Late Guadalupian–early Lopingian marine geochemical records from the Upper Yangtze, South China: Implications for climate-biocrisis events

1 Introduction

The Guadalupian-Lopingian boundary (GLB) transition was a critical time interval during the Earth’s evolution history, associated with global major changes of paleoclimate, paleo-ocean and paleogeography. Many special geological events occurred during this period, such as the onset of long-term ocean hypoxia (Isozaki, 1997; Wei et al., 2016; Zhang B. L. et al., 2021), a lowest sea-level in the Phanerozoic (Ross and Ross, 1987; Haq and Schutter, 2008; Chen et al., 2009), Emeishan Large Igneous Province (LIP) (Xu et al., 2010; Shellnutt, 2014; Zhong et al., 2014), initial breakup of Pangea (Maruyama et al., 2007; Isozaki, 2009; Santosh et al., 2009), Kamura event (Isozaki et al., 2007, 2011) and end-Guadalupian biocrisis (Stanley and Yang, 1994; Shen and Shi, 1996, 2002; Chen et al., 2009; Huang Y. G. et al., 2019), etc.

Accompanied by the violent eruption of the Emeishan LIP during the Late Guadalupian, which causing climatic fluctuations and ocean evolution (Chen et al., 2011; Zhang B. L. et al., 2021; Sun et al., 2022). Emeishan volcanism is also considered to be a potential kill mechanism of the end-Guadalupian biodiversity crisis (Wignall et al., 2009; Bond et al., 2010; Chen and Xu, 2019; Huang et al., 2022). Simultaneously, causal links between volcanic activity, biological crisis, climate change, and other geological events are still under discussion (Wignall et al., 2009; Bond et al., 2010; Cao et al., 2018; Chen and Xu, 2019; Zhang B. L. et al., 2021; Sun et al., 2022). Marine carbonate rocks usually have the original geochemical composition that can approximately represent the paleo-seawater, and record the paleoenvironmental information such as seawater temperature (Zhao et al., 2022), redox conditions (Wignall and Twitchett, 1996; Rimmer, 2004; Tribovillard et al., 2006; Scholz et al., 2011; Takahashi and AuthorAnonymous, 2014; Shembilu and Azmy, 2022) and primary productivity (Tribovillard et al., 2006; Schoepfer et al., 2015). Nevertheless, due to global extensive regression, most regions have transformed from marine strata to non-marine strata (Ross and Ross, 1987; Haq and Schutter, 2008; Shen et al., 2020). There are relatively continuous marine carbonate rocks deposited in the South China at low-latitude region, which provide an excellent material for the paleoenvironment research. Moreover, the abundant marine fossils can constrain the sedimentary time of the strata.

Geochemical indicators in carbonate rocks can usually be used to indicate paleoenvironmental change (Shen et al., 2014; Takahashi and AuthorAnonymous, 2014; Li et al., 2017; Chen et al., 2021). The carbon isotope excursion is closely related to the global carbon-cycle (Kump and Arthur, 1999). Positive drift of δ¹³C is usually linked to the burial of organic carbon, while negative drift of δ¹³C indicates the rapid injection of light carbon materials (Saltzman et al., 2000; Li D. D. et al., 2018). Carbon isotope has rapid responses to environmental changes caused by volcanic eruptions (Arthur et al., 1985; Wignall et al., 2009; Isson et al., 2019), and tend to be globally isochronous. For example, associated with the eruption of the Siberian LIP, the δ¹³C showed a rapid negative drift near the PTB boundary (Huang, 1994; Korte and Kozur, 2010; Shen et al., 2013). The strontium isotopic composition of seawater is mainly driven by two different fluxes (Li and Elderfield, 2013; Peucker-Ehrenbrink and Fiske, 2019; McArthur et al., 2020): 1) the unradiogenic hydrothermal mantle-derived strontium flux. 2) The more radiogenic crustal-derived strontium. Marine carbonates can record the relative variation of mantle-derived Sr flux (provided by volcanic activity) and crust-derived Sr flux (provided by continental weathering) (Wang et al., 2018, 2021; Li et al., 2021).

Here, we present several geochemical proxy signatures (carbon and strontium isotopes, trace elements of carbonate rocks) of Dukou section, northern margin of the Upper Yangtze. To investigate the GLB marine succession, comparing with other reported paleoenvironmental indicators and contemporaneous geological events. This paper discusses the driving mechanism of climatic fluctuation and biological crisis during the GLB transition.

2 Geological setting

During the Guadalupian, South China was situated in the equator at low-latitude region, and surrounded by the Paleotethys (Scotese, 2016) (Figure 1). Marine carbonate rocks rich in biological fossils were deposited in most regions of South China, which are called Maokou Formation in the Upper Yangtze region. Affected by the Emeishan LIP and the Dongwu movement in the Late Guadalupian, a local deep-water trough was formed in the northern margin of the Upper Yangtze (Yang et al., 2021), and shelf facies were developed (Jin et al., 1998; Shen et al., 2013; Hou et al., 2020) (Figure 1B).

FIGURE 1

FIGURE 1. (A) Global paleogeographic map during the GLB transition (Scotese. 2016). (B) Paleogeographic map of South China in the GLB transition, modified from Yao et al. (2015) (1. The Rencunping section 2. The Xiong Jiachang section 3. The Gouchang section 4. The Tieqiao section 5. The Penglaitan section) (C) Geological map of the study area (J₃ = Upper Jurassic, J₂ = Middle Jurassic, J₁ = Lower Jurassic, T₃ = Upper Triassic, T₂ = Middle Triassic, T₁ = Lower Triassic, P₃ = Upper Permian, P₂ = Middle Permian, S = Silurian). (D) Lithologic stratigraphy of the upper Maokou Formation and the lower Wuchiaping Formation at the Dukou section. The conodont biostratigraphy of the investigated section is based on Mei et al. (1994a, 1994b), Yuan (2015) and Shen et al. (2020). Abbreviation: J. postserr.=Jinogondolella postserrata zone, J. shan.=Jinogondolella shannoni zone, J. alt.= Jinogondolella altudaensis zone, J. prexuan.=Jinogondolella prexuanhanensis zone, J. xuanhanensis.=Jinogondolella prexuanhanensis zone, C. dukou.=Clarkina dukouensis zone, C. asymm.= Clarkina asymmetrica zone.

This investigated Dukou section (31°42′11″N, 108°18′35″E) is located in the Xuanhan City on the northern margin of the Upper Yangtze, structurally belongs to the intersection zone between the Daba Mountain NW-trending structural belt and the Huaying Mountain NE-trending structural belt. The Dukou section is one of the relatively complete exposed Permian successions of South China (Jin et al., 1998; Shen et al., 2013; Hou et al., 2020), and in ascending order, the Liangshan, Chihsia, Maokou, Wuchiaping, and Changxing formations are exposed along both sides of the Lishi river (Figure 1C).

The upper Maokou Formation (Capitanian sequence, 0–77 m) and lower Wuchiaping Formation (Early Wuchiapingian sequence, 77–106 m) consist of this study carbonate succession at the Dukou section. The lithology of the studied succession is dominated by light-gray micrite limestone, and the lower part of the Capitanian and Wuchiaping sequences are dark-grey limestone with cherty nodules (Figure 1D). It should be pointed out that that 9 m thick black chert with limy nodules appeared at the top of the Capitanian sequence, and this silica unit is also called the “Kuhfeng Chert Unit” (Qiu and Gu, 1991; Mei et al., 1994a; Hu, 2000). The lithological GLB can be placed at the base of the yellow-green clay rock (Wangpo Shale), the Wangpo Shale is also recognized across South China as the lithological boundary of GLB (Shen et al., 2013; Cao et al., 2018).

The global Late Guadalupian regression and eruption of Emeishan LIP in South China, and resulted in a widespread GLB sedimentary hiatus across most of South China (Mei, 1995; He et al., 2003; Haq and Schutter, 2008; Hou et al., 2020). In the southwestern part of South China, the Emeishan LIP, beginning in the late Guadalupian (Zhou et al., 2002; Shellnutt, 2014; Zhong et al., 2014), had a great influence on the Lopingian depositional system to the east. The terrestrial Xuanwei Formation developed landward of the intermediate zone, the transitional Longtan Formation developed the seaward of the intermediate zone, and the marine Wujiaping Formations mainly developed in the outer zone (Figure 2).

FIGURE 2

FIGURE 2. GLB lithostratigraphic sedimentary sequence (Wang et al., 2020), spatial-temporal distribution of the Emeishan LIP (Shen et al., 2011; Zhong et al., 2014, 2020; Li D. D. et al., 2018; Huang et al., 2018, 2022; Shellnutt et al., 2020), and GLB biostratigraphic correlation (Mei et al., 1994a; 1994b; Mei, 1995; Lai et al., 2008). Chronostratigraphic scale is according to Shen et al. (2022).

GLB marine strata succession in South China consists of the Maokou and Wuchiaping formations (Sun et al., 2010; Shen et al., 2019). Conodont biostratigraphy of South China indicates that the upper Maokou Formation and Wuchiaping Formation correspond to the Capitanian stage (264.28–259.51 Ma) and the Wuchiapingian stage (259.51–254.14 Ma), respectively (Shen et al., 2020, 2022). In the Upper Yangtze, compared with the Shangsi and Chaotian sections, the Dukou section retains four conodont zones in the Capitanian stage (Mei et al., 1994a; 1994b; Yuan, 2015; Shen et al., 2020) (Figure 1D, Figure 2), indicating that the sedimentary record is relatively complete.

Conodont biostratigraphic correlations from the marginal inner zone to the outer zone of the Emeishan LIP, shows that initial eruption began in the J. altudaensis zone of the middle Capitanian (261–262 Ma), and dramatically increased in extent and volume in the J. xuanhanensis zone (260–261 Ma) (Sun et al., 2010; Shen et al., 2019). Increasingly reported high-precision U-Pb ages (Shen et al., 2011; Zhong et al., 2014, 2020; Li Y. et al., 2018; Huang et al., 2018; Shellnutt et al., 2020; Figure 2) constrained the Emeishan LIP to Late Capitanian-Early Wuchiapingian (approximately 261–257 Ma). The new CA-ID-TIMS U-Pb ages from the Binchuan area of inner zone demonstrate that Emeishan volcanism began by 260.55 ± 0.07 Ma and persisted until at least 257.22 ± 0.37 Ma (Huang et al., 2022; Figure 2).

3 Samples and methods

3.1 Sample preparations

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.

FIGURE 3

FIGURE 3. Photomicrographs of micritic to near-micritic carbonates. (A–D): Samples DK-14, DK-32, DK-51, and DK-60).

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.

3.2 Methods

3.2.1 Carbon and oxygen isotopes

Carbon and oxygen isotopes (56 samples) were measured at the Nanjing Hongchuang Exploration Technology Service Co., Ltd, Nanjing, China. Isotope ratios are reported in per mil (‰) notation relative to the Vienna Pee Dee Belemnite standard (VPDB ‰). Samples were calibrated using China standards (GBW04405). The uncertainty (1σ) calculated from standards per run is typically 0.1‰.

3.2.2 ⁸⁷Sr/⁸⁶Sr ratios

A total of 56 samples were analyzed for ⁸⁷Sr/⁸⁶Sr ratios at the State Key Laboratory of Oil and Gas Reservoir Geology and Exploitation of the Chengdu University of Technology, China. The samples (50 mg each) were washed with deionized water, dried, and crushed to less than 0.0750 mm. The ⁸⁷Sr/⁸⁶Sr ratios were measured using a Thermo-Fisher Scientific Triton Plus mass spectrometer.

The measured ⁸⁷Sr/⁸⁶Sr ratios were normalized to an ⁸⁶Sr/⁸⁸Sr ratio of 0.1194 (Nier, 1938), calculated from 150 measurements. An internal precision of ∼5×10^–6 (relative SE) was maintained. The analytical precision was monitored using NBS 987 and the mean measured value obtained for NBS 987 was 0.710,247 ± 0.000008 [2σ, number of samples (n) = 6]. All reported ⁸⁷Sr/⁸⁶Sr data were corrected to the value of the international standard NBS 987 (⁸⁷Sr/⁸⁶Sr = 0.710,248).

3.2.3 Trace elements

Major and trace elements (56 samples) were measured in Chengdu Oriental Mineral Development Technology Research Institute. Major elements were measured by ZSX Primus II X-ray fluorescence spectrometer. Trace element analysis was performed on an American PE 5300 V. The analysis process adopts China standards GBW07314, GBW07315, GBW07316 and USGS basalt standard material 6BHVO-2 for quality control, and uncertainty was generally better than 5%.

4 Results

4.1 Carbon and strontium isotopes

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‰.

FIGURE 4

FIGURE 4. The profiles of carbon, oxygen and strontium isotopes of the Dukou section (The legend is shown in Figure 1D, the triangles are samples that may be affected by diagenesis) (The data of δ¹³C_carb discussed here was published with no detail discussion in our previous paper, the lower 36 data points of Sr isotopes was published in our previous paper).

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.

4.2 Trace elements

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.

FIGURE 5

FIGURE 5. Profiles of (A) Cu_EF, (B) Ba_EF, (C) Ni_EF, (D) V/(V+Ni), (E) Th/U and (F) U_EF of Dukou section, the triangles are samples that may be affected by diagenesis.

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.

5 Discussions

5.1 Evaluation of sample preservation

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 (R² = 0.057; Figure 6A). Thus, obvious diagenesis is not supported by the data.

FIGURE 6

FIGURE 6. Scatter diagrams showing correlations of (A) δ¹³C_carbwith δ¹⁸O, (B) Fe with ⁸⁷Sr/⁸⁶Sr, (C) Fe with Mn, and (D) Mn/Sr with ⁸⁷Sr/⁸⁶Sr in the investigated Dukou section (Red rings indicate data may have been affected by diagenesis).

The poor correlation coefficient between Fe contents and ⁸⁷Sr/⁸⁶Sr values (R² = 0.019; Figure 6B) and the poor correlation coefficient between Fe contents and Mn contents (R² = 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 (R² = 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 (R²= 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).

FIGURE 7

FIGURE 7. Scatter diagrams showing correlations of Sr with (A) V/(V+Ni), (B) Th/U, (C) Cu, (D) Ba, (E) Ni, and (F) U in the investigated Dukou section.

5.2 The negative drifting of δ¹³C_carb during the GLB transition

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.

FIGURE 8

FIGURE 8. Carbon isotope chemostratigraphy of the South China during the GLB transition. The reported high-precision zircon U-Pb age data of the Emeishan LIP are from Zhong et al. (2014); Yang et al. (2018); Huang et al. (2022). The reported carbon isotope data: Penglaitan section is from Chen et al. (2011); Tieqiao section is from Tierney. (2010); Xiongjiachang and Gouchang sections are from Bond et al. (2010); Rencunping section is from Cao et al. (2018). The age of Capitanian is from JTS-2022. 1 = Clarkina postbitteri hongshuiensis, 2 = Clarkina postbitteri postbitteri.

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).

5.3 Dominant cause of ⁸⁷Sr/⁸⁶Sr ratios’ variation during the GLB transition

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).

FIGURE 9

FIGURE 9. The profile of ⁸⁷Sr/⁸⁶Sr during the GLB transition. Note: The data of the Dukou section are from this study and other data are from Veizer et al. (1999), Korte et al. (2006), McArthur et al. (2012) and Wang et al. (2018).

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.

5.4 The integrated geochemical indices indicating the co-evolution of climate-ocean-biodiversity system during the GLB transition

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).

FIGURE 10

FIGURE 10. Paleoenvironmental indicators and geological events during the GLB transition. The profiles of δ¹³C_carb,⁸⁷Sr/⁸⁶Sr, Cu_EF, and U_EF are from the study of the Dukou section, and the CIA values are from Sun et al. (2022), Hg/TOC ratios are from Huang Y. G. et al. (2019), relative seawater temperature is based on δ¹⁸O composition in conodonts and brachiopods from Chen et al. (2011) and Wang et al. (2020). The atmospheric CO₂ concentration is from Foster et al. (2017), warm-water benthic foraminifer diversity from Davydov, (2014), global sea-level change from Haq and Schutter (2008). Geological event: Emeishan LIP (Shellnutt et al., 2012; Zhong et al., 2014; Huang et al., 2018), the GLB extinction (Wignall et al., 2009; Huang Y. G. et al., 2019; Shen et al., 2020), the Kamura events (Isozaki et al., 2007, 2011), dispersal and assembly of Pangea (Maruyama et al., 2007; Isozaki, 2009; Santosh et al., 2009; Shen et al., 2019). The duration of P3-P4 glacial (Metcalfe et al., 2015), sea-level change (Haq and Schutter, 2008). Abbreviations: P3 = the third stage of Permian glaciation, P4 = the fourth stage of Permian glaciation, 1 = Clarkinapostbitteri hongshuiensis, 2 = Clarkina postbitteri postbitteri, geological age and conodont biostratigraphy is based on Shen et al. (2019) and Shen et al. (2022).

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)

6 Conclusion

1) The δ¹³C_carb curve of the Dukou section has a significant shift during the GLB transition, and a negative drift in the Late Capitanian recorded the carbon-cycle perturbation event promoted by primary productivity proxies and redox condition initially. The surge of Emeishan LIP, contemporaneous volcanic arcs and local re-oxidized organic matter under the regression setting may be an superimposed factor for the later rapid negative drift.

2) The ⁸⁷Sr/⁸⁶Sr ratios of the Dukou section remained long-term low values during the entire Capitanian, reached a global minimum value around 262 Ma and appeared a significant increase till approximately 259.4 Ma. Fluctuations of strontium isotopic composition are more dominated by mantle-derived Sr flux.

3) The composited geological records of the climate-ocean system during the GLB transition, promoted carbon cycling perturbation and triggered the end-Guadalupian biocrisis.

Data availability statement

The original contributions presented in the study are included in the article/Supplementary Material, further inquiries can be directed to the corresponding authors.

Author contributions

RL, Writing–original draft, investigation, data curation. SS, Conceptualization, writing–review and editing, Investigation. WX, Data curation, investigation. AC, Conceptualization, writing–review and editing, Supervision, funding acquisition. JO, Reviewing. SY, Investigation. SX, Supervision. ZL, Reviewing. DY, Data curation. MH, Conceptualization.

Funding

This study was supported by the National Natural Science Foundation of China (Grant No. 42272132, No. 42172119 and No. 42050104), the Everest Scientific Research Program of Chengdu University of Technology (Grant No. 2022ZF11402).

Acknowledgments

We thank reviewers and editors for their constructive reviews, and thank Wen Li, Haotian Luo, Ouwen Yang, Fuxiang Li, Yifan Huang for their helpful field works.

Conflict of interest

The authors declare that the research was conducted in the absence of any commercial or financial relationships that could be construed as a potential conflict of interest.

Publisher’s note

All claims expressed in this article are solely those of the authors and do not necessarily represent those of their affiliated organizations, or those of the publisher, the editors and the reviewers. Any product that may be evaluated in this article, or claim that may be made by its manufacturer, is not guaranteed or endorsed by the publisher.

Supplementary material

The Supplementary Material for this article can be found online at: https://www.frontiersin.org/articles/10.3389/feart.2022.1077017/full#supplementary-material

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