Under the petrographic observation of thin sections, avoided carbonate grains (e.g., peloids and bioclasts), only micrite component was sampled (Figure 3). A total of 56 carbonate samples were collected from the Dukou section, of which 40 samples were from the Capitanian succession and 16 samples were from the Wuchiapingian succession. The specific sampling locations can be seen in Figure 1D.

To further minimize the influence of diagenesis on the geochemical signatures, samples were crushed into mm-sized particles to eliminate contamination from the cement veins and recrystallized spots. Samples were ground to 200 mesh in an agate mortar and three subsets were prepared for 1) δ¹³C_carb and δ¹⁸O, 2) ⁸⁷Sr/⁸⁶Sr ratios, and 3) trace element analyses.

During the Capitanian, the δ¹³C_carb presents large fluctuations (Figure 4), which varies from ‒ 3.88‰ to + 5.3‰, with an average of + 3.5‰. In the Early Capitanian, the δ¹³C_carb remains at a relatively stable high value of about + 4‰. In the Late Capitanian, the δ¹³C_carb begins to decrease gradually after reaching a peak value of + 4.78‰, followed by a rapid negative drift which decreases from + 3.2‰ to ‒ 3.9‰. This rapid negative drift is comparable to earlier studies (e.g., Shen et al., 2013, 2020; Yuan, 2015). In the Early Wuchiapingian, the δ¹³C_carb quickly rises from ‒ 3.9‰ to + 2.9‰, and remains at a high value of about + 5‰.

The ⁸⁷Sr/⁸⁶Sr ratios of the Dukou section also show significant fluctuations (Figure 4), which vary from 0.70682 to 0.70733, with an average of 0.70699. In the entire Capitanian, the ⁸⁷Sr/⁸⁶Sr ratios decrease rapidly from 0.70705 to 0.70682, and then remain around 0.70689. Subsequently, the ⁸⁷Sr/⁸⁶Sr ratios rise rapidly to around 0.70713 during the Early Wuchiapingian.

The elemental geochemistry of the investigated carbonates includes the primary productivity (Cu_EF, Ni_EF, Ba_EF) and redox (V/(V+Ni), Th/U and U_EF) (Tribovillard et al., 2006; Shen et al., 2014; Takahashi and AuthorAnonymous, 2014). To normalize the results and describe the enrichment degree of the environmentally sensitive elements, enrichment factors (EF) of selected elements were calculated as: X_EF = (X/Al)_sample/(X/Al)_{average shale}, where X and Al represent the concentrations of the element of interest and the concentrations of average shale are from Wedepohl (1971, 1991). Their profiles are shown in Figure 5 and results are listed in Supplementary Table S1.

According to the systematic variations of trace elements of the Dukou section (Figure 5), we divide it into three stages. The Cu_EF gradually decreases after reaching the maximum in Stage 1, and remains stably low in Stage 2 and the beginning of Stage 3, and subsequently increases. Except that some samples increase in Stage 2, the variation of Ba_EF is similar to that of Cu_EF. Ni_EF fluctuates dramatically in Stage 1 and reaches the maximum at the end of this stage. At the beginning of Stage 2, Ni_EF declines slowly, follows by two subsequent rises, and increases slightly at the end of this Stage. It decreases slowly at the beginning of Stage 3, and gradually increases after stabilization. The value of Ni_EF in Stage 1 is higher than in Stages 2 and 3. In general, the primary productivity proxies seem to exhibit a downward trend in Stage 2. Both V/(V+Ni) and U_EF decrease in Stage 2 and remain at a certain range, which is consistent with the increase of Th/U in Stage 2.

Diagenetic alteration of marine carbonates has been documented to significantly alter the primary geochemical signatures (Veizer, 1983), and the evaluation of carbonate preservation is therefore the foundation needed for the reconstruction of reliable chemostratigraphic profiles and interpretation of the paleoenvironmental conditions. Previous paleoenvironmental studies suggested that the diagenetic alteration at low water/rock interaction ratios, as reflected by minimum recrystallization, does not dramatically alter the signatures of the paleoenvironmental proxies (Veizer, 1983; Azmy et al., 2011). Because the diagenetic fluids are composed of water, the δ¹⁸O is also significantly decreased during diagenesis but the δ¹³C is relatively less susceptible to alteration since the diagenetic fluids do not usually have highly dissolved CO₂ to reset the C-isotope signatures of carbonates unless organic matter is involved (Cochran et al., 2010; Saltzman and Thomas, 2012).

Therefore, the δ¹⁸O values are utilized as a reliable proxy for evaluating the degree of geochemical preservation of carbonates. It is generally believed (Derry et al., 1992; Kaufman and Knoll, 1995) that samples with δ¹⁸O values greater than—eight‰ are commonly accepted as pristine ones. Additionally, the post diagenesis of marine carbonate rocks could cause the simultaneous change of δ¹³C_carb and δ¹⁸O, which led to a positive correlation between δ¹³C_carb and δ¹⁸O (Veizer et al., 1999; Knauth and Kennedy, 2009). However, our data show δ¹⁸O values greater than—eight‰ and a poor correlation coefficient between δ¹³C_carb and δ¹⁸O (R ² = 0.057; Figure 6A). Thus, obvious diagenesis is not supported by the data.

The poor correlation coefficient between Fe contents and ⁸⁷Sr/⁸⁶Sr values (R ² = 0.019; Figure 6B) and the poor correlation coefficient between Fe contents and Mn contents (R ² = 0.421; Figure 6C) suggest a high degree of preservation of near-primary signatures. Additionally, the investigated carbonates have generally low Mn contents (16–355 ppm; Supplementary Table S1) and high Sr contents (58–3880 ppm; Supplementary Table S1), and the Mn/Sr ratios (mainly < 1) exhibit a poor correlation coefficient with their ⁸⁷Sr/⁸⁶Sr values (R ² = 0.018; Figure 6D), also implying insignificant contributions from diagenetic fluids (Li et al., 2021).

As mentioned above, diagenetic alteration of carbonates with progressive burial results in a dramatic depletion in the Sr contents but enrichment in other elements. (Veizer, 1983; Derry et al., 1992; Denison et al., 1994; Li, 2016). Therefore, the Sr ratios are utilized as a dependable proxy for evaluating the degree of geochemical preservation of carbonates. The poor correlations (R ²= 0.099 to 0.001) of Sr with the other proxies (e.g., V/(V+Ni), Th/U, Cu, Ba, Ni, U; Figure 7A–F; Supplementary Table S1) suggest that the studied carbonates succession retain at least near-primary geochemical signatures. (e.g., Shembilu and Azmy, 2022).

The δ¹³C_carb values in the Early Capitanian retain relatively high (∼ +4.5‰) positive plateau. The Cu_EF, Ba_EF, and Ni_EF reveal that there is indeed a relatively high marine primary productivity during the Early Capitanian. Then, the value of δ¹³C_carb shift to a slight negative drifting at around 261.6 Ma. Meanwhile, the primary productivity proxies occurred a slight downward decline. The redox indicators (V/(V+Ni), Th/U and U_EF) recorded in this carbonate platform reflect increased oxidation of surface seawater, which is opposite to the deep water facies that was anoxia and intermittent euxinia evidenced by the coincidence between reduction of framboid size and the extremely negative sulfur isotope values of pyrite (Wei et al., 2016, 2019). To around 260.4 Ma in the Late Captanian, a rapid δ¹³C_carb negative shift occurred, and also had a synchronous response in other regions (Figure 8). Rencunping section in Hunan exhibits two significant negative δ¹³C_carb drifts during the Capitanian (Cao et al., 2018), a negative δ¹³C_carb drift in the J. shannoni zone is related to dolomitization to degree, and that of it in the J. xuanhanensis zone (from + 5‰ to − 2‰) is still primary signal; Xiongjiachang section in Guizhou displays a rapid negative δ¹³C_carb drift (from + 4‰ to ‒ 3‰) in the J. prexuanensis-J.xuanhanensis zone, and neary Gouchang and Houchang sections also show similar negative drift (Bond et al., 2010); Penglaitan and Tieqiao sections of Guangxi presents two significant negative δ¹³C_carb drifts in the J. altudaensi zone and the J. xuanhanensis zone respectively (Tierney, 2010), the latter may correspond to the δ¹³C_carb negative drift in the late Capitanian.

Through investigating volcano-sedimentary record of southwest China, Wignall et al. (2009) revealed that the rapid δ¹³C_carb negative drift coincided with the onset of Emeishan volcanism. According to the ages of conodont biostratigraphy, Sun et al. (2010) proposed that eruption of Eemishan LIP started from the J. altudaensis zone and peaked in the J. xuanhanensis zone. Recently, the new CATIMS U-Pb ages result accurately confirm that the onset time of Eemishan LIP was 260.55 ± 0.07 Ma (Huang et, al., 2022), and the peak time was ∼ 260 Ma (Zhong et al., 2014; Yang et al., 2018). The onset time of the δ¹³C_carb negative drift in this study is consistent with that of the Emeishan volcanic eruption (Figure 8), indicating that Emeishan LIP is undoubtedly one of the important mechanisms driving the δ¹³C_carb negative drift. Meanwhile, the Cu_EF, Ba_EF, and Ni_EF in this interval recorded a decline of marine primary productivity, and V/(V+Ni), Th/U, and U_EF reveal an enhanced oxidation of seawater, which could promote this δ¹³C_carb negative drifting together.

In contrast to the magma eruption scale in Siberia LIP (4×10⁶ km³) and Deccan LIP (2–3 × 10⁶ km³), a total amount of magma in Emeishan LIP is relatively small (0.3–0.6 × 10 ⁶ km³; Shellnutt., 2014), which is might not enough to cause such a large δ¹³C_carb negative drift. In addition to volcanic outgassing, thermal decarbonation reactions also produces a large amount of greenhouse gases. It has been suggested that Emeishan volcanism caused the upwelling of magma to bake the underlying the host sediments (mostly carbonates) to form a large amount of ¹³C-poor CO₂ and CH₄, which entered the atmosphere in a short period (Svensen et al., 2007, 2009; Aarnes et al., 2010; Shellnutt et al., 2012). The widespread volcanic arcs and the initial breakup of Pangea may play an important role in the δ¹³C_carb negative drift during the GLB transition (Maruyama et al., 2007; Macdonald et al., 2019; Zhang et al., 2022b). In addition, in the context of global regression (Ross and Ross, 1987; Haq and Schutter, 2008), the redox indictors (V/(V+Ni), Th/U and U_EF) of this study indicate enhanced marine oxidation in the Late Capitanian. Therefore, the decline of sea-level may cause exposure of continental shelves and re-oxidization of ¹²C-rich organic matter, which will promote such large magnitude of negative drift to some extent.

The δ¹³C_carb values increased rapidly around 259.4 Ma, which can be observed in the Rencunping, Shangsi, Chaotian and Tieqiao sections of South China (Tierney, 2010; Jost et al., 2014; Cao et al., 2018), as well as the sections in Armenia, Turkey, and Europe (Baud et al., 1989; Jost et al., 2014). The cause of this δ¹³C_carb positive drift is considered to be linked to the consumption of atmospheric pCO₂ of low latitude mafic weathering (Emeishan basalt), which led to the onset of the Permian P4 glacial (Yang et al., 2018; Sun et al., 2022).

The strontium isotope variations of this study are consistent with that reported previously (Figure 9; Veizer et al., 1999; Korte et al., 2006; McArthur et al., 2012; Wang et al., 2018). The ⁸⁷Sr/⁸⁶Sr ratios of the Dukou section remained low during the Middle-Late Capitanian, reaching the globally recognized minimum value (0.70682) at approximately 262 Ma (Figure 9). Until approximately 259.4 Ma, the ⁸⁷Sr/⁸⁶Sr ratio also showed a significant recovery (Figure 9).

Prior to the rapid increase of ⁸⁷Sr/⁸⁶Sr ratios, the long-term slow-declining low value indicated the dominant role of mantle-derived Sr Emeishan basalt eruption was considered to be the possible cause of the decrease in the seawater ⁸⁷Sr/⁸⁶Sr ratios (Huang H. et al., 2019; Li et al., 2021). However, the minimum of ⁸⁷Sr/⁸⁶Sr ratio is inconsistent with the duration of the Emeishan LIP (Zhong et al., 2014, 2020; Huang et al., 2018, 2022). It is intriguing that larger Siberian LIP during the end-Permian did not cause such a significant decrease of ⁸⁷Sr^/86Sr ratios (McArthur et al., 2012, 2020; Liu et al., 2013). Therefore, it should be due to other causes that the mantle-derived Sr flux has long been dominant.

From a longer-time scale, the decline of ⁸⁷Sr/⁸⁶Sr ratio during the LPIA has already begun since the Late Carboniferous (McArthur and Howarth, 2004; Kani et al., 2013; Wang et al., 2018). Glacial maximum period (Late Carboniferous-Early Permian) could explain the decrease of crustal-derived Sr to some extent (Chen et al., 2018; Chen et al., 2022). However, it is difficult to provide a reasonable explanation for the continuous decline of ⁸⁷Sr/⁸⁶Sr ratios within the glacial continuous weakening process (Middle Permian).

The dominance of mantle-derived Sr flux could be the important reason for the long-term decline of ⁸⁷Sr/⁸⁶Sr ratios and the minimum value during the Capitanian. This period coincided with the transition from convergence to the disintegration of the Pangea. The Arc volcanism with plate movement (Macdonald et al., 2019), active oceanic ridge magmatism (Zhang B. L. et al., 2021), and their related mafic weathering should provide considerable mantle-derived Sr flux. Marine hydrothermal cherts during the Capitanian were widely distributed in the northern margin of the Upper Yangtze (Murchey and Jones, 1992; Qiu and Wang 2011; Dong et al., 2020), supporting the enhanced supply of mantle-derived material. The long-term low ⁸⁷Sr/⁸⁶Sr value of Dukou section increase rapidly from approximately 259.4 Ma (Figure 9), which corresponds to the attenuation period of weathering intensity. It is difficult to explain this recovery with the change of the crust-derived Sr flux. A significant reduction in mantle-derived Sr flux might mainly be attributed to the rapid weakening of oceanic ridge and arc volcanism associated with active global-scale crust-mantle interaction. In summary, the strontium isotopic composition of this period is more dominated by mantle-derived Sr flux.

The linkage between LIP volcanism and biological crisis events has long been a very interesting topic (Courtillot, 1999; Wignall, 2001; Courtillot and Renne, 2003; Chen and Xu, 2019). Large-scale volcanic activity will release a large amount of volcanic ash, toxic gases, and greenhouse gases in a short period, ultimately leading to environmental degradation and biological crisis (Bond et al., 2010; Day et al. al., 2015; Fan et al., 2020). Based on cases such as the temporal coincidence between the end-Permian mass extinction and the eruption of the Siberian LIP, large-scale volcanic activity is considered to be one of the important factors of the extinction (Burgess and Bowring, 2015; Chen and Xu, 2019).

The GLB transition also experienced a significant extinction event (Jin et al., 1994; Shen and Shi, 1996, 2002; Wignall et al., 2009), which was called " end-Guadalupian mass extinction”. The extent of this extinction event has been controversial (Jin et al., 1994; Stanley and Yang, 1994; Shen and Shi, 1996, 2002; Wignall et al., 2009; Huang Y. G. et al., 2019). However, the decline in the diversity of shallow-marine invertebrates (including corals, brachiopods and ammonoids) still reflects that this extinction event is indisputable (Wang and Sugiyama, 2000; Groves and Wang, 2013; Shen et al., 2020). The trace elements of the samples from the Dukou section show that the enrichment factors such as Cu_EF, Ba_EF, and Ni_EF decreased accompanied by a δ¹³C_carb negative drift in the Late Capitanian (Figure 7), indicating a decrease in marine primary productivity. This decline is in good chronological agreement with the previously proposed GLB biocrisis.

Wignall et al. (2009) proposed that the Emeishan volcanic activity might trigger a biological extinction event based on the temporal coincidence among Emeishan volcanism, biocrisis and carbon isotope negative drift. Bond et al. (2010) reported the loss of keriothecal-walled fusulinaceans and a turnover in calcareous algae in the Upper Maokou Formation at the Xiongjiachang section, and carbonate deposition was frequently interrupted by thick volcanic ash depositional events, revealing a potential link between Emeishan volcanism and biological extinction. The sedimentary, paleoecologic, and geochemical analysis of the Penglaitan section show collapse of metazoan reef system, with two abnormally high mercury concentration/total organic carbon (Hg/TOC) ratios during the GLB transition, supporting that the Emeishan LIP may trigger biotic extinction (Figures 10–f; Huang Y. G. et al., 2019). Accumulating evidence shows that the Emeishan LIP may be one of the important mechanisms of the end-Guadalupian biocrisis (Wignall et al., 2009; Bond et al., 2010; Chen et al., 2011; Chen and Xu, 2019; Sun et al., 2022). Here, we collected a series of paleoenvironmental indicators to integrate our results and carried out a global-scale comprehensive comparative analysis (Figure 10).

A significant δ¹³C_carb negative shift during the GLB transition is consistent with the onset of the Emeishan eruption and the initial breakup of the Pangea. This temporal consistency implies that it is related to the carbon-cycle disturbance caused by the global volcanic-tectonic activity, and is consistent with the GLB extinction. The ⁸⁷Sr/⁸⁶Sr ratios discussed above is dominated by the mantle-derived Sr flux, which also indicates that the global volcanic-tectonic activity caused by Earth's internal dynamics is the triggering factor for this biocrisis event. Declines in biodiversity are often a precursor to extinction events (Stanley, 2016). Biodiversity decreased significantly during the Emeishan LIP (Chen et al., 2011; Davydov, 2014; Zhang B. L. et al., 2021). In particular, the latest high-resolution paleontological data strongly confirms that species diversity declined rapidly during the Late Capitanian (Fan et al., 2020).

The reconstructed atmospheric CO₂ curve reveals there was a transient climate warming event in Late Capitanian (Figures 10–h; Foster et al., 2017). Meanwhile, the conodont δ¹⁸O_apatite results of the Penglaitan section show that the seawater temperature increased about 6 ∼ 8 °C (Figures 10–g; Chen et al., 2011), and the brachiopod δ¹⁸O_calcite results of the Xikou section show that the seawater temperature increased about 4°C (Figures 10–g; Wang et al., 2020), both revealing a significantly transient warming event. Recently, the CIA curve established by the acid-insoluble residues of carbonate rocks (Figures 10–e; Sun et al., 2022), compared with the latest high-precision age of the Emeishan LIP (Huang et al., 2022), supporting a high-temperature plain in the main peak period (approximately 260 Ma), which may represent a transient warming event (Figure 10). This climate event may have been an important cause of the end-Guadalupian crisis. The redox proxies from this study support there was an enhanced seawater oxidation during the Late Capitanian (Figures 10–d), the onset of that is consistent with the end of the P3 glacial. Thereafter it in turn transitions to a reductive increase from the onset of P4 glacial. Obviously, the transient climate warming may be an important reason for the enhanced oxidation of the shallow-water shelf during this event, and inhibit the preservation of organic matter, which is precisely coupled with the carbon isotope negative drift. In contrast to the enhanced oxidation of the shallow-water shelf represented by the Dukou section, due to the transient climate warming, the ocean circulation may have weakened associated with a decrease in the temperature difference between high and low latitudes. Deep-water facies beyond the shelf may exacerbate oceanic water column stratification and marine anoxia, and even euxinia are not conducive to the survival and reproduction of organisms (Yan et al., 2013; Wang et al., 2022).

It was discussed previously that the onset of δ¹³C_carb negative drifts and the strontium minimum was inconsistent during the Capitanian (Figure 10). However, across the GLB boundary, both showed a rapid positive drift almost simultaneously, and paced climate transition from interglacial to P4 glacial (Frank et al., 2015; Metcalfe et al., 2015; Davydov et al., 2022). But the eruption of the Emeishan LIP was not completely over, and the climate transitioned into the P4 glacial quickly. This may be due to the weathering of low-latitude mafic rocks and the possible volcanic ash aerosol effect (Huang et al., 2018; Yang et al., 2018), which led to the climate cooling. Consequently, the enrichment factors for trace elements in our study showed that primary productivity went into a tendency of recover in the Early Wuchiapingian, which proves that the biocrisis had been temporarily alleviated (Figures 10–c). (Azmy and Lavoie, 2009, Azomani et al., 2013, Kaufman et al., 1992, Shen and Mei, 2010)