Guadalupian (Middle Permian) δ13Corg changes in the Lower Yangtze, South China

Hengye Wei· Ziao Geng · Xuan Zhang

Abstract The Middle Permian Guadalupian witnessed significant environmental changes in the Phanerozoic such as large-scale sea-level drop, supercontinent Pangaea assembly, and transition from Early Permian glaciation to Late Permian non-glacial intervals. Carbonisotope tracers can provide insights for these environmental changes. The δ13C studies of the entire Guadalupian epoch are rare, and most of them has focused on near the end of Guadalupian and carbon isotopes of inoragnics.Here,we present carbon isotopic compositions of organic matters in the Guadalupian from two sections(Chaohu and Xiaolao)in the Lower Yangtze area, South China. Our results show that δ13Corg profiles in the Guadalupian show a peak in the Roadian and a gradual negative shift from the Roadian to the middle Capitanian. These trends can be matched by δ13C changes of carbonate rocks or organic matter in South China and other places in the world, representing a global carbon cycle signal. The Roadian positive peak was probably due to high productivity which was caused by upwelling during cooling time. The gradual negative shift of δ13C was caused mainly by a decrease of organic matter burial on land and in the ocean, resulting from global sea-level drop and anoxia-caused benthos decline, respectively. The less important causes for the gradual δ13C negative shift are volcanic-gases releasing,decreased mountain belts,and the resultant reduced silicate weathering-consumption of CO2.The gradual negative shift of δ13C coincides with the gradual extinction in the Guadalupian. Therefore, global sea-level drop and marine reducing conditions may be the main causes of the gradual extinction in the Guadalupian.

Keywords Organic-carbon isotope · Carbon cycle · Mass extinction · Kuhfeng Formation · Chaohu · Xiaolao

1 Introduction

The Middle Permian Guadalupian epoch was a transition time from Early Permian glacial interval to Late Permian non-glacial interval(Veevers and Powell 1987).Associated with this environmental change, a significant mass extinction occurred from the mid-Capitanian to the Guadalupian–Lopingian (G–L) boundary (Shen and Shi 2009; Bond et al. 2010a). This mass extinction is called Guadalupian mass extinction (Clapham 2015), and also called end-Guadalupian (Jin et al. 1994; Shen and Shi 1996,2002)or mid-Capitanian mass extinction(Bond et al.2010a, b). Before this mass extinction, biodiversity increased in the Early Guadalupian (Roadian) and then declined gradually from the Middle Guadalupian (Wordian)to the middle Capitanian(Late Guadalupian)(Groves and Wang 2013).Except for the anoxia(Zhang et al.2015;Wei et al. 2016), large-scale volcanism such as the Emeishan large igneous province(LIP)had been suggested to be the main potential cause for this mass extinction(Wignall et al. 2009; Huang et al. 2018; Chen and Xu 2019). The volcanic greenhouse gases from the Emeishan LIP disturbed the global carbon cycle and caused a large carbon-isotope negative excursion from the mid-Capitanian to the G–L boundary(Wang et al.2004;Bond et al.2010b;Shen et al.2013),though the global nature of this negative excursion had been questioned (Jost et al. 2014). Did this carbon-isotope negative excursion take place from a longterm carbon cycle stability or reflect carbon cycle perturbations which started from the Early and Middle Guadalupian? Addressing this question will assist in the understanding of the causes for this mass extinction,as the Early-Middle Guadalupian marine environmental conditions before the mass extinction event seems to be essential(Musashi et al. 2010). The ocean processes that interact with the global carbon cycle affect the chemical- and physical-conditions of bottom water where benthic faunas lived during the Guadalupian crisis.

The transfer of organic carbon in the ocean is forced byabiological pump regulated by climate changes (Goodwin et al.2008).During this process,organic carbon hiding in the sediments keeps oxygen in the atmosphere (Hemingway et al. 2019). Therefore, the stable isotopic composition of organic matter is often used for the reconstruction of paleoclimates and paleoceanography in the past (Arthur et al. 1988; Popp et al. 1997; Hayes et al.1999;Dawson et al.2002;Galimov 2006).The covariance of carbon-isotope composition from carbonate rocks and organic matter can be used to support for global carbon changes (Dean et al. 1986; Freeman and Hayes 1992;Magaritz et al. 1992; Holser 1997; Kump and Arthur 1999).Unfortunately,most of the carbon isotope studies in the Guadalupian had been inorganic-carbon isotope analyses (Wang et al. 2004; Isozaki et al. 2007; Bond et al.2010b;Shen et al.2013;Liu et al.2017;Cao et al.2018).A few studies of organic-carbon isotopic compositions had focused on the G–L boundary or the Capitanian (Kaiho et al. 2005; Bond et al. 2015), the Early Guadalupian (Liu et al. 2018), or the Early-Middle Guadalupian with lowresolution samples (Birgenheier et al. 2010; Zhang et al.2020).

We present here organic-carbon isotope stratigraphy in the Guadalupian from two separated sections (Chaohu and Xiaolao)in the Lower Yangtze area,South China.We also analyzed TOC and H contents to constrain the organiccarbon isotopic variations. Our results will help to analyze the global carbon cycle during the Guadalupian, and to constrain the environmental changes during the Guadalupian biocrisis for a better understanding of mass extinction causes.

2 Geological setting

During the Guadalupian,the South China block was a large tropical isolated platform located in the east of the Paleo-Tethys Ocean (Fig. 1a). The South China block is composed of the carbonate Yangtze Platform in the west and the arenaceous Cathaysia Platform in the east. The Lower Yangtze area was located in the northeast part of the South China block (Fig. 1b). This area was framed by the Jiangshao fault in the east and the Tanlu fault in the west(He et al. 2014). During the Middle Permian, the depression of this area formed the Lower Yangtze Basin, which deposited deep-water facies of the black cherts in the Kuhfeng Formation (Li et al. 2008).

The two studied sections at Chaohu (25° 01′44′′N,113° 44′18′′E) and Xiaolao (30° 56′47′′N,118° 0′52′′E) were located in the deep basin and the platform slope environments, respectively (Fig. 1b). At the Chaohu section, the Cisuralian Chihsia Formation, the Guadalupian Kuhfeng, and the Yinping formations outcrop in ascending order(Fig. 2),where the Kuhfeng Formation is the main studied interval. The upper Chihsia Formation consists of gray bioclastic limestones with abundant crinoids, foraminifers, fusulinids, ostracods, brachiopods, and some gastropods and bivalves. The contact between the Chihsia and the Kuhfeng is separated by an unconformity,suggesting a hiatus. At the base of the Kuhfeng Formation,brown mudstones intercalated with several layers of bentonites are of the age 272.95 ± 0.11 Ma and 271.038 ± 0.097 Ma (Wu et al. 2017), suggesting the base of the Guadalupian epoch (Shen et al. 2019). The lower Kuhfeng Formation consists of black muddy cherts with abundant phosphorite nodules. Kametaka et al. (2005)inferred that these phosphorite nodules formed in dysoxic environments with low terrestrial detritus supply. The middle and upper Kuhfeng Formation consists of thin-bedded black cherts with high organic matter content (Fig. 2). The Kuhfeng Formation contains abundant radiolariansand sponge spicules.Radiolarian assemblage studies suggest that the Kuhfeng Formation spans the Roadian to the middle Capitanian stages(He et al.1999;Kametaka et al.2009).At the top of the Kuhfeng Formation, several cm-scale bentonite layers have an age of 261.6 ± 1.5 Ma (Zhang et al.2019). The contact between the Kuhfeng and Yinping formations is conformable.The Yinping Formation consists of dark-gray shales which turn into mixed colors of purple,dark-gray and white altered by modern weathering.

At the Xiaolao section, the Cisuralian Chihsia Formation and the Guadalupian Kuhfeng Formation outcrop in ascending order (Fig. 2). The upper Chihsia Formation consists of medium- to thick-bedded grey limestones with abundant foraminifers, bryozoans, ostracods, and green algae. The contact between the Chihsia and Kuhfeng formations is conformable. The lower Kuhfeng Formation consists of thin- to medium-bedded cherty limestones and black shales with phosphorite nodules which can be correlated with the phosphorite nodules interval in the lower Kuhfeng Formation at Chaohu section, ～85 km away.The middle Kuhfeng Formation consists of gray to darkgray thin-bedded cherts with rare radiolarians and common sponge spicules.The upper Kuhfeng Formation consists of gray medium-bedded nodular cherts containing sponge spicules and ostracods. According to the stratigraphic characteristics of the Lower Yangtze area (Li et al. 2008),the upper Kuhfeng Formation is the transition interval between the Kuhfeng and the Wuxue formations,the latter is upper Capitanian limestone deposition. Therefore, the Kuhfeng Formation at Xiaolao spans from the Roadian to the middle Capitanian.

Fig. 1 a Guadalupian global paleogeography. SC South China. NC North China. IC Indochina. T Tarim. CG Central Guizhou. Base map courtesy of Ron Blakey(http://jan.ucc.nau.edu/～rcb7/).Note:70°counterclockwise of South China block relative to its modern orientation(Nie 1991). b Guadalupian paleogeography of South China (modified from Wei et al. (2016) and Chen et al. (2018)). CH Chaohu, XL Xiaolao. The base map with coordinate was adapted from Google maps

3 Methods

We collected samples from the upper Chihsia to the whole Kuhfeng formations.The samples were cleaned,dried,and then ground in an agate ball mill. Powdered samples were digested by 2 N HCl to remove carbonate and then treated with HF to remove chert at 50 °C for more than 48 h. The residues were washed using 18.2 MΩ water until pH was neutral, and then centrifuged and dried for measurements of total organic carbon (TOC) and H contents, and bulk organic-carbon isotope ratios. The residues sample splits were used to analyze TOC and H contents by a Vario ELIII elemental analyzer at the Key Laboratory of Tectonics and Petroleum Resources of Ministry of Education, China University of Geosciences, Wuhan. Absolute precision for TOC and H is better than 0.3 %. For organic-carbon isotopic composition measurement, the residues sample splits were combusted at 800 °C and the released CO2was used to determine13C/12C ratios using a Thermo Scientific MAT 253 at the Nanjing Institute of Geology and Palaeontology,Chinese Academy of Sciences.All the data are reported in per mille (‰) relative to the Vienna Pee Dee Belemnite(V-PDB)standard.The precision was better than ± 0.2 ‰based on the duplicate analyses of samples and standards of charcoal (GBW04407, - 22.43 ‰) and urea(IVA3380217, - 40.81 ‰).

Fig. 2 Lithological logs and field photos at the Chaohu and Xiaolao sections, South China

4 Results

At the Chaohu section,TOC contents range from 0.03%to 0.11 %in the Chihsia Formation and range from 0.48%to 21.48 % (average = 4.52 %) in the Kuhfeng Formation(Table 1). The TOC profile shows two peaks in the lower and upper Kuhfeng Formation, respectively (Fig. 3). H contents range from 0.01 % to 0.05 % in the Chihsia Formation, and range from 0.15 % to 1.30 % (average = 0.37 %) in the Kuhfeng Formation. Atomic H/C ratios range from 1.21 to 5.66 in the Chihsia Formation and range from 0.26 to 5.56 (average = 1.87) in the Kuhfeng Formation. Bulk organic-carbon isotopic compositions(δ13Corg) range from - 28.7 to - 24.8 ‰ (average = - 26.9 ‰). The δ13Corgprofile shows a peak in the lower Kuhfeng Formation and then a gradually decreasing trend upward (Fig. 3).

At Xiaolao section,TOC contents range from 0.15%to 0.33 %in the Chihsia Formation,and range from 0.03%to 0.66 % (average = 0.22 %) in the Kuhfeng Formation(Table 2).The TOC profile shows a weak decreasing trend upward (Fig. 3). H contents range from 0.01 % to 0.04 %in the Chihsia Formation,and range from 0.01%to 0.62 %(average = 0.21 %) in the Kuhfeng Formation. Atomic H/C ratios range from 0.32 to 34.34 (average = 12.12).Bulk organic-carbon isotopic compositions (δ13Corg) range from - 27.9 to - 24.1 ‰ (average = - 25.8 ‰). The δ13Corgprofile shows a peak in the lower Kuhfeng Formation and then a gradually decreasing trend upward,which is similar to the δ13Corgprofile trend at Chaohu(Fig. 3).

5 Discussion

5.1 Diagenetic effect on δ13Corg

Owing to selective preservation of organic matter through microbial decomposition during early diagenesis, sedimentary bulk organic matter δ13Corgundergoes diagenetic alteration (Benner et al. 1987; Lehmann et al. 2002).However,there is a slight change(less than 1–2 ‰)of bulk organic-carbon isotopic composition during early diagenesis (Galimov 2006). Thermal alteration during late diagenesis tends to remove13C-depleted hydrocarbons and increase the bulk organic carbon isotope compositions(Hayes et al. 1999;Tocque´ et al.2005). Biomarker studies on the organic matter in the Kuhfeng Formation in the Lower Yangtze area show a small change of thermal maturation (Geng and Wei 2019). This suggests that the significant δ13Corgvariations (＞2.5 ‰) here are likely unrelated to thermal alteration. Lehmann et al. (2002)concluded that bulk organic-carbon isotopic compositions are well correlated with primary isotope values,suggesting that sedimentary bulk organic-carbon isotope composition recorded surface waters signals. Furthermore, thermal alteration may only elevate absolute δ13Corgvalues but not destroy secular trends (Des Marais et al. 1992). Therefore,the carbon composition of the organic matter here can be used to trace primary environmental changes(e.g.,Ostrom et al. 1997; Hodell and Schelske 1998; Hayes 2004; Galimov 2006).

5.2 Global environmental controls on δ13Corg

In addition to diagenesis effect, several primary environmental factors could cause shifts in bulk organic-carbon isotopic composition, including the mixture of marine organic and terrestrial C3carbon, initial δ13C of carbon source, photosynthetic fractionation and rate of organic matter burial (Arthur et al. 1985; Hayes et al. 1999;Newton and Bottrell 2007). First, in terms of organic matter provenance, terrestrial C3plant photosynthesis yields more(e.g. ～10 ‰) depleted δ13C values of leaves than marine photosynthesis owing to the13C-depleted carbon source in the atmosphere relative to DIC and faster diffusion of CO2in the air relative to that in water(Newton and Bottrell 2007; Ohkouchi et al. 2015). Therefore, proportions of marine and terrestrial organic matter influence organic-carbon isotopic composition in sediments (Arthur et al. 1985; Hayes et al. 1999), and marine organic matter δ13C is higher than terrestrial organic matter (Deines 1980). However, marine organic carbon that is Cretaceous or older is isotopically lighter than land-derived carbon(Dean et al.1986;Galimov 2006;see also Cui et al.2017).Permian higher plants show heavier δ13C values than coeval marine-derived organic matter (Hayes et al. 1999;Krull 1999; Ward et al. 2005; Kraus et al. 2013). This opposite trend can be interpreted by (1) marine organiccarbon δ13C values older than Cenozoic are more negative than marine organic carbon in most Cenozoic sediments because of higher atmospheric CO2concentrations and thus higher isotopic fractionation (Arthur et al. 1985; Popp 1989; Hayes et al. 1999); (2) the whole-plant δ13C values in sedimentary organic matter are heavier than leaves(Badeck et al. 2005; Cernusak et al. 2009); (3) gymnosperms δ13C values are more13C-enriched than angiosperms which bloom in Cenozoic (Diefendorf et al. 2010),and (4) vascular plants which bloom in Carboniferous and Permian are more13C-enriched than non-vascular plants because of the13C-enriched roots and woody stems (Cui et al. 2017).

Table 1 TOC, atomic H/C ratio and organic-carbon isotopic composition data at Chaohu section, South China

Table 1 continued

In this study, we use atomic H/C ratios to indicate organic matter sources to constrain organic-carbon δ13C variations.H/C ratios mainly reflect organic matter sources(Peters 1986; Talbot and Livingstone 1989) and thermal maturity (Baskin 1997). Marine autochthonous plankton organic matters consisting of proteins, carbohydrates, and lipids have high H/C ratios, while terrestrial higher plant organic matters consisting of lignin, cellulose, and sklero proteins have low H/C ratios(Stuermer et al.1978;Deines 1980). The thermal maturity proxies of biomarkers of the Kuhfeng Formation in the Lower Yangtze area show a small variation upward (Geng and Wei 2019), and thus may not contribute to the significant variation of H/C ratios in this study. Cross-plots between organic matter δ13Corgand H/C ratios show weak or no correlation in this study(Fig. 4), suggesting that organic matter sources exert a small impact on the δ13Corgvariations here. The terrestrial organic detrital grains in the Kuhfeng Formation are small and limited (Tang and Wei 2020), and the phosphorite nodules in the lower Kuhfeng Formation also suggest low terrestrial detritus input(Kametaka et al.2005).Therefore,the gradually decreasing δ13Corgtrends upward in the Middle Permian Guadalupian here is unlikely to be due to changes of organic matter sources, but other factors.

Combining the δ13Corgdata at Chaohu and Xiaolao, the δ13Corgtrend of the Guadalupian in the Lower Yangtze area (South China) is matched by the inorganic-carbon δ13Ccarbtrend in Guizhou (South China, Buggisch et al.2011)(Fig. 5),suggesting that they are subject to the same forcing (Kump and Arthur 1999). Both of them show similar trends of a peak in the lower Roadian (see also Zhang et al. 2020) and then a gradual negative shift of δ13Cupward (Fig. 5). Besides, inorganic-carbon isotopic composition from the Roadian to the Wordian in Japan(Musashi et al. 2010) and organic-carbon isotopic composition in the Guadalupian in Australia with low-resolution sampling (Birgenheier et al. 2010) also show a similar trend.Thus,the covariance of δ13C measured on carbonate rocks and organic matter can argue for δ13C changes during the global carbon cycle (Arthur et al.1985; Magaritz et al.1992; Holser 1997).

The global carbon isotopic composition changes in the Guadalupian are mainly affected by initial δ13C values of DIC, marine photosynthetic fractionation and rate of organic burial (e.g., Arthur et al. 1985; Arens et al. 2000).Marine photosynthetic fractionation is related to DIC concentration (Dean et al. 1986) and organic matter amount during the biological pump process (Newton and Bottrell 2007). The Roadian cooling (Saunders and Reichow 2009; Chen et al. 2013) enhanced the oceanic upwelling intensity (Zhang et al. 2019), increasing the primary productivity and thus a peak in both inorganic-and organic-carbon isotopic compositions during this time(Fig. 5). Increasing pCO2and gradual global warming occurred from the late Roadian to the end-Guadalupian(Saunders and Reichow 2009; Chen et al. 2013), which resulted from basalt eruption in the opening of the Neo-Tethys Ocean (Muttoni et al. 2009), widespread subduction-related volcanism (Saunders and Reichow 2009), and the Emeishan large igneous province (LIP) eruption (Ali et al. 2005). Associated with the assembly of supercontinent Pangaea, the decreasing mountain belts (and the resultant reducing CO2consumption during silicate rocks weathering) also contribute to the increasing CO2in the Middle Permian (Saunders and Reichow 2009). The volcanic-release12C-enriched CO2would lower initial δ13C of marine DIC and contribute to the decreasing trend of δ13C from the Roadian to the Capitanian (Fig. 5). Increasing pCO2in the Guadalupian atmosphere could increase marine DIC concentration(e.g.,Brewer 1997;Wu et al.2008).Greater CO2availability can enhance the extent of carbonisotope fractionation by marine plankton (Dean et al.1986), and thus produces more negative δ13C values of organic matter. This may contribute to the observed negative shift of organic-carbon isotopic composition from the late Roadian to the Capitanian in the Guadalupian(Fig. 5).However, both the increased CO2concentration and the lower initial δ13C value of DIC which resulted from volcanic-release CO2can only account for minor changes of 1–2 ‰ in the δ13C of marine organic matter (Dean et al.1986). The ＞2.5 ‰ negative shift of δ13Corgin the Guadalupian (Fig. 5) can be mainly controlled by other factor(s).

Fig. 5 δ13Corg and δ13Ccarb correlation between the Lower Yangtze area (South China) and Guizhou (South China), and the atmospheric CO2 concentration (Saunders and Reichow 2009), sea-level changes (Haq and Schutter 2008) and fusulinacean species diversity (Groves and Wang 2013) in the Guadalupian. The timescale and carbonate-carbon isotope of Guizhou are from Shen et al. (2019) and Buggisch et al. (2011).Organic-carbon isotope composition curves in this study are smoothed using LOESS smoothing procedure

During the zenith process of the supercontinent Pangaea assembly in the Guadalupian (Isozaki 2009), global sealevel drop (Haq and Schutter 2008) caused drainage of coastal swamplands and reduced organic matter burial(Berner 2005; Peters-Kottig et al. 2006), resulting in more12C enrichment in DIC and thus a decrease of δ13CDICand δ13Corgvia photosynthesis in the Guadalupian as seen in this study (e.g., Birgenheier et al. 2010; Buggisch et al.2011, 2015). Meanwhile, the decomposition and oxidation of organic matter to CO2during the sea-level drop could also contribute to the long-term δ13C negative shift in the Guadalupian (e.g., Berner 2002), although some reports(Shields and Mills 2017; Zhang et al. 2020) suggest that a reduction of carbonate production during regression may drive a positive δ13C excursion. The latter may offer a smaller contribution because dissolved inorganic carbon is smaller than dissolved organic carbon in rivers(Prokushkin et al. 2011). Dissolved organic matter (DOM) can form aggregates such as zooplankton shells and fecal pellets through faunas uptake (e.g., Alldredge and Silver 1988;Passow and Carlson 2012) to be fast-sinking particles (see review in Wakeham and Lee, 2019). The carbon-isotopic compositions of DOM are depleted in13C(Galimov 2006).Therefore, the sinking and burial efficiency of DOM can influence the carbon-isotopic compositions of marine carbonate and organic matter (D’Hondt et al. 1998). During the Guadalupian, the gradual decreasing diversity of benthic faunas led to decreased uptakes of DOM and thus may also contribute to the small δ13C negative shift in the Capitanian in the Guadalupian (Fig. 5).

5.3 Implications for the Guadalupian mass extinction

The gradual negative shift of δ13C in the Guadalupian(Fig. 5) and rapid δ13C negative excursion at the end-Guadalupian(Wignall et al.2009;Bond et al.2010b;Shen et al. 2013) coincide with gradual extinction (Fig. 5,Groves and Wang 2013) and rapid extinction (Bond et al.2010a; Shen and Shi 2009), respectively. This suggests a causal relationship. As mentioned above, the gradual negative shift of δ13C from the Roadian to the middle Capitanian is mainly caused by decreased organic matter burial on land and in the ocean,which result from the global sealevel drop and oxygen-depletion-caused benthos decrease,respectively. The secondary causes for the gradual δ13C negative shift are volcanic-gases release and decreased mountain belts, and thus decreased silicate weatheringconsumption of CO2during this period (Chen and Xu 2019).Therefore,the gradual glacio-eustatic sea-level drop(Haq and Schutter 2008; Qiu et al. 2014) and reducing conditions of marine water masses (Wei et al. 2016,2019)may be the causes for the gradual diversity-decreasing process in the Early and Middle Guadalupian. The significant δ13C negative shift from the middle-late Capitianian to the end-Guadalupian inherits from the gradual negative shift in the Early and Middle Guadalupian, probably sharing the same causes for δ13C variations of the Early-Middle Guadalupian. In other words, the significant δ13C negative shift in the Late Guadalupian may be caused by the rapid global sea-level drop (Haq and Schutter 2008)and anoxia events (Zhang et al. 2015; Wei et al.2016, 2019), which may be also the main causes for the rapid mass extinction during this time. Oxygen depletion can trigger sublethal to lethal effects (Riedel et al. 2008)and global sea-level drop caused geographic habitat loss(Haq and Schutter 2008).The combined effect of these two events likely led to the mass extinction.

6 Conclusions

We reconstructed marine organic-carbon isotopic composition changes in the Guadalupian through the study of two sections(Chaohu and Xiaolao)in the Lower Yangtze area,South China. Our results show that the δ13Corgchanges in the Guadalupian can be matched by δ13C changes of carbonate rocks or organic matter in South China and other places in the world. Both inorganic- and organic-carbon isotopic compositions have similar trends in the Guadalupian,showing a peak in the Roadian and a gradual negative shift from the late Roadian to the middle Capitanian. The positive peak in the Roadian was probably due to high productivity caused by intensive upwelling from cooling during this time. The gradual negative shift of δ13C was mainly caused by decreased organic matter burial on land and in the ocean, which result from the global sea-level drop and anoxia-caused benthos decreasing, respectively.The secondary causes for the gradual δ13C negative shift are volcanic-gases releasing,decreased mountain belts and thus decreased silicate weathering-consumption of CO2.The gradual negative shift of δ13C coincides with the gradual extinction in the Guadalupian. Therefore, global sea-level drop and marine reducing conditions may be the main causes of the gradual diversity decrease in the Early and Middle Guadalupian.

AcknowledgementsWe thank two anonymous reviewers for their constructive comments. This work was funded by the National Natural Science Foundation of China (No. 41762003).

Compliance with ethical standards

Conflict of interestThe authors declare that there is no confliction of interest regarding the publication of this paper.