Petrogenesis and Tectonic Implications of the Qulihai Pluton in the Northern Margin of the Yili Block, NW China

The Central Asian Orogenic Belt (CAOB) is the largest Phanerozoic accretionary orogenic belt on Earth between Eastern Europe, Siberia, Tarim, and North China Cratons [1, 2] (Figure 1(a)). It is the product of long-term subduction and reduction of the Paleo–Asian Ocean. CAOB formation involves the assembly of numerous microblocks, island arcs, oceanic islands, and accretionary complexes, as well as the production of massive magmatic rocks [3, 4]. The CAOB is the most significant area of Phanerozoic continental crustal growth on Earth [1-3, 5-9].

The Tianshan Orogenic Belt, an important part of the CAOB, extends from Uzbekistan in the west to Xinjiang and Gansu in China in the east [9]. The ophiolites distributed along Jinghe, Kuitun River, and Bayingou may represent the remnant of the North Tianshan Ocean during the late Paleozoic [10-12]. The ophiolites developed in the Southern Tianshan Suture Zone can be divided into north and south belts based on their locations: the ophiolites in the northern belt are mainly distributed along the northern boundary of the Southern Tianshan in the area of Uwamen–Guluogou–Dalubayi–Qiongkushitai, while those in the southern belt are exposed along the Tonghua Mountain–Yushugou–Serikeya–Kule Lake area. These ophiolites are mainly regarded as products of the South Tianshan Oceanic Basin [6, 13-16].

However, the closure time of the North Tianshan Ocean remains debatable. Some researchers suggest that its closure occurred at the end of the Devonian and that the Carboniferous-early Permian volcanic rocks were formed related to the super-mantle plume [17, 18], but others believe that the volcanic rocks of the Dahalajunshan Formation exposed in the northern Yili Block (YB) display geochemical characteristics of continental arcs, instead of rift volcanic rocks [19-26]. In addition, the age statistics of magmatic rocks show that the Paleozoic magmatism at the northern margin of the West Tianshan Mountains is episodic. It lacks magmatism between 340 and 320 Ma [27]. Based on this phenomenon, some scholars have suggested that the North Tianshan Ocean subducted at a low angle beneath the YB during 340–320 Ma and that the eruption of magmatism in the context of extension after 320 Ma may be related to the roll-back of the subducted slab [28].

The recognition mentioned above demonstrates the complexity of the Paleozoic tectonic evolution in the northern West Tianshan Orogen (WTO). Paleozoic intrusive rocks with ages of 477–288 Ma abound in the WTO [29-38], with the Hercynian granites being the most widely distributed and largest in scale [39]. Granites formed in different tectonic processes, including subduction, collision, and postcollision, have distinct geochemical features; thus, research on granites is significant in revealing the orogenic process [40, 41].

Here, we present petrology, geochemistry, zircon U–Pb dating, and Lu–Hf isotope studies on the Qulihai pluton in the northern margin of the YB and discuss its petrogenesis and geological significance to provide new evidence on the tectonic evolution of the northern WTO.

According to previous studies, the North and South Tianshan oceans existed to the north and south of the Yili–Central Tianshan Block, respectively, during the Paleozoic. Due to the subduction of the North and South Tianshan oceans, the Tarim Craton and Junggar Terrane eventually collided with the Yili–Central Tianshan Block on the north and south sides, respectively, forming the Tianshan Orogenic Belt. With 88°E as the boundary, the Tianshan Orogenic Belt in China is divided into two parts: the East Tianshan Orogen and WTO [42, 43]. The WTO is situated between the Junggar Terrane and Tarim Craton and can be subdivided into three tectonic units from north to south: the North Tianshan Accretionary Complex (NTAC), the Yili–Central Tianshan Block, and the South Tianshan Belt [8, 44, 45] (Figure 1(b)).

The NTAC strikes near the NWW-SEE direction, separated from the Yili–Central Tianshan Block by the North Tianshan Fault to the south, and thrusts northward over the southern Junggar Basin with stacked thrusts [44]. The NTAC comprises Middle Devonian andesite–dacite–rhyolite association and clastic rocks [46], Carboniferous sandstones, limestones, conglomerates, volcanic lavas, tuffs, and late Devonian–early Carboniferous ophiolites [47]. Among them, the Devonian–Carboniferous sequence is mainly turbidites [47], and abundant late Devonian–early Carboniferous radiolarians and conodonts were found in the siliciclastic rocks [15]. Wang et al. [45] proposed that turbidites may have been deposited in a fore-arc basin in the late Carboniferous. The ophiolites are exposed as blocks within the turbidites. The typical ophiolite section is in the Bayingou area, comprising mantle peridotite, sheet dyke, gabbro, diabase, pillow lava, and siliciclastic rocks [10, 11, 15, 48]. The zircon U–Pb ages of gabbro and plagiogranite from Bayingou and Jinghe ophiolite were 344–325 and 382 Ma, respectively [10, 11, 49]. The Permian and post-Permian strata are mainly in fault contact with Carboniferous strata in the NTAC [50]. The depositional environment changed gradually from marine to terrestrial in the early Permian and became typical terrestrials of massive molasses in the late Permian [51].

The YB is a triangle-shaped area that wedges out eastward, which might be originally separated from the Tarim Craton [52-54]. Precambrian rocks are predominantly distributed on the north and south margins of the YB, while the central area is covered by massive late Paleozoic–Cenozoic sediments [44, 47]. Except for the Precambrian metamorphic crystalline basement rocks, Cambrian-Lower Ordovician clastic rocks and carbonates and middle Ordovician-Lower Silurian arc-type calc-alkaline volcanic rocks were also exposed on the northern margin of the YB [44]. The Middle–Upper Silurian is mainly a flysch deposition, composed of siltstone, siliceous rocks, and mudstone. The Devonian contains a suite of clastic deposits unconformably overlying the pre-Devonian strata. Carboniferous volcanic and sedimentary rocks are extensively exposed on the northern margin of the YB, primarily composed of tuffs, sandstones, shales, and volcanic rocks [9]. The northern margin of the YB was in intraplate settings in the Permian, while the depositional environment was gradually transformed from marine to terrestrial. It mainly exposed terrestrial molasse construction [51, 55]. Some Paleozoic intrusive rocks are exposed on the northern margin of the YB, including large amounts of granite [34, 56]. The Hercynian granites are intruded into the Upper Paleozoic strata, which are constrained by the surrounding faults with an NWW-SEE directional distribution [57].

The studied Qulihai pluton intruded into the Paleozoic strata on the northern margin of the YB with an ellipse shape and spread along the east–west direction, which is about 26 km from east to west and 10 km from north to south (Figure 2). It primarily contains gray-black fine-grained quartz diorite, gray fine-grained quartz monzonite, and flesh-red syenogranite. The surrounding strata include Devonian and Carboniferous sequences [39]. The upper Devonian Tuosikuerta Formation mainly features muddy siltstone, mudstone, and minor tuff. The Carboniferous strata include the Lower Carboniferous Dahalajunshan Formation and Upper Carboniferous Naogaitu Formation. The Lower Carboniferous Dahalajunshan Formation contains calc-alkaline intermediate-acid volcanic clastic rocks and intermediate-basic lavas, representing an active continental margin environment. The Upper Carboniferous Naogaitu Formation is predominantly acidic pyroclastic rocks intercalated with sandstones and sand conglomerates (Figure 2). According to field investigation and petrographic studies, the samples collected in this study can be classified into three types: quartz diorite, quartz monzonite, and syenogranite. The quartz diorite is in the central part, surrounded by quartz monzonite and syenogranite. The quartz diorite also occurs as enclaves in the syenogranite at the boundary (Figure 3(a) and 3(b)).

The gray-black quartz diorite is fine-grained and has a massive structure, mainly comprising plagioclase (40%–45%), amphibole (20%–25%), quartz (15%–20%), K-feldspar (5%–10%), and biotite (5%). The plagioclase grains are mostly euhedral of various sizes, presenting a plate and flake in shape, while some of them are altered to clay minerals. The amphiboles range from 0.3 to 0.5 mm, mostly euhedral in shape, while partly chloritized. The quartz grains range from 0.05 to 0.2 mm with anhedral in shape and surface clean. The K-feldspars are generally ca. 50 µm in length with euhedral shapes. They have Carlsbad twins and are partly serialized. The biotite grains are anhedral in shape (Figures 3(c) and 3(d)).

The quartz monzonites are deeply gray and fine-grained, primarily containing plagioclase (35%–40%), K-feldspar (30%–35%), quartz (10%–15%), biotite (5%–10%), and amphibolite (5%–10%). The plagioclase grains range from 0.2 to 0.6 mm with a euhedral shape. Some of them are altered, and some grains develop polysynthetic twins. The K-feldspars are 0.2–0.6 mm in size with a sheet and plate shape. They developed Carlsbad twins, and some are serialized. The quartz grains are mostly anhedral with a clean surface. The biotite grains are dark brown-light brown with an anhedral shape. The amphiboles are mostly anhedral in shape, presenting as flaky and granular. Some of them are chloritized (Figure 3(e)).

The flesh-red syenogranite is medium fine-grained to medium coarse-grained, mainly featuring K-feldspar (50%–55%), quartz (20%–25%), plagioclase (10%–15%), minor biotite (5%), and chlorite (5%). The quartz grains range from 0.3 to 0.5 mm and are partly smaller than 0.1 mm. The surface of quartz grains is clean, showing wavy extinction. The K-feldspars range from 0.3 to 2 mm with a euhedral shape, mostly presenting a sheet shape. The plagioclase grains range from 0.5 to 2 mm with a euhedral shape. Some have polysynthetic twins and ring-band structures, while most are sericinated. The biotite grains are anhedral in shape, presenting long strips and dark brown in interference color. The chlorite grains present scaly aggregates in shape, and some of them contain tiny quartz grains, while some parts are biotite alteration products (Figure 3(f)).

Whole-rock geochemical composition studies were conducted at Wuhan Sample Solution Analytical Technology Co., Ltd., Hubei, China. The compositions of the major elements were analyzed using a Japan Primus II X-Ray Fluorescence Spectrometer (XRF-1500). Loss on ignition was determined by the weight method, and the analytical precision was generally better than 2%–3%. Trace and rare Earth elements were determined by Agilent 7700e ICP-MS, and the samples were analyzed by acid dissolution methods. The analytical precision (according to GSR-1 and GSR-2 national standards) was better than 5% for the element mass fraction exceeding 10 × 10⁻⁶ and better than 10% for the element mass fraction below 10 × 10⁻⁶. The detailed analytical procedures were described by Zhou et al. [58].

The zircon selection was performed at Xi'an Ruishi Geological Technology Co., Ltd. First, enriched zircons were selected by crushing, screening, and electromagnetic methods, after which zircons with relatively well crystalline shapes, less cleavage, and fewer internal inclusions were selected under the microscope. The selected zircons were embedded in epoxy resin and then polished down until their surface was exposed. Subsequently, transmission and reflection light and cathodoluminescence (CL) images were taken by electron microscopy. Representative zircon grains and areas were selected for U–Pb dating based on transmission, reflection light, and CL images.

The in situ U–Pb dating of the zircons was completed at the Institute of Geology and Mineralogy, Tianjin Geological Survey Center. A Neptune LA-MC–ICP-MS equipped with a 193 nm ArF excimer laser-ablation system was used for the analysis. The analytical spots were 35 µm in diameter, and helium was used as the carrier gas for the ablation material. Detailed analytical methods and equipment parameters were described by Li et al. [59]. TEMORA was used as the external standard zircon for zircon U–Pb isotope dating. The trace element content of zircon was determined using NIST 610 as the external standard and ²⁹Si as the internal standard. The isotope ratio and elemental content analysis of the samples were processed using ICPMSDataCal. The weighted average age was calculated, and zircon age concordia diagrams were generated using Isoplot (ver3.00) [60].

In situ zircon Lu–Hf isotope analyses were performed at the Key Laboratory of Orogenic Belts and Crustal Evolution, Ministry of Education, School of Earth and Space Science, Peking University. The instrument used was an MC-ICP-MS equipped with a GeoLas HD 193 nm ArF excimer laser. The test spot was near the U-Pb isotope data analysis point. The test was conducted in single-point exfoliation mode, and the analytical spot was 50 µm in diameter. Helium was used as the carrier gas for the exfoliation material, and the international standard zircon 91500 and Plesovice were used as the standard samples for analysis. The analyzed ¹⁷⁶Hf/¹⁷⁷Hf value of 91500 was 0.282306 ± 27 (2σ, n = 8), which is consistent within the error range with the reference value of 0.282307 ± 0.000031 (2σ) [61]. The analyzed ¹⁷⁶Hf/¹⁷⁷Hf value of Plesovice was 0.282485 ± 23 (2σ, n = 8), which is also equal to the reference value of 0.282482 ± 0.000013 (2σ) [62]. The decay constant of ¹⁷⁶Lu used to calculate the initial ¹⁷⁶/¹⁷⁷Hf values was 1.865 × 10⁻¹¹ when calculating the Lu–Hf isotopic composition [63]. For εHf(t) calculation, we used a chondritic model with ¹⁷⁶Lu/¹⁷⁷Hf of 0.0332 and ¹⁷⁶Hf/¹⁷⁷Hf of 0.282772 [64]. Single-stage Hf model ages (T_DM1) were calculated using present day ¹⁷⁶Lu/¹⁷⁷Hf of 0.0384 and ¹⁷⁶Hf/¹⁷⁷Hf of 0.28325. The two-stage model age (T_DM2) was calculated using the average continental crust value ¹⁷⁶Lu/¹⁷⁷Hf of 0.015 [65].

Table 1 shows the major and trace element compositions of the studied samples.

The quartz diorite samples are characterized by a SiO₂ content of 58.99–60.43 wt.% and total alkalis (Na₂O + K₂O) of 7.25–7.83 wt.%, plotting in the diorite and subalkaline fields in the TAS diagram (Figure 4(a)). They have an Al₂O₃ content of 16.64–17.19 wt.%, a CaO content of 4.33–4.94 wt.%, and K₂O/Na₂O ratios of 0.35–0.68, belonging to the metaluminous, high-K calc-alkaline series (Figures 4(b) and 4(d)). Their MgO and Fe₂O₃^T contents range from 1.99 to 2.19 wt.% and 5.95 to 6.69 wt.%, respectively, corresponding to the Mg^# values of 43–44. The rare earth element (REE) content of the quartz diorite is 170–178 ppm. They show enrichment in light rare earth elements (LREEs) and depletion in heavy rare earth elements (HREEs) with (La/Yb)_N ratios of 7.61–10.1 and (Ga/Yb)_N ratios of 1.41–1.72 (Figure 5(a)). They display a right-dipping rare Earth distribution pattern, with moderate negative Eu anomalies (δEu = 0.69–0.87) in the chondrite-normalized diagram. Their primitive mantle-normalized trace element distribution patterns are characterized by prominent enrichment in large-ion lithophile elements (LILEs), such as Rb, Th, U, and Pb, but they are depleted in high-field strength elements (HFSEs), such as Nb, Ta, P, and Ti (Figure 5(b)).

The quartz monzonites yield moderate SiO₂ (64.91 and 65.27 wt.%) and total alkali (Na₂O + K₂O) contents of 7.73–8.86 wt.%, plotting in the quartz monzonite and subalkaline fields in the TAS diagram (Figure 4(a)). They have Al₂O₃ of 14.91–15.24 wt.%, CaO of 2.09–2.75 wt.%, and K₂O/Na₂O ratios of 1.21–1.31. They are metaluminous with A/CNK ratios of 0.9–1.08 and belong to the shoshonite series (Figures 4(b) and 4(d)). They have slightly varying MgO and Fe₂O₃^T concentrations of 1.62–1.88 wt.% and 4.50–4.90 wt.%, with corresponding Mg^# values of 45–47. The REE content of quartz monzonite is 214–231 ppm. They are enriched in LREEs rather than in HREEs, with (La/Yb)_N values of 9.28–11.4 and moderate negative Eu anomalies (δEu = 0.56–0.62; Figure 5(a)). On the primitive mantle-normalized trace element diagrams, the quartz monzonite samples display depletions in Nb, Ta, P, and Ti and enrichments in LILEs, such as Rb, Th, and U (Figure 5(b)).

The syenogranite samples have a high SiO₂ content of 75.14–76.45 wt.% and total alkali (Na₂O + K₂O) values of 7.02–8.72 wt.%. They are displayed in the granodiorite and subalkaline fields on the TAS diagram (Figure 4(b)). They have Al₂O₃, CaO, and K₂O/Na₂O contents of 12.42–13.10 wt.%, 0.19–1.03 wt.%, and 1.05–1.55, respectively, belonging to the metaluminous, high-K calc-alkaline rocks (Figures 4(b) and 4(d)). They have relatively lower MgO and Fe₂O₃^T values of 0.07–0.28 wt.% and 0.94–1.86 wt.%, respectively, corresponding to the Mg^# values of 12–31. The REE content of the syenogranite samples is 149–222 ppm. They are also enriched in LREEs with the (La/Yb)_N ratios of 6.39–12.3, with pronounced to moderate negative Eu anomalies (δEu = 0.26–0.54; Figure 5(a)). They display relative enrichment of Rb, Th, U, and Pb and depletion in Nb, Ta, and Ti on primitive mantle-normalized trace element diagrams (Figure 5(b)).

Generally, quartz diorite and quartz monzonite samples show consistent distribution patterns on the chondrite-normalized REE patterns and primitive mantle normalized trace element diagrams (Figure 5). In contrast, the syenogranite distribution pattern differs slightly from the former two (Figure 5). Overall, the three samples are enriched in LILEs, such as Rb, Th, U, and Pb and depleted in HFSEs with negative Nb, Ta, P, and Ti anomalies (Figure 5(b)).

A total of sixty zircon grains from the samples of the Qulihai pluton were analyzed for U–Pb dating. Figure 6 shows typical CL pictures of the analyzed zircon grains. The analytical results are listed in Table 2 and further illustrated in the concordia diagrams (Figure 7). All the analytical zircons displayed euhedral inshapes, with length/width ratios of 3:1–1:1. The CL images show that most of the zircons have magmatic oscillatory zoning (Figure 6).

Twenty zircons from the syenogranite samples were analyzed. Their Th and U contents are 103–827 ppm and 115–557 ppm, respectively, with Th/U ratios of 0.57–2.98, indicating an igneous origin. Twenty concordant analyses were obtained from twenty zircon grains. Spot TS18.36.11 and TS18.36.18 are relatively older, with ²⁰⁶Pb/²³⁸U ages of 391 and 446 Ma, respectively. The remaining eighteen zircons have U–Pb ages of 343–349 Ma, with a weighted average of 346 Ma (Figure 7(a)).

Nineteen concordant ages were obtained from the twenty analyzed zircon grains of quartz diorite. Spot TS18.38.12 and TS18.38.17 yielded ²⁰⁶Pb/²³⁸U ages of 330 and 351 Ma, respectively. The other seventeen zircons yielded ages of 337–346 Ma, with a weighted average age of 342 Ma (Figure 7(b)). The twenty zircons have Th contents of 57–687 ppm and U contents of 100–652 ppm, with Th/U ratios of 0.56–1.62, indicating their magmatic origin.

All twenty zircons analyzed for quartz monzonite are concordant. Their Th and U contents are 48–189 ppm and 101–217 ppm, respectively, with Th/U ratios of 0.47–0.87, suggesting an igneous origin. The twenty zircons have U–Pb ages of 340–349 Ma, defining a weighted average age of 344 Ma (Figure 7(c)).

We selected ten zircon grains with concordant U–Pb ages from syenogranite, quartz diorite, and quartz monzonite for Lu–Hf isotope analyses, respectively. The results are listed in Table 3 and illustrated in Figures 8 and 9. The εHf(t) values were calculated using the weighted mean ages as the crystallization ages of the zircons. The ¹⁷⁶Lu/¹⁷⁷Hf ratios for the thirty zircons ranged from 0.000549 to 0.002329, which represents the Hf isotopic composition of its system when formed [66].

The initial ¹⁷⁶Hf/¹⁷⁷Hf value of zircons from the syenogranite sample was 0.282665–0.282806, with the εHf(t) values varying from +3.3 to +8.4 (average +5.5; Figures 8(a) and 9) and T_DM1 and T_DM2 ages of 647–853 Ma and 817–1138 Ma, respectively.

For the zircons from quartz diorite, the initial ¹⁷⁶Hf/¹⁷⁷Hf value was 0.282659–0.282748, with the εHf(t) values of 3.3–6.3 (average +4.7; Figures 8(c) and 9). Their T_DM1 ages cluster between 727 and 856 Ma, with T_MD2 ages relatively concentrated between 949 and 1144 Ma.

Zircon grains from quartz monzonite have relatively homogeneous Hf isotope compositions, with initial ¹⁷⁶Hf/¹⁷⁷Hf values of 0.282637–0.282725 and εHf(t) values of 2.7–5.8 (average +4.2; Figures 8(e) and 9). The corresponding T_DM1 and T_DM2 ages were 741–867 Ma and 982–1182 Ma, respectively.

The quartz diorite, quartz monzonite, and syenogranite do not contain dark alkaline minerals such as fayalite. Their 10,000 × Ga/Al ratios are 2.47–2.57, 2.32–2.40, and 2.22–2.69, respectively, lower than the global average of 3.5 for A-type granites [67]. Almost all the samples of the Qulihai pluton plot into the field of non-A-type granites in the (K₂O + Na₂O)/CaO versus Zr + Nb + Ce + Y and Zr versus 10,000 Ga/Al diagrams (Figure 10(a) and 10(b)), indicating that all three type rocks are non-A-type granites. The quartz diorite, quartz monzonite, and syenogranite in the Qulihai pluton fall within the I- and S-type granite fields on the (K₂O + Na₂O)/CaO versus Zr + Nb + Ce + Y and Zr versus 10,000 Ga/Al diagrams (Figure 10(a) and 10(b)). They do not contain characteristic minerals of S-type granite such as cordierite. They plot within the I-type granite field on the Rb/Zr versus SiO₂ diagram (Figure 10(c)) and show the evolution trend of I-type granite on the Th versus Rb diagram (Figure 10(d)). In addition, they also contain minerals such as hornblende and biotite. The above features indicate that quartz diorite, quartz monzonite, and syenogranite belong to the I-type granite series. On the Harker diagram, all samples of Qulihai pluton display a negative correlation between P₂O₅ and SiO₂ content, similar to the I-type granites in the Australian Lachlan Orogenic Belt [68]. The above analyses indicate that the Qulihai pluton belongs to the I-type calc-alkaline granite series.

The εHf(t) values of syenogranite, quartz diorite, and quartz monzonite in the studied Qulihai pluton ranged from +3.3 to +8.4 (average +5.5), +3.3 to +6.3 (average +4.7), and +2.7 to +5.8 (average +4.2), respectively. All their samples fall into the field between chondrite and depleted mantle on the εHf(t) versus Age diagram (Figure 9), and their two-stage model ages are Meso-Neoproterozoic. According to previous studies, the high and positive εHf(t) values (up to +11.4) of the early Carboniferous granites in West Tianshan indicate the presence of a juvenile lower crust [69-71]. The formation of many late Paleozoic granites exposed in the YB was often attributed to the partial melting of the Precambrian basement that existed in the YB [72]. Their εHf(t) and εNd(t) values are generally positive [73, 74], which suggests that the young mantle material exerts significant effects during the formation of these granites. Furthermore, the northern margin of the YB experienced intense magmatic events during the Neoproterozoic [75], resulting in the widespread existence of Neoproterozoic magmatic rocks in the Tianshan area [76]. This is further evidenced by the presence of residual zircons of the Neoproterozoic period in some early Paleozoic magmatic rocks in the Tianshan area [32, 77]. Considering that these previous analyses are similar to the present analysis, we believe that the Qulihai pluton was probably derived from the Meso-Neoproterozoic juvenile crust.

The quartz diorites and quartz monzonites display similar evolutionary trends on the chondrite-normalized REE patterns and primitive mantle-normalized trace element diagrams (Figure 5), with similar the Hf isotopes. Based on the above analysis, we think quartz diorite and quartz monzonite have the same magmatic source. Conversely, the syenogranite pattern differs from those of quartz diorite and quartz monzonite on the chondrite-normalized REE patterns and primitive mantle-normalized trace element diagrams (Figure 5). Furthermore, it differs from quartz diorite and quartz monzonite in major and trace element contents and displays a discontinuous trend in some geochemical diagrams (Figures 11 and 12). Thus, syenogranite differs from quartz diorite and quartz monzonite in magma sources.

The quartz diorites and quartz monzonites are characterized by moderate SiO₂, high K₂O, and low Cr and Ni contents. They all fall into the amphibolite melt field on the Al₂O₃/(FeO^T + MgO + TiO₂) versus Al₂O₃ + FeO^T + MgO + TiO₂ and CaO/(MgO + FeO^T + TiO₂) versus CaO + MgO + FeO^T + TiO₂ diagrams (Figure 12). Their chondrite-normalized REE distribution patterns show no significant depletion of MREEs, suggesting their potential derivation from the dehydration melting of amphibolite. However, their Zr/Hf values range from 37.5 to 40.8, close to the primitive mantle (Zr/Hf ≈ 36.3) [16]. According to experimental petrological studies, an Mg^# value exceeding 40 indicates that the melt may involve mantle components [36, 78, 79]. The Mg^# values of quartz diorite and quartz monzonite are 43–44 (>40) and 45–47 (>40), respectively, suggesting input of mantle-derived magma. Nevertheless, their Mg^# values are below the pure mantle melt values, indicating that the mantle components in quartz diorite and quartz monzonite are relatively limited. Meanwhile, these samples also show the trend of magma mixing on the FeO^T versus MgO and Al₂O₃/MgO versus SiO₂/MgO diagrams (Figure 13). The above analysis suggests that quartz diorite and quartz monzonite may be derived from the partial melting of the Meso-Neoproterozoic juvenile crust, accompanied by the addition of some mantle-derived components.

The syenogranite samples have relatively higher SiO₂, K₂O, and very low Ni and Cr contents. On the Al₂O₃/(FeO^T + MgO + TiO₂) versus Al₂O₃ + FeO^T + MgO + TiO₂ diagram and CaO/(MgO + FeO^T + TiO₂) versus CaO + MgO + FeO^T + TiO₂ diagram, the syenogranite samples mainly fall within the field near the melting of graywackes, suggesting a dominant magma source of continental crust (Figure 12). The syenogranite has Nb/Ta, Zr/ Hf, and Th/Ce ratios of 9.86–13.4 (average 11.2), 25.7–38.1 (average 30.6), and 0.30–0.63, respectively, similar to the continental crust (Nb/Ta ≈ 11.4, Zr/Hf ≈ 35.8, and Th/Ce ≥ 0.2) [16, 80, 81]. The syenogranite samples display Rb/Sr ratios of 1.80–4.63 (average 3.42), consistent with the crustal-derived granites (Rb/Sr > 0.5). They have Mg^# values of 12−31. The above characteristics indicate that the syenogranite is derived from pure crustal source material. The syenogranite samples also display a magmatic mixing trend on the Al₂O₃/MgO versus SiO₂/MgO diagram and FeO^T versus MgO diagram, suggesting that it might be the acid end member of magmatic mixing (Figure 13). The above analysis shows that the syenogranite was derived from the partial melting of Meso-Neoproterozoic juvenile crust, without the participation of mantle source components.

The negative correlations between the SiO₂ content and MgO, CaO, Na₂O, TiO₂, and P₂O₅ contents on the Harker diagram suggest that they may have experienced some degree of fractional crystallization. The moderate to strongly negative Eu anomalies and the Harker diagram displayed a negative correlation of SiO₂ with Al₂O₃ and CaO, indicating the fractional crystallization of plagioclase (Figures 5 and 11). The positive correlation trend between K₂O and SiO₂ on the Harker diagram suggests the absence of substantial fractional crystallization of K-feldspar and biotite (Figure 11). Sr, P, and Ti depletion may indicate fractional crystallization of Ti–Fe oxide, plagioclase, and apatite [82, 83].

In summary, the quartz diorite and quartz monzonite of the Qulihai pluton are mainly derived from the partial melting of the Meso-Neoproterozoic juvenile lower crust with the participation of mantle components, while the syenogranite is the product of the partial melting of the Meso-Neoproterozoic juvenile crust.

The Paleozoic is a crucial period in the evolution of the northern margin of the YB. The North Tianshan Ocean, a part of the Paleo–Asian Ocean, was active during the Paleozoic. Studies have shown that the North Tianshan Ocean began to subduct beneath the northern margin of the YB to the south in the early Ordovician [84]. This coincides with the early Ordovician island arc volcanic rocks exposed in the Wenquan area and the magmatic zircon ages determined for the eastern part of the northern margin of the YB, as well as the age of the Tabale blueschist [9]. After the early Ordovician, the North Tianshan Ocean continued to subduct beneath the YB to the south. The late Paleozoic magmatic activity stages indicate that the North Tianshan Ocean subduction was extremely strong during the late Devonian–early Carboniferous [3, 20, 85, 86]. The Lower Carboniferous Dahalajunshan Formation volcanic rocks are the product of continuous southward subduction of the North Tianshan Ocean [33]. Furthermore, the zircon SHRIMP reveals that the age of the Bayingou ophiolite is 325 Ma, with radiolarians found in its siliciclastic rocks. These results indicate that the North Tianshan Ocean was not yet closed in the early Carboniferous [6]. In contrast, the early Permian purplish-gray continental volcanic-sedimentary rocks exposed in the Nileke area are typical bimodal volcanic suites [87]. The Permian and Carboniferous strata of the YB are generally in angular unconformable contact [50]. These findings suggest that the North Tianshan Ocean might close in the latest Carboniferous [43]. The north margin of the YB has transformed from the subduction to the extensional environment since the Permian, forming alkaline and bimodal volcanic rocks that are characteristic of the rift [88-90].

The zircon U–Pb ages of the quartz diorite, quartz monzonite, and syenogranite of the studied Qulihai pluton are 342, 344, and 346 Ma, respectively. They are enriched in LILEs, such as Rb, Th, U, and Pb and depleted in HFSEs, such as Nb, Ta, P, and Ti, suggesting that the Qulihai pluton has geochemical features consistent with arc magmatic rocks formed in the subduction setting. These features suggest that their formation was related to the southward subduction of the North Tianshan Ocean during the early Carboniferous.

Wang et al. [91] found that the zircon εHf(t) age, whole-rock εNd(t) age, and zircon saturation temperature of the late Paleozoic granites exposed in the northern margin of the YB changed during the early Carboniferous (~350 Ma). These features include (1) Carboniferous granites having higher zircon saturation temperatures than Devonian granites, (2) Carboniferous granites having significant positive εHf(t) and εNd(t) values, and (3) Carboniferous crustal thicknesses slightly thinner than Devonian crusts. Therefore, the subduction pattern of the North Tianshan Ocean was thought to change from advancing subduction to retreating subduction on the northern margin of YB during the early Carboniferous (~350 Ma). Our data also support the above perspective (Figure 14). Li et al. [12] considered that the granite ages were gradually younger from the southern to the northern margin of the Borohoro Mountains during the late Devonian–early Carboniferous. This may indicate the existence of retreating subduction in the North Tianshan Ocean. Li et al. [92] considered that the volcanic rocks of the Upper Carboniferous Yishijilike Formation show a “bimodal” feature and may represent the extensional stage after the continental collision. Some scholars also believe that the late Carboniferous magmatic rocks have the characteristics of island arc volcanic rocks, and the “bimodal” magmatic rocks may be the product of the extensional environment caused by the retreat of the subducting slabs [93].

Generally, retreating subduction results in an extensional setting that leads to the upwelling of the asthenosphere and generates significant heat, causing the magma to form under high-temperature conditions [94, 95]. The zircon saturation temperatures of the quartz diorite, quartz monzonite, and syenogranite of the studied Qulihai pluton are 760°C–771°C (average 766°C), 810°C–837°C (average 819°C), and 768°C–831°C (average 786°C), slightly exceeding those of the late Devonian–early Carboniferous granites of the Western Tianshan formed in the typical subduction setting (713°C –783°C) [96]. This variation suggests the presence of thermal anomalies during the Qulihai pluton formation in the early Carboniferous. This thermal anomaly event may be related to the extensional background caused by the plate retreat (Figure 15). Based on the above analysis, we infer that the Qulihai pluton was formed in the tectonic setting of the southward retreating subduction of the North Tianshan Ocean during the early Carboniferous.

1. The zircon U–Pb ages of quartz diorite, quartz monzonite, and syenogranite in the Qulihai pluton are 342, 344, and 346 Ma, respectively.

2. The Qulihai pluton is mainly metaluminous moderate-high potassium calc-alkaline, I-type granite with arc magmatic features. It was formed in an active continental margin tectonic environment, where the North Tianshan Ocean was subducted southward beneath the YB during the early Carboniferous.

3. The quartz diorite and quartz monzonite in the Qulihai pluton originated from the partial melting of the Meso-Neoproterozoic juvenile lower crust, accompanied by the addition of mantle-derived materials. The syenogranite was derived from the partial melting of the Meso-Neoproterozoic juvenile lower crust. The three samples experienced some degree of fractional crystallization in the late diagenetic process and finally formed the Qulihai pluton.

4. The samples from the Qulihai pluton are characterized by relatively high zircon saturation temperatures. Combined with the regional data, the North Tianshan Ocean may have been in retreating subduction during the early Carboniferous.

The data for this study can be downloaded directly from this manuscript or requested by contacting the authors.

The authors declare that they have no conflicts of interest.

This research was supported by the National Natural Science Foundation of China (42072267), the West Light Foundation of the Chinese Academy of Sciences (XAB2020YW03), the Fundamental Research Funds for the Central Universities, CHD (300102272103, 211427230085), Opening Foundation of State Key Laboratory of Continental Dynamics, Northwest University (19LCD06), and the Youth Innovation Team of Shaanxi Universities.