The Pennsylvanian Composite Volcanism in the Bogda Mountains, NW China: Evidence for Postcollisional Rift Basins

The postcollisional setting is distinctive and not integrated into the model of plate tectonics (e.g., [1]). However, how to distinguish the characteristic of postcollisional magmatism remains questionable [1, 2]. Liegeois [1] proposed that geochemical discrimination diagrams cannot characterize the postcollisional tectonic setting due to a variety of magma types generated coevally. The postcollisional magmatism is so complex, but they still share some common characteristics: (1) They mainly belong to the high-K calc-alkaline to shoshonitic series, of which the felsic volcanic rocks are usually peraluminous and have attributes of A2-type granites [3–8]. (2) Their magma sources are generated from the crust or lithospheric/asthenospheric mantle during the preceding subduction and collision period [1, 8, 9]. Moreover, the sources commonly contain a sizeable juvenile component which has similar geochemical and isotopic attributes to the lithospheric mantle (e.g., [1] and reference therein). (3) They are usually linked to large horizontal movements along major shear zones [1, 10–13]. The most typical examples of the postcollisional setting are Neogene Tibet (northern Tibet at Songpan-Ganzi, Kunlun, and western Qiangtang terranes, southern Tibet at Lhasa terrane; [1, 4, 6–8, 14, 15]).

The Central Asian Orogenic Belt (CAOB) is one of the largest Phanerozoic accretionary orogens which was formed by a long accretionary history that interacted through microcontinents, island arcs, seamounts, and accretionary wedges (Figure 1; [16]). In the CAOB, scholars suggested a Late Paleozoic postcollisional process, but the beginning time and the geodynamic mechanism are controversial [17, 18]. Some have believed that the subduction processes have continued to the Pennsylvanian [19–21] and subsequently transferred to the postcollisional extensional setting in the Permian [17, 22–24]. Besides, scholars also thought that the Tarim-Tianshan-Junggar ocean was still active, and the subduction processes have lasted to the Permian-Triassic [25–28]. Others have suggested that the Pennsylvanian is a tectonic transitional regime from arc to postcollision due to their transitional characteristics in geochemistry [20, 29–33]. Due to the controversy on the beginning time of the postcollisional process, the Late Paleozoic (especially the Pennsylvanian) geodynamic mechanism of the CAOB is disputed. The Late Paleozoic Bogda Mountains (BMs) in the southern CAOB have been interpreted to represent a subduction-related island arc [34, 35] or back-arc [13, 17], mantle plume [36], ridge subduction [37, 38], postcollisional extension ([39]; Zhang et al., 2014; [40]), and local strike-slip faulting [41, 42]. These disputes suggest the need to find more solid evidence to interpret the tectonic evolution of the CAOB.

The Liushugou and Qijiagou Formations of the BMs in the southern CAOB (Figure 2), belonging to the Late Paleozoic [41, 43], could give constraint on this controversial period. The completed successions, exposed along the western and eastern BMs, provided an excellent opportunity to uncover their petrogenesis and geodynamic background. Thus, this study reports new geochronological, geochemical, and isotopic data for the volcanic and sedimentary rocks to study their petrogenesis, and provides a model to explain the Late Paleozoic geodynamic background of the North Xinjiang, as well as that of the CAOB.

The CAOB is a large accretionary orogenic belt surrounded by the continent Baltica in the northwest, the Siberian Craton in the northeast, and the Tarim and North China Cratons in the south (Figure 1(a); [44]). The BMs, as the southern segment of CAOB, function as a pivotal tectonic belt separating the Turpan-Hami Basin from the Junggar Basin (Figure 1(b); [45]). The BMs, located in the northern part of the NTS (Figure 1(b)), are mainly comprised of Devonian to Quaternary sedimentary rocks with some igneous rocks [19]. Of these, the Carboniferous bimodal volcanic rocks are widely distributed [13, 46–48]. The sedimentary rocks consist mainly of fine-grade clastic rocks, limestones, and volcanic breccias, with no ophiolites [49]. The present average elevation of the BMs is above 4000 m, which is induced by the late Cenozoic India-Asia collision [50]; the Carboniferous strata were deposited in an intracontinental rift basin [36] or a subduction-related forearc or back-arc basin [13, 17, 34, 35, 41].

The Haxionggou section, located in the central peak of the western BMs, is composed of the Mississippian Qierguositao Group, and the Pennsylvanian Liushugou and Qijiagou Formations from bottom to top, with no Devonian or even older strata (Figure 2(a); [43]). The Mississippian Qierguositao Group is located mainly at the peak of the BMs and is not involved in this study. The Liushugou Formation is dominated by pillow basalts with interstitial limestones (Figures 3(a) and 3(b)), peperites (Figure 3(c)), vesicular basalts (Figure 3(d)), basaltic andesites, andesites, dacites, rhyolites, volcanic breccias, and tuffs, with a small amount of sandstones and siltstones. The Qijiagou Formation consists of a set of clastic and carbonate rocks with fossils of brachiopods, gastropods, corals, and sea lily stems. A typical Bouma sequence is developed at the bottom, and the whole formation is composed of marine sediments [51]. Notably, in the southern BMs of the Baiyanggou area, the unconsolidated limestone and magma clastic rocks are mixed to form peperites, and the lenticular limestones are distributed in pillow basalts [41].

The Qijiaojing section, located in the eastern BMs, consists primarily of the Mississippian Qijiaojing Formation, and the Pennsylvanian Liushugou and Qijiagou Formations from bottom to top (Figure 2(b); [43]). The Qijiaojing Formation, dominated by marine volcano-sedimentary construction, is composed mainly of marine sandstones, siltstones interbedded with shales, siliceous rocks, and basalts in the lower part of the siliceous rocks, and with massive or pillow basalts and rhyolites interlayered with acidic tuffs, sandstones, and shales in the upper part. There is also some rhythmic eruption cycle composed of basalts and rhyolites (bimodal volcanic suites) reported in the Qijiaojing area of the eastern BMs [46, 52]. The Pennsylvanian Liushugou Formation is dominated by volcanic rocks, marine clastic rocks, and volcaniclastic rocks with a small amount of volcanic lavas (Figures 3(e)–3(g)). The volcanic rocks mainly consist of basalts, dacites, and rhyolites, similar to a “bimodal” volcanic assemblage (Figure 3(e); [13, 52]). The Pennsylvanian Qijiagou Formation mainly consists of felsic volcanic rocks mixed with volcaniclastic rocks (Figures 3(h) and 3(i)). The marine biodetritus is distributed in sandstones and limestones (Figure 3(h)).

In the Haxionggou section, the volcanic rocks of the Liushugou Formation mainly include basalts, andesites, dacites, and rhyolites with some peperites (Figures 4(a)–4(e)). In the Qijiaojing section, the Liushugou Formation consists of bimodal volcanic rocks with minor or no andesites (Figures 4(f)–4(i)).

The basaltic rocks are greyish-green or black grey; have a massive, pillow or vesicular structure; and exhibit no porphyry or less porphyritic textures with 10–30% content of clinopyroxene and plagioclase (Figure 4(a)). The clinopyroxene phenocrysts are euhedral and are up to 2 mm long. The plagioclase phenocrysts are subhedral, have a long columnar shape, and are 0.2–1.0 mm long. The groundmasses consist of fine-grained plagioclases, clinopyroxenes, and opaque oxides (Figures 4(a) and 4(g)).

The gabbros are greyish-green and have a gabbro texture in which the basic plagioclases and clinopyroxenes are irregularly interlaced with each other (Figure 4(f)). Clinopyroxene phenocrysts are subhedral, granular shaped, and 0.2–1.0 mm long. Basic plagioclase phenocrysts are euhedral to subhedral, have a long columnar shape, and are 0.2–1.0 mm long. The plagioclase and clinopyroxene phenocrysts have similar grain shapes, probably the products of the same period of crystallization.

The andesites and basaltic andesites are grey or greyish-green and commonly show a porphyritic texture with clinopyroxene, plagioclase, and hornblende phenocrysts (Figure 4(b)). The clinopyroxene phenocrysts are euhedral to subhedral, and some crystals have epidote alteration. The plagioclase phenocrysts are subhedral and tabular shaped, and some crystals are altered. The hornblende phenocrysts are few, commonly euhedral, and 0.1–0.5 mm long. The matrix consists mainly of plagioclase, clinopyroxene, volcanic glass, and opaque oxides, some of which have a trachytic texture (Figure 4(b)).

The dacites are grey and greyish brown and show nonporphyritic to porphyritic textures with plagioclase, hornblende, and quartz phenocrysts (Figures 4(c) and 4(h)). The plagioclase phenocrysts are subhedral and 0.1–1.0 mm long, and some of them have a polysynthetic twin and a Carlsbad twin. The matrix is mainly composed of plagioclases, quartzes, and volcanic glasses.

The rhyolites are generally grey or greyish-white and have an apparent rhyolite structure, with nonporphyritic cryptocrystalline and porphyritic textures (Figures 4(d) and 4(i)). The phenocrysts consist mainly of quartzes and plagioclases, with a content of about 10%. The quartz phenocrysts are hypidiomorphic to xenomorphic, with a diameter of 0.1–0.5 mm. The plagioclase phenocrysts are euhedral, tabular shaped, and 0.1–1.0 mm long, with a polysynthetic twin and zonal texture. The matrix mainly consists of quartzes, plagioclases, and glasses.

Besides, there are also some peperites in the Liushugou and Qijiagou Formations which were formed by magma-sediment mingling (magma and unconsolidated limestone, [41]). The main components of the flowing slurry are plagioclases and quartzes, and the host sediments are mainly calcites. Some slurry particles have condensation edges and recrystallized calcite baking edges (Figure 4(e)).

The studied volcanic rocks were collected from the Haxionggou and Qijiaojing sections in the BMs (Figure 2). Two samples (16HXG-12 and 16 HXG-13) from the Haxionggou section and one sample (16QJJ-15) from the Qijiaojing section have been collected for Zircon U–Pb dating and Hf isotope analyses. Fourteen samples from the Haxionggou section (all of these are collected from the Liushugou Formation) and 27 samples from the Qijiaojing section (9 samples are collected from the Liushugou Formation and 18 samples from the Qijiagou Formation) have been collected for bulk-rock geochemical analysis and Sr–Nd isotope analyses.

Zircon U–Pb isotope analyses were conducted using a Laser Ablation Inductively Coupled Plasma-Mass Spectrometry (LA-ICP-MS) instrument (Agilent 7500c ICP-MS coupled with a 193 nm ArF excimer laser) at the Key Laboratory of Orogenic Belts and Crustal Evolution, Ministry of Education, Peking University. Detailed analysis procedures follow Yuan et al. [53]. Isotopic ratios and elemental contents of samples were calculated by the GLITTER 4.0 program [54]. Common lead correction was calculated by the program given by Andersen [55]. The weighted average age and Concordia plot were drawn using ISOPLOT 3 [56]. Zircon in-situ Lu–Hf isotopic analysis is based on the preceding zircon U–Pb isotope, and the test was conducted at the State Key Laboratory of Continental Tectonics and Dynamics, Institute of Geology, Chinese Academy of Geological Sciences using the Neptune Multicollector Plasma-Mass Spectrometer and New Wave UP-213 UV Laser Ablation Multiple Collector Inductively Coupled Plasma-Mass Spectrometer. See Hou et al. [57] for operating conditions and detailed analysis procedures of relevant instruments.

The 12 basaltic rocks, 11 andesitic rocks, and 18 felsic rocks were milled to a mesh size of less than 200 for elemental geochemical analysis. Major element determinations were performed by inductively coupled plasma optical emission spectrometry analysis at the China University of Geoscience (Beijing). Preliminary treatment and measurement of trace elements were undertaken at OBCE. Then, these bulk samples were analyzed by the VG Axiom multicollector, high-resolution ICP-MS. Detailed method descriptions follow Liu et al. [58].

Bulk-rock Sr–Nd isotope separation and purification of Rb, Sr, Sm, and Nd were completed in the OBCE mainly by the conventional ion-exchange method. The isotope ratios of Sr and Nd were measured by LA-MC-ICP-MS in OBCE. The ratios of ⁸⁷Rb/⁸⁶Sr and ¹⁴⁷Sm/¹⁴⁴Nd were calculated mainly based on trace test results of Rb, Sr, Sm, and Nd. Mass fractionation is mainly conducted by normalizing the tested ⁸⁷Rb/⁸⁶Sr and ¹⁴⁷Sm/¹⁴⁴Nd with ⁸⁶Sr/⁸⁸Sr (0.1194) and ¹⁴⁶Nd/¹⁴⁴Nd (0.7219), respectively. Standard rock sample BCR-2 (basalt; $87Rb/86Sr=0.704992±7$ (2σ, $n=94$⁠); $147Sm/144Nd=0.512634±1$ (2σ, $n=97$⁠)) was used to evaluate the degree of isolation and purification of Rb, Sr, Sm, and Nd.

In comparison, we also collected the major and trace element data from the Andean Arc, the Cascade Arc, the Lesser Antilles Arc, Hawaii, the Izu Arc, the East African Rift, and the Lau Back-Arc Basin for comparison. The data are collected from PetDB (http://www.earthchem.org/petdb) and GEOROC (http://georoc.mpch-mainz.gwdg.de/georoc/). Due to a large number of data, we performed principal component analysis (PCA). PCA is a statistical procedure that uses an orthogonal transformation to convert a set of observations of possibly correlated variables into a set of values of linearly uncorrelated variables called principal components (PCs). The detailed analysis software employed is PAST—Palaeontological Statistics [59, 60]; introductions follow Wang et al. [42]. The main parameters are set as follows: matrix—correlation; groups—between group; and missing values—mean value imputation.

Three felsic volcanic rocks were collected from the Haxionggou section for zircon LA-ICP-MS U–Pb dating. 16HXG-12 and 16HXG-13 were chosen from the western Liushugou Formation (Haxionggou section). 16QJJ-15 was selected from the eastern Liushugou Formation (Qijiaojing section). All these analytic data are listed in Supplementary Table 1. Almost all the zircons are euhedral to subhedral, show typical oscillatory growth zoning, and have $Th/Uratios>0.2$ (0.29–1.36), indicating a magmatic origin [61]. Nineteen analyses from 16HXG-12 (N43°49.596$′$⁠, E87°57.360$′$⁠) were concordant, yielding ²⁰⁶Pb/²³⁸U ages ranging from 317 Ma to 302 Ma with a weighted mean ²⁰⁶Pb/²³⁸U age of $310±2Ma$ (Figure 5(a); $MSWD=1.5$⁠, $n=19$⁠). This age was interpreted as the crystallization age of the dacite. All the twenty zircons from 16HXG-13 (N43°49.677$′$⁠, E87°57.247$′$⁠) yielded concordant ²⁰⁶Pb/²³⁸U ages ranging from 317 Ma to 290 Ma. Except for 4 zircons of preeruptive units (inherited zircon) and 3 zircons of posteruptive units, the rest of the analyses gave a weighted mean ²⁰⁶Pb/²³⁸U age of $301±3Ma$ (Figure 5(b); $MSWD=1.6$⁠, $n=13$⁠), which represents the formation time (syneruptive time) of the tuff. Twenty zircon grains from 16QJJ-15 (N43°40.979$′$⁠, E91°38.937$′$⁠) yielded concordant ²⁰⁶Pb/²³⁸U ages ranging from 371 Ma to 294 Ma. Except for 6 zircons of preeruptive units (inherited zircon) and 1 zircon of posteruptive units, the rest of the analyses gave a weighted mean ²⁰⁶Pb/²³⁸U age of $308±4Ma$ (Figure 5(c); $MSWD=1.4$⁠, $n=13$⁠), which was interpreted as the formation time (syneruptive time) of the tuff.

The results of zircon Hf isotopic data are listed in Supplementary Table 2 and shown in Figure 5. The zircons from 16HXG-12, 16HXG-13, and 16QJJ-15 have positive $εHft$ values which were mainly concentrated in the ranges from +8.0 to +9.8, +13.1 to +15.6, and +10.8 to +13.2, respectively (Supplementary Table 2). All the zircons have $TDM1$ model ages from 644 Ma to 324 Ma, and $TDM2$ model ages from 737 Ma to 326 Ma (Supplementary Table 2). The $TDM1$ model ages have a major peak at 624–554 Ma, and a secondary peak at ca. 405 Ma (Figure 5(d)). The $TDM2$ model ages have a major peak at 700–638 Ma, and a secondary peak at ca. 446 Ma (Figure 5(d)). After excluding seven incorrect data (16HXG-13-07, 16HXG-13-10, 16HXG-13-12, 16HXG-13-17, 16HXG-13-20, 16QJJ-15-05, and 16QJJ-15-13), the remaining 30 zircon grains have relatively lower $εHft$ values which plot below the depleted mantle evolutionary line (Figure 5(e); Supplementary Table 2).

The result of bulk-rock major and trace elements from the Liushugou and Qijiagou Formations in the BMs are presented in Supplementary Table 3. The samples of the Liushugou Formation from the Haxionggou section have SiO₂ contents ranging from 47.31 to 76.87 wt.%, classified as basalt, andesite, basaltic trachyandesite, trachyandesite, dacite, trachydacite, and rhyolite in the Total Alkalis and Silica (TAS) diagram (Figure 6(a)), and subalkaline basalt, andesite, dacite/rhyodacite in the Zr/TiO₂–Nb/Y diagram (Figure 6(b)). On the K₂O–SiO₂ and FeO^T–Na₂O+K₂O–MgO (AFM) diagrams, most samples plot in the high-K and medium-K calc-alkaline areas, and a minority plot in the tholeiite areas (Figures 6(c) and 6(d)). In the Harker diagrams, SiO₂ values correlate negatively with MgO, Al₂O₃, CaO, TiO₂, and Fe₂O₃ contents, with no clear correlations with Na₂O, P₂O₅, Sr, Ni, and Co contents (Figure 7). The basaltic rocks have moderate light rare earth element (LREE) enrichment with no Eu anomalies in the chondrite-normalized diagram (Figure 8(a)). On the N-MORB-normalized spider diagrams, the basaltic rocks show enrichment of large-ion lithophile elements (LILEs; e.g., Rb, Ba, Sr, and U) and depletion of high-field-strength elements (HFSEs; e.g., Nb and Ta) (Figure 8(b)). The andesites, dacites, and rhyolites have similar trace element patterns to the basaltic rocks which exhibit enriched LREEs and LILEs and depleted HREEs and HFSEs (Figures 8(c)–8(h)). Furthermore, the felsic rocks have a little more LREE fractionation and negative Eu anomaly (Figures 8(e)–8(h)).

The samples of the Liushugou Formation from the Qijiaojing section are composed mainly of basalts, trachybasalts, dacites, and rhyolites, forming as the bimodal volcanic rocks (Figure 6(a)). The basaltic rocks show moderate LREE and LILE enrichment with Nb depletion (Figures 8(a) and 8(b)). The dacites and rhyolites show LREE and LILE enrichment with Nb, Ta, and Ti depletion (Figures 8(e)–8(h)). The major and trace characteristics are similar to the Haxionggou section but have relatively higher REE contents (Figures 6–8).

The samples of the Qijiagou Formation from the Qijiaojing section have SiO₂ contents ranging from 47.31 to 76.87 wt.% and are classified as basalt, basaltic andesite, basaltic trachyandesite, trachyandesite, dacite, trachydacite, and rhyolite in the TAS and Zr/TiO₂–Nb/Y diagrams (Figures 6(a) and 6(b)). Moreover, they are mostly from the calc-alkaline series in the K₂O–SiO₂ and AFM diagrams (Figures 6(c) and 6(d)). The other major and trace characteristics are similar to those of the Liushugou Formation (Figures 7 and 8).

The bulk-rock initial ⁸⁷Sr/⁸⁶Sr ratios (⁠$ISr$⁠) and $εNdt$ values are calculated based on the results of the zircon U–Pb ages (Figure 9), and the results are listed in Supplementary Table 4.

The basaltic rocks of the Liushugou Formation from the Haxionggou section have high $ISr$ (0.7048–0.7058) and positive $εNdt$ values (from +6.7 to +7.6), and the felsic volcanic rocks have similar Sr–Nd isotopic characteristics to the basaltic rocks with $ISr$ values ranging from 0.7051 to 0.7059 and $εNdt$⁠) values from +3.0 to +8.0. The basaltic rocks of the Liushugou Formation from the Qijiaojing section have $ISr$ values ranging from 0.7035 to 0.7054 and $εNdt$ values from +7.7 to +8.1, and the dacites and rhyolites have $ISr$ values ranging from 0.7039 to 0.7066 and $εNdt$ values from +6.8 to +7.7. The basaltic rocks of the Qijiagou Formation from the Qijiaojing section have high $ISr$ (0.7038–0.7042) and positive $εNdt$ values (+6.1 to +8.0), and the felsic volcanic rocks have similar Sr–Nd isotopic characteristics to the basaltic rocks with $ISr$ values ranging from 0.7032 to 0.7044 and $εNdt$ values from +5.3 to +7.2.

The major and trace elements of representative tectonic settings and the Carboniferous-Permian basaltic rocks from the BMs were all collected for PCA analysis (Figure 10). PC1 accounts for 41.8% of the variance and shows high positive loadings for Fe₂O₃^T, MgO, and TiO₂ and negative loadings for SiO₂ and Al₂O₃ when ten major elements are used as variables (Figures 10(a), 10(c), and 10(e)). PC2 (34.0% of variance) gives distinct positive loadings for Na₂O, P₂O₅, and K₂O and negative loadings for CaO. When the trace elements are used as variables, PC1 (49.2% of variance) has strong positive loadings for Rb, Sr, Ba, Nb, Ta, Zr, Hf, and LREEs compared to the negative loadings of V and Cu. PC2 (21.5% of variance) yields positive loadings for Sc, Pb, and HREEs and shows negative loadings for Cu, Co, Ni, Cr, and Zn (Figures 10(b), 10(d), and 10(f)).

Compared to the data of representative tectonic settings (collected from PetDB and GEOROC), the Carboniferous-Permian basaltic rocks from the BMs share some similar features with the Andean Arc, the Cascade Arc, the Lau Back-Arc Basin, the Izu Arc, and the Lesser Antilles Arc which were considered as continental arc, back-arc basin, and island arc in some aspects (Figure 10). But the Pennsylvanian basaltic rocks are rather complex and have composite compositions (Figure 10). On the principal compositional biplot diagrams, the Mississippian basaltic rocks show the same characteristics as the arc-related basaltic rocks, which overlap with the Lau Back-Arc Basin, as well as with the Andean Arc, the Cascade Arc, and the Lesser Antilles Arc in large part (Figure 10). Although some of the Pennsylvanian basaltic rocks from the BMs plot in the areas of the Andean Arc, the Cascade Arc, the Lau Back-Arc Basin, the Izu Arc, or the Lesser Antilles Arc, none of the tectonic settings overlap completely. All these features reveal that the Pennsylvanian basaltic rocks have a wide variety of magma types, and the geochemical features probably were mobilized from the original source or decoupled from the tectonic setting.

The Carboniferous strata are subdivided as the Mississippian Qiergusitao Group (C₁qr) and the Pennsylvanian Liushugou (C₂l) and Qijiagou (C₂q) Formations. These formations mainly consist of sandstones, mudstones, limestones, bioclastic limestones, tuffs, basalts, and andesites intercalated with rhyolites. Also, they have different fossil assemblages developed in these rocks. The biological fossils include marine brachiopods, corals, gastropods, and crinoids, indicating a shallow marine depositional environment [40, 43, 51]. Moreover, the pillow lavas, limestones, and siliceous rocks, together with peperites, suggest that the volcanic and volcaniclastic rocks were the result of subaqueous volcanism (Figures 3 and 4; [41]). Besides, the Mississippian Qiergusitao Group is composed predominantly of shoreline sandstones and mudstones, with minor or no pillow lavas and peperites. Above the Qiergusitao Group, the pillow lavas and peperites began to appear in the Pennsylvanian Liushugou and Qijiagou Formations. Up to the upper part of the Qijiagou Formation, the clastic rock units show a conglomerate-sandstone-siltstone-siliceous mudstone sequence, with the characteristics of rhythmically bedded turbidite sequences [51]. Therefore, the Carboniferous succession in the BMs is deposited in a progressively deepening basin, from a shallow marine to a semideep marine environment.

To assess the influence of the postmagmatic alteration and crustal contamination before speculating as to their source is very important (e.g., [62, 63]). Basaltic rocks are a good indicator for their mantle source and petrogenesis (e.g., [64]). The studied basaltic rocks have relatively high LOI contents (2.58–8.22 wt.%), revealing that the effect of alteration cannot be ignored. PCA analysis suggests that Zr and Ti are some of the least mobile elements (e.g., [65]); MnO, Fe₂O₃^T, and TiO₂ are the immobile elements; and Na₂O, K₂O, and Al₂O₃ are the mobile elements (Figures 10(a), 10(c), and 10(e)). Likewise, REEs, Rb, Sr, Th, Nb, Ta, and Hf are immobile elements due to their positive correlations with Zr, while Ba, Cs, and V are mobile elements as they are inversely correlated to Zr in the PCA biplots (Figures 10(b), 10(d), and 10(f)). Thus, the mobile elements cannot be used for the study of petrogenesis. Still, the immobile elements (e.g., REEs, Zr, and Ti) can reflect their petrogenetic processes.

Crustal contamination generally causes a decrease in $εNdt$⁠, Nb/La ratios, and Nb/Th ratios but causes an increase in $ISr$ because crustal materials have low $εNdt$ and Nb values but have high $ISr$ and Th contents (Rudnick and Fountain, 1995). All the studied volcanic rocks have relatively homogeneous Sr–Nd isotopic compositions, and their $εNdt$ and $ISr$ values do not vary with Mg^#, SiO₂, and Al₂O₃ (Figure 11), which are inconsistent with crustal contamination. Also, the constant Nb/La and Nb/Th ratios irrespective of SiO₂ indicate insignificant crustal contamination.

Although the genesis of the basaltic rocks is controlled mainly by the mantle component due to their low SiO₂ (47.0–51.2 wt.%) contents and high Mg^# (mostly >50), the lower Cr (21.3–661 ppm), Co (26.4–56.6 ppm), and Ni (5.1–178 ppm) contents related to primary mantle-derived magmas [66] indicate that their parental magma has experienced various degrees of fractional crystallization. The negative correlations between MgO, CaO, Al₂O₃, Fe₂O₃^T, and SiO₂, together with the positive correlation between Cr and Ni, are compatible with the fractionation of clinopyroxene and olivine (Figure 7). The basaltic rocks have positive Eu and Sr anomalies (Figure 8) and have no correlation between Na₂O, K₂O, and SiO₂ (Figure 7), suggesting minor or no fractionation of plagioclase. There is no correlation between Dy/Yb and Mg^# further reflecting the little effect of amphibole fractionation ([67]; Supplemental Table 3).

The basaltic rocks have positive $εNdt$ (3.0–8.1) and $εHft$ (4.3–17.7) values with relatively lower $ISr$ values (0.703–0.707), suggesting that they were derived from an isotopically depleted mantle source (Figure 9). In the chondrite-normalized REE patterns and N-MORB-normalized trace element spider diagrams, these basaltic rocks show the geochemical affinity between oceanic island basalt (OIB) and enriched middle ocean ridge basalt (E-MORB) (Figure 8). Compared to OIB, their lower Nb/Ta, Nb/La, Ce/Pb, and Nb/Yb ratios also indicate that they unlikely originated from an OIB-like mantle plume but mostly had a MORB-like mantle source (Figures 12(a) and 12(b)). However, enrichment in LREEs and LILEs and depletion in HREEs and HFSEs with negative Nb–Ta anomalies of the samples are indicative of arc-related magmatism and reflect a depleted mantle source previously hybridized by fluids or melts from a subducted slab [68]. On the Th/Yb–Nb/Yb diagram (Figure 12(c)), some of the basaltic rocks plotted above the MORB-OIB array, further suggesting the involvement of an arc-related component [64]. On the Zr/Y–Zr diagram, most basaltic rocks plot in an intraplate setting and a minority plot in an island arc setting, implying the influence of an arc-related component (Figure 12(d)). The depletion in Nb and Ta is attributed to their derivation from a mantle source modified by slab-derived fluids (Figure 12(e)) and sediment-derived melts (Figure 12(f)). These characteristics probably indicate that the magma is derived from a subduction-related fluid/melt-modified mantle source.

The basaltic rocks are characterized by flat chondrite-normalized HREE patterns (Figure 8(a)), indicating that they were derived from a spinel-garnet mantle. This can be further evaluated by REE contents and ratios (e.g., Sm/Yb and Sm; [69]). On the Sm/Yb–Sm diagram, the basaltic rocks from the Qijiaojing section plotted near the (garnet+spinel)-bearing lherzolite melting trend, indicating a spinel-garnet lherzolite mantle source with partial melting of 10%–25%, from a depth of about 60–80 km [70]. On the other hand, the basaltic rocks from the Haxionggou section have lower Sm/Yb ratios and plot below the garnet-spinel lherzolite (1 : 1) melting curve, implying a magma source in a transition zone between spinel lherzolite and garnet-spinel lherzolite (Figure 13). The degree of partial melting is about 15%–35%, from a depth of slightly about 60 km, which contains more spinel. Most basaltic rocks from the Qijiaojiing section have higher Sm/Yb ratios than those from the Haxionggou section, suggesting that they are generated from the partial melting of relatively deep mantle sources with LREE-enriched mantle components (Figures 8(a) and 8(b)).

The Pennsylvanian felsic volcanic rocks in the BMs included basaltic andesites, basaltic trachyandesites, andesites, trachyandesites, dacites, trachydacites, and rhyolites, constituting a successive magmatic evolution sequence (Figure 6(a)). The basaltic and felsic volcanic rocks share similar REE patterns and have identical Sr–Nd isotopic compositions (Figures 10 and 11), implying that they probably share a common mantle-derived parental magma. Chen et al. [46] and Xie et al. [71] have suggested that the felsic volcanic rocks are generated by coeval basaltic magma, but the modelling results using MELTS did not support the fractionation hypothesis [13]. Instead, the Carboniferous zircons from these felsic volcanic rocks have positive $εHft$ values (16HXG-12, 16HXG-13, and 16QJJ-15 have positive $εHft$ values concentrated in the range from +8.0 to +9.8, +13.1 to +15.6, and +10.8 to +13.2, respectively), young two-stage Hf model ages (⁠$majorpeaks<800Ma$⁠), and approximate $TDM1$ and $TDM2$ ages (Supplementary Table 2) suggesting that they are primarily derived from a juvenile continental crust. The published data also show that isotopes of the juvenile continental crust are almost indistinguishable from those of the depleted mantle source (Figure 5(e)). Therefore, the felsic volcanic rocks were probably generated by partial melting of the juvenile arc crust as suggested by Zhang et al. [13].

The Pennsylvanian felsic volcanic rocks in the BMs have high SiO₂ and K₂O content, belonging to the high-K calc-alkaline and shoshonitic series (Figure 6(c)). Most samples also exhibit high Zr, Rb, and Ga/Al, akin to A-type granites (Figure 14(a); [72]). These A-type-like felsic volcanic rocks can be classified into the A2 Group (Figure 14(b)), suggesting that their source has been through a cycle of continent-continent collision or island arc magmatism [73, 74]. On the tectonomagmatic discrimination diagram [75], the felsic volcanic rocks were plotted within an area of overlap among the volcanic arc (VAG), ocean ridge (ORG), and within-plate (WPG) granite fields (Figure 14(c)). Besides, almost all samples are plotted in the postcollisional area on the Rb–Y+Nb diagram (Figure 14(d)).

The mantle plume [36], island arc [34, 35, 71], subduction-torn-type rift [48], and back-arc rift [13] models have been proposed to explain the Carboniferous-Permian tectonic setting of the BMs. In recent years, the Mississippian (ca. 345–330 Ma; [13, 46]), Pennsylvanian (ca. 315 Ma; [76]; this study), and Cisuralian (ca. 295 Ma; [2]) bimodal volcanic rocks were recognized in the Heishankou-Dashitou, Sepikou, and Qijiaojing regions, respectively. Bimodal volcanism commonly occurs in an extensional setting related to intraplate, postcollisional, or back-arc rifting tectonic regimes ([13] and references therein). Xia et al. [36] have proposed that Carboniferous-Permian volcanism is a large igneous province related to an intracontinental rift, whereas more and more researchers demonstrate that the mantle plume is not appropriate due to the different geochemical features, low magma temperature, and long erupted period [13, 22]. Instead, accumulating evidence suggests that the Mississippian volcanic rocks were formed in an arc-related setting (island arc or back-arc; [13] and references therein), and the Permian high-K calc-alkaline and alkaline rocks were related to the postcollisional setting [22]. The main controversy is concentrated on the Pennsylvanian period [34–36] due to their individual characteristics.

Firstly, the Pennsylvanian volcanic rocks of the BMs are composed of various types of high-alumina basaltic rocks, peperites, bimodal volcanic rocks, and A2-type rhyolites, almost all of which have arc-related signatures, such as Nb–Ta negative anomalies (Figure 8). Secondly, except for the basaltic rocks of the Liushugou Formation from the Haxionggou section which have the characteristics of island arc basalts, others plot in the areas of within-plate basalts or E-MORB (Figures 12(b)–12(d)). The felsic volcanic rocks are predominated by high-K calc-alkaline and shoshonitic rocks with characteristics of A2-type granite (Figures 14(a) and 14(b)), and most of them are plotted in complicated tectonic settings, including within-plate, volcanic arc, syncollisional, and ocean ridge settings (Figures 14(c) and 14(d)). Thirdly, petrogenesis analyses revealed that the Pennsylvanian basaltic rocks are mainly derived from a lithospheric mantle source metasomatized by slab-derived fluids and sediment-derived melts, so this naturally explains why these volcanic rocks have arc-like signatures. The zircons acquired from the Pennsylvanian felsic volcanic rocks have lower zircon saturation temperatures (⁠$TZr$ is about 810°C) compared with those from the Mississippian felsic volcanic rocks (⁠$TZr$ is about 960°C) [2, 13, 46, 71]. We agree with the viewpoint that the Mississippian volcanic rocks of the BMs have erupted in a back-arc basin [13, 41, 46]. However, the Pennsylvanian volcanic rocks of the BMs are rather composite and have three unique characteristics which we mentioned above. These three characteristics indicate that they were formed in an intraplate extensional setting and were affected by preceding subduction-related fluids. The collected evidences suggest that it is very challenging to explain these phenomena using the model of plate tectonics. Instead, the postcollisional setting is more suitable for deciphering their geodynamic background. This is consistent with the features of felsic volcanic rocks, which plot in the areas of postcollisional setting (Figures 14(c) and 14(d)). Therefore, we proposed that the Pennsylvanian volcanic rocks of the BMs are formed in a series of postcollisional rift basins (Figure 15), and the eastern BMs have smaller rifting which have a relatively thicker crust and lithospheric mantle (Figure 15).

The postcollisional setting can also be demonstrated by the 316 Ma Sikeshu stitching pluton [39], widespread 315–270 Ma A-type granitoid and mafic intrusions in the continental arc [77], Pennsylvanian-Permian molasse deposits ([78], Zhang et al., 2014), and 310 Ma intracontinental bimodal volcanic rocks [58]. The geodynamic mechanism is probably related to the large-scale rotation and displacement that occurred in the southwestern CAOB during the Pennsylvanian to the earliest Mesozoic [17, 79–81] or during local strike-slip faulting [41].

The Pennsylvanian volcanic rocks of the BMs have a composite composition deposited in a shallow to semideep marine environment with subaqueous volcanism and a progressively deepening process. Of these, the basaltic rocks are derived from the isotopically depleted mantle source (lithosphere mantle) hybridized by subduction-related fluids and sediment-derived melts. The felsic volcanic rocks are derived from the partial melting of the juvenile continental crust. The volcanic and volcano-sedimentary rocks were probably related to a series of postcollisional rift basins, and the eastern BMs were in a relatively shallower basin due to their smaller rifting.

All the raw data are listed in Supplementary Tables 1–4.

The authors declare no conflict of interest regarding this publication.

We are grateful to Zhaojie Guo, Qiugen Li, Yuming Qi, Jian Ma, Yizhe Wang, Jiaxuan Leng, Yue Jiao, Qingyun Li, and Qi Zhao for their assistance in the field and in indoor experiments. This study was financially supported by a National Science and Technology Major Project of China grant (2017ZX05008-001).