (1.13–8.68 wt% and 4.23–12.21 wt%, respectively). The gabbros had the highest CaO concentrations (10.66–15.06 wt%). TiO₂ levels in the dunites were virtually non-existent, but increased in clinopyroxene peridotites (0.00–0.11 wt%) and pyroxenites (0.07–0.14 wt%), and reached a maximum in the gabbros (0.02–0.30 wt%). The average Mg# (Mg/(Mg + Fe) atomic ratio) accordingly decreased from 0.90 in dunites, 0.90 in clinopyroxene peridotites, followed by 0.87 in pyroxenites to 0.82 in the marginal gabbros (Supplementary Table S1).

In terms of trace elements, dunites contain enriched levels of large-ion lithophile elements (LILE) (like Cs), high field strength elements (HFSE) (U, Zr, Hf), and Er. Also, significantly negative anomalies in Nb, Ba were observed, and Ti to a lesser extent (Nb/La ratios from 0.12 to 0.19, Th/Yb ratios from 0.39 to 2.78, Ta/La from 0.03 to 0.18 and Nb/Th from 0.01 to 0.55; Supplement 1). The gabbroic rocks shared many geochemical similarities with pyroxenites. They had higher trace element abundances and distinctive Ba, Pb, Cs, and Ti anomalies than the ultramafic rocks, which had lower Rb and Ba levels (Figure 7).

Different from the dunites in the Alaskan complex with identical flat REE patterns (Su et al., 2014; Habtoor et al., 2016), the dunites in Hongshishan display higher REE abundances (average 6.04 ppm) and enriched light REE (LREE) values relative to heavy REE (HREE) with high La_N/Yb_N ratios (average 4.49; Supplementary Table S1), negative Eu anomalies and pronounced positive Er anomalies with a typical V (or U)-shaped pattern (Song and Frey, 1989; Wang et al., 1996). The clinopyroxene peridotites showed negative Eu anomalies and wide compositional HREE variations. Pyroxenites and gabbros showed nearly flat REE patterns and positive Eu, Er anomalies (Figure 7). The positive Eu anomalies in the pyroxenites and gabbros indicated the presence of plagioclase in the crystallized rocks, which was in accord with other petrographic investigations. The positive Sr anomalies were consistent with the positive Eu anomalies and implied plagioclase accumulation.

One gabbro sample collected from the south margin of the mafic-ultramafic complex was selected for zircon U–Pb dating. Zircons separated from the sample ranged from 30 to 100 μm and were generally euhedral, colorless, and transparent. The cathodoluminescence (CL) images with internal growth zoning indicated a magmatic crystallization origin. (Figure 8). Sixteen-grain effective data formed a concordant group in the Concordia diagram with a weighted mean age of 366.1 ± 1.6 Ma (MSWD = 1.17), the age reflects the gabbro emplacement time. This age predates previous gabbro geochronological data and is regarded as an estimate of the Hongshishan complex crystallization age. Supplement 2 gives the analytical data of the Hongshishan gabbro Zircon U-Pb age.

The microprobe analysis data for chromian spinels, olivines, clinopyroxenes, hornblendes in the Hongshishan mafic-ultramafic complex are listed in Supplementary Table S3.

Chromian spinels in veins and massive chromites have average Al₂O₃ wt% levels of 0.97–25.81 (average 16.6 wt%), Cr₂O₃ (40.52–58.23, average 48.8), Fe₂O₃ (0–2.4, average 4.2), FeO (12.97–29.15, average 19.3), TiO₂ (0.06–0.46, average 0.2) and MgO (2.24–14.73, average 9.8), with narrow Cr# ranges (from 0.52 to 1.00; 0.70 on average) and Mg# (0–0.63; 0.38 on average). Various density degrees of chromian spinels in dunites have wide oxide ranges with Al₂O₃ ranging from 0 to 29.27 wt% (average 10.07 wt%), Cr₂O₃ from 0.20 to 57.74 (average 36.63 wt%), Fe₂O₃ from 2.97 to 70.97 (average 21.72 wt%), FeO from 15.9 to 31.86 (average 23.83 wt%), TiO₂ from 0 to 1.59 (average 0.36wt%), MgO from 0.27 to 14.73 (average 6.19 wt%), and Cr# from 0.44 to 1.0 (0.79 on average) and Mg# from 0 to 0.55 (0.25 on average). A portion of non-magnetized spinels compositions fell inside the field of ophiolitic podiform chromites (Figures 9A,D), while plots of Mg# vs. Fe³⁺/(Fe³⁺+Cr +Al) and Cr# vs. Fe²⁺#s (Figures 9B–D) show that most data approximate the field of the Alaskan-type complex worldwide and nearly approached the field of the Xiadong Alaskan-type complex (Barnes and Roeder, 2001; Habtoor et al., 2006; Su et al., 2012). The chromites of Hongshishan trend towards Fe³⁺-rich composition (Figure 2), a typical arc-related trend.

Olivines from the dunite are chemically homogeneous, with FeO (average 8.15 wt%), MnO (average 0.09 wt%) and NiO (average 0.35 wt%); they have Forsterite compositions that vary between Fo_90–93. Olivines from the clinopyroxene peridotites show slightly higher FeO and MnO levels (average 8.77 and 0.13 wt%, respectively) and the same NiO (average 0.35 wt%). They have Forsterite compositions that vary from Fo_88–92, all of which are higher than olivines in Abu Hamamid (Fo_74–81), Gabbro-Akarem (Fo_69–87), Genina Gharbia (Fo_80–86) and Alaskan-type complex (Khedr and Arai., 2016). The Hongshishan olivines plot of NiO, MnO versus Fo falls within in the Alaskan Global trend field but out of the Abyssal peridotite and Fore-arc peridotite fields (Figures 10B,C). A graph of FeO_T vs. Fo shows a systematic negative correlation (Figure 10D), and a typical fractional crystallization trend similar to olivines in the Alaskan-type complex (Figures 10A–C).

The Clinopyroxenes (Cpxs) from the clinopyroxene peridotites are represented by Wo (45–50) and display narrow Mg# (92.4–96.4). The Cpxs contain average amounts of Al₂O₃ (1.07 wt%), Cr₂O₃ (0.31 wt%), FeO (1.99 wt%) and higher amounts of CaO (23.94 wt%). Increased amounts of Al₂O₃ and Na₂O (samples of D01-5-3-1, D01-5-3-2) may be attributed to the neogenic pargasite in diopsides. In the pyroxene classification diagram (Figure 11A), nearly all Hongshishan clinopyroxenes fall within the diopside and show similar compositional variations with those of Poshi complexes and Alaskan-type complexes (Su et al., 2013). In the diagram of Al₂O₃ vs. Mg# (Figure 11B), most clinopyroxenes from Hongshishan complexes that belong to the second Cpx are similar to those in Alaskan-type intrusions (Khedr and Arai, 2016).

Table 1 shows PGE concentrations for all rock types. Total chromite PGE concentrations varied from 17.41 to 218.90 ppb. Just like most ophiolitic podiform chromites, which contain Os, Ir and Ru levels between 0.1 and 0.01 times chondritic (Leblanc 1991) and lower concentrations of Pt and Pd, all samples, except Cr9-11, showed an enriched Ir-subgroup (IPGE = Os, Ir, and Ru) and a depleted Pd-subgroup (PPGE = Rh, Pt, and Pd). The (Pd/Ir)_N ratio of the chromites ranged from 0 to 0.06 (except Cr9-11, Pd/Ir = 8.26) and the chondrite-normalized spider diagram showed steep right-facing sloped patterns, mostly similar to those of the PGE-rich chromites of the Wadi Al Hwanet ophiolite in Saudi Arabia (Ahmed et al., 2012) and the chromitites of the Luobusa ophiolite in Tibet (Zhou et al., 1996) (Figure 12).

Petrologically and mineralogically, the Hongshishan complex shows similarities with the Xiadong complex in the Middle Tianshan Terrane (Su et al., 2012; Su et al., 2014) and the Tuerkubantao intrusion in West Junggar (Deng et al., 2015a; Deng et al., 2015b). The Hongshishan complex contains no cross-cutting and intrusive rock unit relationships with distinct geochemical features, this implied the complex was probably not formed by multiple magmatic pulses (Su et al., 2012, 2014). The zonal textural features demonstrated the dunitic core resulted from crystal mush intrusions of deeper-seated cumulate fractionations (Tistl et al., 1994). In our study, compared with Alaskan-type complexes (Irvine, 1974; Himmelberg and Loney, 1995; Krause et al., 2007; Su et al., 2012), the Hongshishan complex has a concentrically zoned structure characterized predominantly by olivine, chrome spinel, clinopyroxene, and hornblende and an absence of orthopyroxene and plagioclase. Nearly all chrome spinels were less hydrous and were free of fluid inclusions, which implied the absence of a magmatic origin of chromite-bearing peridotite silicates in hydrous parental melts or scarce hydrous melts trapped during spinel growth.

Based on field observations, petrography and major element composition variations, the mineral crystallization sequence of the complex proceeded as follows: early cumulus olivine + chrome spinel formed chromite-bearing dunite, followed by crystallization of olivine + chrome spinel + clinopyroxene in spinel-poor clinopyroxene peridotite, and plagioclase + Cpx crystallized to form gabbro (Green et al., 2004; Krause et al., 2007; Khedr et al., 2020).

Generally, the Mg# values of the silicates indicated a fundamental range that followed normal fractional crystallization trends. Moreover, dunites have the highest levels of Ni, Cr, and Mg as compared to other complex units, and probably represented early magmatic products with olivine fractionation (Himmelberg and Loney, 1995). Significantly, from the ultramafic core to the mafic rim, through pyroxenite and the gabbroic margin, the major oxides of the Hongshishan complex have systematically negative correlation trends between major oxides (Al₂O₃, SiO₂, CaO) with Mg# and positive correlations between Mg# with MgO, NiO, and Cr₂O₃; these represent an accumulation of spinel and mafic minerals and the change of fractional crystallization degrees from the ultramafic core outwards from a common parental magma chamber of Uralian–Alaskan-type complex (Himmelberg and Loney, 1995; Farahat and Helmy, 2006; Habtoor et al., 2016).

For Uralian–Alaskan-type complexes, decreases in Fo and Cr# in the coexisting chromite helped monitor the early olivine and chromite crystallization stages. Olivine fractionation was significant during the early stage of magma ascension and evolution, as Mg preferentially partitioned into olivine accumulation instead of coexisting in a melt or other silicate minerals (Green et al., 2004; Teng, 2017). Crystallization of large olivine amounts to form an olivine-rich peridotite enriched the parental melt with Cr, Fe, and Al and triggered the crystallization of Fe, Al-rich chromite to form chromite. Crystallization of the chromite, and vice versa, increased magma Mg levels. High Fo values in olivine and the high Cr/(Cr + Al) in the spinel implied a parental magma rich in MgO but poor in Al₂O₃ (Krause et al., 2007). Additionally, increased Fe and Mn levels in dunite olivines to the one in the spinel-poor clinopyroxene peridotite likely relates to fractional crystallization (Khedr et al., 2020). Diagrams of major oxides versus Fo for olivines in dunite and clinopyroxene peridotite from the Hongshishan complex appeared within the Alaskan Global trend fields and showed a typical fractional crystallization trend similar to olivines in typical Alaskan-type complexes (Figure 10).

The Chondrite-normalized rare Earth element patterns are shown in Figure 7. All samples showed highly variable concentrations of HREE but slight enrichments of total REE levels and other incompatible elements relative to primitive mantle values from the ultramafic core to the marginal gabbroic rocks. The strong positive anomalies of Ba and Sr and the slight positive Eu anomaly from gabbros and pyroxenite whole rock samples supported a cumulative origin, i.e., accumulation of plagioclase. The REE patterns of the peridotites depended on the melting degree and mineralogy involved when the initial residue and enriched melt were formed. Residue composition formed by varying degrees of non-modal fractional melting of the same source. LREE-enriched patterns developed if melting occurred in the spine1 stability field. U-shaped REE patterns represent mixing of these residues with different degrees of incipient melts (Song and Frey, 1989). Additionally, serpentinization, carbonation, and hydrous fluid metasomawotism by LREE-enriched melts act as primary causes of LREE/HREE fractionation, LREE enrichment in carbonated peridotite rocks (Becker et al., 2001; Tian et al., 2011; Boskabadi et al., 2020), and subduction-related fluid, melts necessarily account for LILE-enrichment and HFSE-depletion of peridotite rocks in this complex.

Our preliminary investigation, combining the current field survey with these new observations, revealed that the petrological and mineralogical features of the Hongshishan complex were comparable with a Ural–Alaskan type complex.

Even though interactions with late interstitial, percolating melts, fluids, and subsolidus equilibration with neighboring grains can modify the elemental concentration in minerals as well as the mineral composition in cumulate rocks (Krause et al., 2007; Murphy, 2013); however, we did not observe any evidence of percolation and melt-rock reactions in the field or on a microscopic scale. Additionally, even though crustal contamination potentially increases Th/Yb ratios and reduces Nb/La, Ta/La, and Nb/Th ratios (Neal et al., 2002; Pearce, 2008), negative Eu anomalies and REE fractionation may imply crustal assimilation. Because the trace and REE element patterns of ultramafic cores and the marginal gabbros have sub-parallel trends, it implied that the different rock types of the Hongshishan complex are comagmatic, and initially generated by fractional crystallization from a common basaltic melt. The negative Eu anomalies of dunite accompanied by LREE enrichment were interpreted by a disequilibrium melting model, when and where, melting began in the garnet lherzolite facies through the spinel facies and concluded in the plagioclase facies. Phases in the source melted; however, the melt did not equilibrate with residual minerals. The negative Eu anomalies meant that some partial melting occurred in the plagioclase facies. Disequilibrium melting of plagioclase produces a residue with a negative Eu anomaly because of the positive Eu anomaly caused by suppressed crystallization of plagioclase (Prinzhofer and Allègre, 1985; McDonough and Frey, 1989; Wang et al., 1996). In addition, the absence of plagioclase in dunite and clinopyroxene peridotite likely caused the negative Eu anomaly due to low Sr concentrations, and subsequent serpentinization, carbonatization and metasomatism related subduction resulted in the higher total REE abundance, LREE enrichment with negative dunite Eu anomalies.

Cr-spinels in Hongshishan dunites showed ¹⁸⁷Os/¹⁸⁸Os isotopic ratios from 0.1251 to 0.1274 and ¹⁸⁷Re/¹⁸⁸Os ratios from 0.0066 to 0.0842. The Re/Os isotopic compositions of Cr-spinel samples appeared in the chromite field, which represents a residual peridotites trend. Just like a Gaositai complex with no significant crustal contamination in North China (Tian et al., 2011), the mean ¹⁸⁷Os/¹⁸⁸Os ratio for the chromites from Hongshishan showed that magmas originated from the mantle. The radiogenic Os isotopic compositions of the chromites meant the parental magmas of the Hongshishan complex suffered minor contamination coming from crustal components during magma volution. So crustal assimilation may not be necessary to significantly change the geochemical compositions of the parental magmas (Reiners et al., 1996; Batanova et al., 2005; Burg et al., 2009; Su et al., 2014).

Tectonically, some researchers have proposed that Alaskan-type complexes formed at subduction zones, representing arc magmas (Taylor, 1967; Irvine, 1974; Helmy and El Mahallawi, 2003), arc-root complexes (DeBari and Coleman, 1989; Brugmann et al., 1997; Helmy et al., 2014), at the change from the arc setting to the extensional regime (Tistl et al., 1994; Mues-Schumacher et al., 1996; Chen et al., 2009; Tian et al., 2011; Spandler and Pirard, 2013; Helmy et al., 2015) or even related to mantle plume (Ishiwatari and Ichiyama, 2004; Pirajno, 2004; Farahat and Helmy, 2006).

The Alaskan-type complexes of Xiadong in Central Tianshan and Tuerkubantao in West Junggar are arc-related, and interpreted as the product of partial melting of metasomatized lithospheric mantle triggered by subduction of an oceanic plate or slab window created by the subduction of an oceanic ridge (Xiao et al., 2010; Su et al., 2012; Su et al., 2014; Deng et al., 2015a; Deng et al., 2015b). In the early Carboniferous, Hongshishan, Liutuoshan areas might have evolved into an initial small oceanic basin (Zuo et al., 1990a; Zuo et al., 1990b; Zhao et al., 1994; Wang et al., 2014), but had not reached the degree of a mature mid-ocean ridge (Gong et al., 2003; Huang and Jin, 2006a; Huang and Jin, 2006b; Shi et al., 2017) because the geochemistries and age data came from basalts and gabbros in the Carboniferous Lvtiaoshan Formation, characterized by the occurrence of the bimodal volcanic rocks and lack of data on the mafic and ultramafic rocks in the complex (Peng et al., 2016). Virtually, the Hongshishan ultramafic-mafic complex lacked mantle peridotite, mafic cumulate, as well as pillow lava, so the ultramafic and mafic rocks were not the member compositions of a typical ophiolite (Yang et al., 2010; Wang et al., 2013; Peng et al., 2016).

We deduced that subduction or an arc magma setting could account for the upwelling of the depleted asthenosphere and rapid ascending magma flow which was less hydrous or free of opaque inclusions. Os isotopic compositions of chromian spinels in chromite suggested minor magma crustal contaminations. Formation of chromite was thought of as related to the evolution and chemical composition change of the parent magma (Tian et al., 2011). We attributed the enriched Fe³⁺ chromite compositions, LILE-enrichment, and HFSE-depletion of peridotite rocks and the neogenic clinopyroxenes in the clinopyroxene peridotites to suffering subduction modification and enrichment.

In the Alaskan-type complexes, chromites occurred as bands, seams, massive, veined or disseminated, and were accompanied by cumulate dunites (Himmelberg and Loney, 1995; Garuti et al., 2002; Garuti et al., 2003; Krause et al., 2007; Tian et al., 2011; Khedr et al., 2020). Cumulus olivines in host dunites and banded chromites with varying degrees of disseminated chromian-spinel grains were the typical characters of fractional crystallization (Khedr et al., 2020). Chromites may enter clinopyroxene crystals in an isomorphic form instead of an olivine form and crystalize in a chromian-spinel form during cumulation of olivine, so chromites form in cumulate dunite rather than in clinopyroxene-riched rocks. As the olivines fractionally crystallized, chromite compositions and relatively volatile components concentrated in residual magma. The emergence of high concentrations of volatiles largely delayed melt crystallization times and promoted the formation of ore pulp from ore-forming components as a massive and veined chromite orebody. Thus, those two different processes accounted for chromite formation, where the former crystallized from melting and released chrome-bearing minerals, which accumulated through fractional crystallization. In the latter, ore pulps carried by volatiles cut across the earlier disseminated chromite ores. The textures in Hongshishan chromites show that the disseminated chromite crystallized earlier than the massive and veined chromites. The Uralian-Alaskan-type chromites are characterized by the predominant composition of Cr³⁺→Fe³⁺ substitution, which indicated an oxygen fugacity (fO₂) variation during chromite precipitation (Johan, 2006). Crystallization and separation of Alaskan-type magma at high oxygen fugacity promoted transformation of the magma Fe²⁺ components into Fe³⁺, which impeded the combination of FeO with MgO and SiO₂ to produce poor a CaO-poor orthopyroxene, which instead occurred as magnetite and dispersed throughout the magma to form an iron-rich cumulatite.

Alaskan-type complexes are well-known as sub-economic Cu–Ni–PGE mineralization hosts, other than chromite deposits (Helmy and Mogessie, 2001; Helmy, 2004). However, the parental magmas of Alaskan-type complexes are water-rich/hydrous with high oxygen fugacities and less crustal assimilation, which might explain the absence of Ni–Cu sulfide mineralization that requires a reducing environment for most Alaskan-type complexes (Pettigrew and Hattori, 2006; Su et al., 2013). The Hongshishan complex contains a unique Ir-Ru-rich chromite deposit along the southern border of the Altaids orogenic belt. The first PGE-bearing sulfides crystallized within the early cumulus olivine and chromian spinels as primary inclusions with variable but generally high levels of PGE (17.41–218.90 ppb). They display similar patterns with IPGE enrichment (Os, Ir, Ru), PPGE depletion (Pt, Pd), and Ni/Cu depletions. Partial melting of the upper mantle controls PGE levels and their distribution in mantle rocks, and the migrating ability of IPGE is significantly weaker than PPGE (Wood, 1987; Zhou et al., 2014). IPGEs are compatible elements in monosulfide solid solutions (MSS), while PPGEs are incompatible with MSS when MSS forms by fractional crystallization of sulfide melts. High levels of IPGEs occurred in MSS whereas PPGE increased in residual sulfide melts, leading to the differentiation of PGE (Mungall et al., 2005; Cui et al., 2020).

PGE-bearing chromite mineralization in addition to the Hongshishan complex has not been reported in a mafic-ultramafic complex along the southern border of the Altaids orogenic belt. The identification of the Hongshishan Alaskan-type complex indicated that this kind of complex has great potential for chrome-PGE prospecting in the Beishan orogenic collage.