Journal Pre-proof Global carbon cycle perturbations triggered by volatile volcanism and ecosystem responses during the Carnian Pluvial Episode (late Triassic) Ziheng Li, Zhong-Qiang Chen, Feifei Zhang, James G. Ogg, Laishi Zhao PII: S0012-8252(20)30450-5 DOI: https://doi.org/10.1016/j.earscirev.2020.103404 Reference: EARTH 103404 To appear in: Earth-Science Reviews Received date: 24 June 2020 Revised date: 27 September 2020 Accepted date: 9 October 2020 Please cite this article as: Z. Li, Z.-Q. Chen, F. Zhang, et al., Global carbon cycle perturbations triggered by volatile volcanism and ecosystem responses during the Carnian Pluvial Episode (late Triassic), Earth-Science Reviews (2020), https://doi.org/10.1016/ j.earscirev.2020.103404 This is a PDF file of an article that has undergone enhancements after acceptance, such as the addition of a cover page and metadata, and formatting for readability, but it is not yet the definitive version of record. This version will undergo additional copyediting, typesetting and review before it is published in its final form, but we are providing this version to give early visibility of the article. Please note that, during the production process, errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain. © 2020 Published by Elsevier. Journal Pre-proof Global carbon cycle perturbations triggered by volatile volcanism and ecosystem responses during the Carnian Pluvial Episode (Late Triassic) Ziheng Lia, Zhong-Qiang Chenb *, Feifei Zhangc, James G. Oggb, d, e, Laishi Zhaoa * a State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences, Wuhan, Hubei 430074, China State Key Laboratory of Biogeology and Environmental Geology, School of Earth Sciences, of b China University of Geosciences, Wuhan, Hubei 430074, China State Key Laboratory of Mineral Deposits Research and School of Earth Sciences and ro c State Key Laboratory of Oil and Gas Reservoir Geology and Exploitation, Chengdu re d -p Engineering, Nanjing University, Nanjing 210023, China e lP University of Technology, Chengdu, Sichuan 610059, China Department of Earth, Atmospheric and Planetary Sciences, Purdue University, West na Lafayette, IN 47907-2051, USA Abstract Jo (LSZ)) ur (*Corresponding authors; emails: zhong.qiang.chen@cug.edu.cn (ZQC); lszhao@cug.edu.cn The Carnian Pluvial Episode (CPE) was a dramatic climatic event during the early Late Triassic. The CPE has been recognized worldwide and is marked by the termination of carbonate platform successions and by pronounced negative δ13C excursions (denoted as the CPE excursion). The onset of the CPE has been proposed to be linked with the volatile eruption of the Wrangellia Large Igneous Province (W-LIP). However, this extreme climatic event remains disputed in terms of its precise global correlation, timing of onset, duration, and global magnitude. We compiled a database of 13 conodont biozone-controlled Journal Pre-proof stable-isotope reference sections throughout the Tethyan region. After a reexamination of previously published conodont taxonomy from each section, the statistics on the conodont assemblages/zones yield a global set of 17 high-resolution conodont Unitary Association Zones (UAZs) spanning the entire Carnian successions. A set of age-tie points placed an age model on this global UAZ scale. We added paired δ13Ccarb and δ13Corg data from two sections in South China to other records in this global database, and then normalized all of carbon-isotope datasets. A pronounced negative δ13Ccarb excursion with a magnitude of -2 ± 0.5 ‰ (from an average of ~3 ‰ to ~1 ‰) is evident in these normalized trends and is ro recognized around the entire Tethys realm. The CPE excursion coincides with conodont -p UAZ-4 (~234 Ma) through UAZ-8 (within the Mazzaella carnica and Paragondolella re praelindae zones, Julian2) and has a duration of ca. 1.5 Myr. We note that an apparent lP delayed onset of the CPE excursion in South China relative to elsewhere in the Tethys as suggested by previous studies may have been an artifact caused by incomplete na conodont-stratigraphic records and/or irregularities in taxonomic identification. A carbon and phosphorus cycle model highlight the strong relationship between the eruption of the W-LIP ur and CPE excursion; however, the estimated amount of CO2 volatiles released by the W-LIP Jo would directly account for only ~25% of the total amount of light-carbon required to explain the combined magnitude and duration of the global CPE excursion. Keywords: Carbon isotopes; carbonate factory demise; conodont Unitary Association Zones; Wrangellia LIP; C-P cycle model; South China 1. Introduction Journal Pre-proof The Triassic was a critical period for the evolution of marine ecosystems. The attainment of a Middle Triassic stable ecosystem (Chen and Benton, 2012) was temporarily disrupted by an extreme climatic event known as Carnian Pluvial Episode (CPE) during the mid-Carnian of early Late Triassic (Simms and Ruffell, 1989). The CPE is a strongly humid period characterized by large runoff and enhanced continental weathering in many terrestrial settings (Ogg, 2015; Ruffell et al., 2016). An abrupt lithological change, the Reingraben turnover of (Schlager and Sch llnberger 1974), from carbonate-dominated to siliciclastic-dominated facies has been recognized around the Tethys realm (details see Fig. 1C). The CPE is also Jo ur na lP re -p ecosystems (e.g. Simms and Ruffell et al., 1989). ro characterized by significant biotic crises and/or turnovers in both marine and terrestrial Figure 1. Summary of major climatic changes, and the global stratigraphic correlations of the Carnian Stage. (A) Triassic climatic trends with major humid (dark blue) and arid (orange) intervals projected onto the Triassic age model of Ogg and Chen (2020). (B) Carbon isotope trends, conodont Unitary Association Zones (UAZs), and magnetostratigraphy of the entire Carnian Stage. (C) Approximate stratigraphic correlations of key CPE reference sections and/or regions based on biostratigraphy profiles and major lithologic transitions. Occurrences of widespread amber are shown by orange-colored droplets. Datum sources: (1-2) Stefani et al. (2010) and Seyfullah et al. (2018); (3) Zhang et al. (2020); (4-5) Zhang et al. (2015) and Shi et al. (2017); (6-12) Kozur and Bachmann, (2010), Hornung and Brandner (2005), Dal Journal Pre-proof Corso et al. (2015), Dal Corso et al. (2018), Furin et al. (2006), Rigo et al. (2007), Muttoni et al. (2014); (13) Nakada et al. (2014); (14) Lukeneder et al. (2012); (15-16) Hornung et al. (2007b) and Bhargava et al. (2004); (17) Sun et al. (2019); (18-19) Marsicano et al. (2015), Ezcurra et al. (2017); (20) this study. Abbreviations: NCA = Northern Calcareous Alps; TR = Transdanubian Range; Fm. = Formation. During the past two decades, an increasing attention of geoscientists to this extreme climatic event resulted in the discovery of many other coeval responses during the relatively of short CPE humid interval (Fig. 1; Stefani et al., 2010), including: (1) Expansion of the ro hygrophytic vegetation around the Tethys realm. On the basis of palynological associations -p and the distribution of terrigenous levels, the CPE was subdivided into several pulses of re humid climate in some regions (Roghi et al., 2010; Mueller et al., 2016b). However, palynological records also suggested that Carnian Pluvial Episode could not be recognized in lP inner parts of Pangea (Baranyi et al., 2018). (2) Extensive distribution of paleo-karst features na (e.g. Stefani et al., 2010; Shi et al., 2017). (3) An abundance of amber during the CPE. This is considered to be its first appearance on a global scale (Seyfullah et al., 2018). (4) Variation in ur the mineralogical compositions of pelagic chert (Nakada et al., 2014). (5) Disappearance of Jo oolites in many sections across the South China, indicating demise of the ooid carbonate factory (Li et al., 2017; Jiang et al., 2019). 1.1. Critical period for the evolution of Mesozoic ecosystems: a biotic crisis or a turnover？ Life took a long time (~8 Myr) to recover after the end-Permian mass extinction, and a stable, complex ecosystem did not re-emerge until the beginning of the Middle Triassic (Chen and Benton, 2012; Zhang et al., 2018). Even though the Carnian Age of the early Late Triassic was a time of wide-ranging evolutionary innovation, a Tethys-wide Carnian Crisis Journal Pre-proof and global CPE interval may have been one of the most severe biotic crises in the history of Earth life (Hornung et al., 2007b; Bond and Grasby, 2017). The CPE interval witnessed the major extinction of reef builders, bivalves, crinoids, bryozoans, ammonoids, and conodonts (Simms and Ruffell, 1989; Simms et al., 2018; Orchard, 2013, 2014; Chen et al., 2016; Rigo et al., 2018). Conodonts may have undergone a two-phase crisis during the CPE, with the first occurring during the early Julian and the of second at the Julian-Tuvalian boundary (Fig. 2; Chen et al., 2016). In addition to extinctions, the CPE also coincided with the appearance and surge in abundance of several important ro Mesozoic lineages, including pelagic calcispheres and scleractinian corals in the ocean -p (Stanley, 2003; Furin et al., 2006; Preto et al., 2017; George et al., 2018), and early dinosaurs re and mammals on land (Bernardi et al., 2018; Benton et al., 2018). In particular, the lP Ischigualasto dinosaurian forms, which are thought to be the oldest undisputed members of the Dinosaurian (Benton, 1991; Ezcurra et al., 2017), are radioisotopically dated as Carnian na in age (Marsicano et al., 2015; Langer et al., 2018), and appear to be strongly associated with the CPE (Bernardi et al., 2018). ur A challenge is that the biostratigraphic age calibration of the CPE in a few sections of Jo carbonate-siliciclastic facies appears to differ between current conodont and ammonoid biostratigraphy. For example, in the Maantang Formation unit2 of Sichuan Basin, the observed conodonts indicate an earliest late Carnian age (Tuvalian substage), whereas the ammonoids indicate the lower Carnian (Julian2) substage (Shi et al., 2017). The onset timing and the duration of the CPE, and associated biotic evolutions have been dated differently from region to region by various authors (e.g. Zhang et al., 2015, 2018; Simms et al., 2018). A contributing factor is that there is not yet a stable conodont zonation. Currently, nine conodont zones are recommended for the entire Carnian Stage in Tethys realms (Fig. 2; Rigo Journal Pre-proof et al., 2018), whereas eleven conodont zones are recognized worldwide (Orchard, 2013, 2014, na lP re -p ro of but see Chen et al., 2016). Jo ur Figure 2. Potential correlations of conodont zonations (Tethys and North America), ammonoid zones (Tethys), conodont generic distribution, and richness of conodont genera and species. Modified after Chen et al. (2016) and Rigo et al. (2018). Abbreviations: P. = Paragondolella, Pa. = Parvigondolella, Pa. p = Parvigondolella prominens, Pa.b = Parvigondolella beattyi, Ma. = Mazzaella, Me.d = Metapolygnathus dylani, H. = Hayashiella, C. = Carnepigondolella, E. = Epigondolella, Ne. = Neocavitella, Ac. = Acuminatella. The purple background represents the estimated CPE interval. 1.2. Environmental perturbations and volatile volcanism Environmental perturbations during the CPE are synchronous with one or more pronounced negative δ13C excursions with an amplitude of ~2–4‰ (Dal Corso et al., 2012, 2018; Muttoni et al., 2014; Mueller et al., 2016a; Sun et al., 2016, 2019; Miller et al., 2017). Oxygen-isotope analyses from conodont bioapatite indicated that seawater temperatures Journal Pre-proof during the CPE may have warmed by 2-5 ℃ (Trotter et al., 2015). Pyrite framboid analysis revealed that multiple anoxia prevailed in marine environments during the mid-Carnian (Sun et al., 2016). Previous studies have speculated that frequent and intense tectonic movements (Hornung and Brandner, 2005), volcanic activity (Furin et al., 2006), or their combined effects during the Late Triassic may be the main cause for the CPE. Furin et al. (2006) noted that the few radioisotopic dates on Carnian successions indicate of that the CPE may coincide with the eruption of the Wrangellia Large Igneous Province (W-LIP), which today is mainly accreted to the northeastern margin of the Pacific Ocean. ro Therefore, they proposed that the W-LIP was the likely trigger of the CPE. Dal Corso et al. -p (2012) obtained pronounced negative carbon-isotope excursions in organic carbon derived re from both total organic matter and terrestrial/marine fossil molecules from the CPE interval. lP They suggested that the required release of a great amount of light carbon was triggered by volatile volcanism, and the CPE δ13C excursions were directly caused by the eruption of the na W-LIP. Following these initial studies, negative excursions in both δ13Ccarb and δ13Corg have been globally reported from the CPE interval and assumed to be linked with the W-LIP in ur origin (Sun et al., 2016, 2019; Preto et al., 2017; Dal Corse et al., 2018). Jo This implied model for the onset of the CPE is similar to the scenario of the marine carbon cycle around the Permian-Triassic transition (Payne et al., 2004; Bond and Grasby, 2017), in which the injection of relatively light CO2 into the ocean-atmosphere system contributed both to a warmer, more humid world and to the carbon-isotope excursions. However, the CPE carbon-isotope excursions reported in most studies are represented by bulk δ13Corg data of whole rocks, which also raises the possible risk that the variation in the marine vs. terrestrial components of organic carbon during the lithologic facies change may have biased the interpretation of the magnitude of the original marine carbon-isotope signals. Journal Pre-proof Thus, the origin for all δ13Corg negative excursions needs strong evidence; and a modelling between carbon isotopes and volatile volcanism may test the role of the W-LIP. 1.3. Goals of this study However, many aspects of the CPE remain enigmatic. The debated issues mainly focus on three questions: (1) what are the characteristics and significance of the Reingraben of turnover from carbonate-dominated to siliciclastics-dominated facies in eastern Tethys region? (2) what are the timings and tempo of the CPE carbon isotope excursions? (3) what was the ro magnitude of the carbon isotope excursions and their relationship with volatile volcanism? -p In order to examine some aspects of these issues in our study, we compiled a re high-resolution conodont biostratigraphic and a carbon-isotopic stratigraphic database of lP reference sections throughout the Tethys realms (Appendix A. Supplementary data) to precisely document the timing and extent of the Carnian Pluvial Episode (Fig. 3). We na investigated two additional Carnian sections from Sichuan, South China: Guanyinya (HWG) and Qingyangou (HWQ) to enhance this compilation. Major and trace element data and TOC ur contents derived from these two South China sections were also employed to quantify aspects Jo of the Reingraben turnover in the eastern Tethys region. Figure 3. Global paleogeography during the Carnian (ca. 237-227 Ma) (modified after Li et Journal Pre-proof al., 2019), showing localities of 15 reviewed reference sections and two additional study sections; see the text of Section 3.1 for details on data and sources for these 17 sections. NC = North China block; SC = South China block; Neo-T = Neo-Tethys; strati. = stratigraphy. To provide the necessary precise age constraints for global correlation of the CPE carbon isotopic excursions, we employed the conodont unitary association zone (UAZ) method, which is a powerful biostratigraphic tool for age control of stratigraphic units (Guex, 1991; of Brosse et al., 2016). Based on conodont biostratigraphic datasets collected from many key ro published reference sections, this enabled the establishment of worldwide high-resolution -p UAZs throughout the entire Carnian Stage. These UAZs, together with magnetostratigraphy (Zhang et al., 2020) and an updated δ13Ccarb database, were employed to constrain the onset re timing and duration of the CPE excursion. The paired δ13Ccarb and δ13Corg data from the two lP sections, combined with published carbon isotope data, suggest a single global (Tethys-wide) and synchronous δ13Ccarb negative excursion during the CPE interval (Julian2). na We then applied a carbon-phosphorus (C-P) cycle model (Clarkson et al., 2018; Zhang-F et ur al., 2020) to this global carbon-isotope dataset to investigate the marine carbon cycle during Jo the CPE, to test the possible relationship between the CPE excursion and contemporary volatile volcanisms (i.e., W-LIP), and to unravel the potential environmental impacts of the CPE (e.g., enhanced temperature, ocean anoxia, and appearance of the keystone species such as dinosaurian and calcareous nannoplankton). Hence, we can conclude that the CPE interval played an important role in the evolution of early Mesozoic ecosystems. 2. Geologic and stratigraphic settings of the South China sections The South China block was located in the eastern Paleo-Tethys during the Late Triassic (Scotese, 2017), and its central through western region was occupied by the Yangtze Platform. Journal Pre-proof The Upper Triassic marine successions are widely exposed in the western part of the Yangtze Platform and on its northern margins. Marine CPE successions are well recorded in the Qingyangou (HWQ; N31°27′46.85′′ E104°09′35.40′′) and Guanyinya (HWG; N31°28′10.32′′ E104°08′49.52′′) sections of northwestern Sichuan Province (Fig. 3). The HWQ and HWG sections are both located in the Hanwang area of the modern northwestern Sichuan Basin, which was formerly the northern margin of the Upper Yangtze of Platform (see Shi et al., 2017 for detailed location map). The Carnian strata of the Maantang Formation was subdivided into four units (Shi et al., 2017; Jin et al., 2018), and the transition ro between Unit-1 and Unit-2 record the Reingraben turnover with a lithologic change from a -p lower reef/bioherm facies limestone to upper offshore facies of mudstone and siltstone. This re interval of the Maantang Formation yields abundant conodonts of Quadralella lP polygnathiformis, Paragondolella praelindae, and Mazzaella carnica as well as ammonoids of Thisbites sp. and Anatomites sp., thereby indicating a mid-Carnian age (Shi et al., 2017; na Jiang et al., 2019). Jo ur 3. Materials and analytical methods 3.1. Research materials A total of 13 sections with δ13Ccarb profiles were compiled (Fig. 3): (1-2) the South China study sections of this study, (3-4) Longchang section (LC, Sun et al., 2016) Yongyue section (YY, Sun et al., 2019), (5) Waidi Mayhah (WM, Oman) sections (Sun et al., 2019), (6) Guling2 section (GL, Hornung et al., 2007a), (7) Aghia Marina (AM, Greece) section, (8-10) Pizzo Mondello section (PM, Italy), Lavarella, and Fedares sections (LA and FD, Muttoni et al., 2014; Keim et al., 2006), (11) Silicka Brezonva section, Slovakia (SB, Korte et al., 2005), Journal Pre-proof and (12-13) Lunz-Polzberg and Freygutweg sections, Germany (LP, and FT, Hornung and Brandner, 2005; Hornung et al., 2007b). A total of 14 sections were compiled to generate high-resolution conodont unitary association zones (UAZs): LC, WM, HWG, HWQ (Jiang et al., 2019), LSK, WY, AM, PM, FT, YY, (14) Guri Zi (GZ, Albania) sections (Muttoni et al., 2014), (15-16) British Columbia and Black Bear Ridge sections (Orchard, 2014), and (17) Guling 1 section (India, Bhargava of et al., 2004). A total of 8 sections with δ13Corg profiles were compiled: HWG, and HWQ sections from ro this study, LC, Lunz section (Mueller et al., 2016a), and Heiligkreuz, Milieres-Dibona, Cave re -p del Predil sections, and Balaton, Bfü-1 and Met-1 cores (Dal Corso et al., 2018). lP 3.2. Unitary association method The Unitary Association method (Guex, 1991) is a quantitative and deterministic method na that produces biozonations based on discontinuous maximal association zones. This method is described in detail by Guex (1991). The dedicated statistical tool is integrated within the ur well-known freeware paleontological analysis software package of Past (Hammer, 2013; also Jo see Brosse et al., 2016, and reference therein). Our purpose in applying the UA method is to provide a global correlation tool for the Carnian Stage, because many of the conodont zonation sequences in the compiled sections are incomplete (e.g., sections that lack the occurrence of Mazzaella carnic and/or Paragondolella praelindae). In the UA method, species with a discontinuous range in a section are assumed to range from their local first occurrence (FO) to their last occurrence (LO) to obtain a complete and logical conodont assemblage, especially during the CPE interval. Journal Pre-proof The newly derived UAZs spanning the entire Carnian Stage are partially constrained by the reporting of the biostratigraphic data and taxonomy as provided in the published conodont-stratigraphic studies. Accordingly, revisions were required to ensure the consistency and standardization of the taxonomic data, including: (1) the FO of the Quadralella noah in Longchang section (Sun et al., 2016) was revised upward from ~12 m to a level at ~29 m; (2) the Para. praelindae zone in the Yongyue section that had been marked of by affinis species in Sun et al. (2019) was excluded; and (3) all taxa that had qualifiers of undetermined (indet.), compared (cf.), and transitional (trans.) were excluded from the ro database. The current Carnian conodont stratigraphic database is not of the same magnitude -p as some other stratigraphic databases (e.g. the detailed Permian-Triassic transition compiled re by Brosse et al., 2016). Therefore, we cannot employ some more advanced statistics such as lP detailed cross section comparison. We cannot fully rule out the possibilities that artifacts of sample spacing and taxonomic assignments may slightly deviate the first and/or last ur the UAZ method. na occurrences of some conodont taxa in some sections, which may further impact the results of Jo 3.3. Sampling and preparation A total of 201 samples were collected for analysis in this study: 108 samples from the 73-m-thick outcrop at HWG, 93 samples from the 36-m-thick outcrop at HWQ. Samples were collected at high resolution (~10 to 30-cm spacing) through the transition of the lithology and at a uniform spacing of ~50 to 200 cm throughout the rest of each section. Following the removal of weathered surfaces and obvious veins, the chosen pieces from the crushed samples were powdered and passed through a ~200 mesh for further testing. Journal Pre-proof 3.4. Major and trace elements Major and trace element analysis used ~10 g of powdered samples. Major and trace element determination was done using standard X-ray fluorescence (XRF) and with an Agilent 7500a ICP-MS instrument, respectively, at the China University of Geosciences in Wuhan (CUG). The precision of the major and trace element analyses is estimated to better than ca. 1 % and 5 %, respectively. The analytical results are tabulated in Appendix A. of Supplementary data. ro 3.5. Carbon isotopes -p In preparation for carbonate carbon isotope (δ13Ccarb) analysis, ~80 to 120 μg of powder re samples were placed in a 10 ml Na-glass vial, sealed with a butyl rubber septum, and reacted lP with 100 % phosphoric acid at 72 °C after flushing with helium. The evolved CO2 gas was analyzed for δ13C and δ18O using a MAT253 mass spectrometer at the State Key Laboratory na of GPMR CUG. Isotope values are reported as per mil variation (‰) relative to Vienna Pee Dee belemnite (VPDB) standard. The analytical precision was better than 0.04‰ for δ13C ur and δ18O based on duplicate analyses of the national reference standard GBW-04416 (δ13C = Jo +1.61 ‰; δ18O = −11.59 ‰) TB-2 (δ13C = -6.06 ‰; δ18O = -24.12 ‰). For organic carbon isotope (δ13Corg) analysis, aliquots of powdered samples of between 5 and 30 mg (depending on TOC content) were first decarbonated in 50 ml tubes overnight through acid fumigation using concentrated HCL (12 M). After drying, one or more drops of 1 M HCL were added to the captures to ensure complete carbonate removal. The carbonate-free residue was then washed by 18.2 MΩ water for three-to-five times to change the solution to neutral pH (≥ 6.0). After drying at 30°C for 14 h, samples were stored in airtight 50 ml tubes for later analysis. Organic carbon isotopes were analyzed using a MAT253 mass spectrometer at SKL-GPMR. Isotope values are reported in per mil relative to Journal Pre-proof VPDB. Standards of GBW-04408 (δ13Corg = -36.91 ‰), GBW-04407 (δ13Corg = -22.43 ‰), ZT-4 (δ13Corg = -27.00 ‰) and Urea (δ13Corg = -28.25 ‰) were used to monitor the external and internal uncertainties, which were better than 0.2 ‰ for δ13Corg and 0.1 % for TOC. To monitor the reproducibility of organic-rich and organic-poor samples, duplicates were analyzed for every twelfth sample; and their reproducibility in δ13Corg and TOC was better of than 0.5 ‰ and 0.2 %, respectively. 3.6. C-P cycle model ro The total of 181 δ13Ccarb and 168 δ13Corg values generated in this study were combined -p with published δ13Ccarb data (n = 468) after normalization for burial diagenetic trends (after re Hasiuk et al., 2016). A simple carbon and chosphorus (C-P) cycle associated model lP (Clarkson et al., 2018; Zhang-F et al., 2020) was applied to this suite of normalized and smoothed δ13Ccarb values. The C-P cycle model is a simplified version of the GEOCARB and na COPSE family of models (Berner and Kothavala, 2001; Bergman et al., 2004). This model focuses on shorter-term carbon cycle processes (e.g. ~0.5 to 5 Myr), and is controlled by ur normalized forcing parameters, including degassing (D), uplift driving erosion (E), extent of Jo land colonization by vegetation (V), lithologic or vegetation effects on silicate rock weatherability (W), and CO2 emission from LIPs volcanism (LIP) (after Clarkson et al., 2018). In the steady state, the values of these forcing were adopted from the COPSE model (Bergman et al., 2004). We assume that the first four forcings remain constant (D, E, V, W = 1) during the relatively short span of the Carnian (237–227 Ma spanning 10 Myr), and the model was only perturbed by CO2 emissions into the atmosphere from LIP volcanism. We note that another ‘B’ forcing from marine biological evolution has been suggested in COPSE/GEOCARB that affects the apportioning of carbonate burial between deep and shallow seas; which, in turn, has a positive correlation with the flux of degassing of carbonate Journal Pre-proof carbon. This B forcing is a delayed feedback that is enhanced when an increase in calcareous plankton abundance shifts the fraction of inorganic carbon buried as deep-water carbonates; which is later recycled back into the atmosphere-ocean system by thermal decomposition during subduction (Volk, 1989). Therefore, this B forcing should also be associated with the CPE, which saw the initial rise of significant calcareous nannoplankton (Preto et al., 2013). However, the magnitude of this B forcing may initially have been rather weak during the of middle Carnian when only 4 families of coccoliths and nannoliths appeared, compared to its importance in the Jurassic-Cretaceous as calcareous nannoplankton diversified into 30 ro families (Bown et al., 2004) and pelagic carbonate began to dominate oceanic sediments. The -p increasing abundance of calcareous nannoplankton would increase the flux of degassing re carbonate carbon, thereby increasing the CO2 values and temperature (GEOCARBIII, Berner lP and Kothavala, 2001). Several significant environment changes are predicted using the model after inducing na disturbances to the steady-state conditions, including: (1) temperature change, controlled by changes in CO2 and solar luminosity (GEOCARB III); (2) oceanic anoxia, controlled by O2 ur changes through remineralization of organic matters and by organic matter production, which, Jo in turn, is dependent on P concentrations; (3) atmospheric oxygen, controlled at a simplified quantitative level by the feedback of organic carbon burial (Clarkson et al., 2018). According to the magnitude and duration of the CPE excursion, we first used δ13Ccarb data normalized to burial-diagenetic trends as a required forcing to compare to that estimated from the CO2 emission flux from W-LIP volcanism during the CPE interval. We then tried a simplified representation of this volcanic emission in the model simulations as a linear increase followed by linear decrease at same rate during 234–233 Ma. This model highlights the relationship between CPE excursions and volatile volcanism; hence, invokes the potential response of the ecosystem to the volcanism during the CPE interval. Journal Pre-proof 4. Results 4.1. “Reingraben turnover” in South China: Characteristics of major element and TOC contents The Reingraben turnover is considered as the most characteristic feature of the CPE. As of suggested by Hornung et al. (2007b) and Rigo et al. (2007), the Reingraben turnover should have different characteristics in different regions according to regional climatic belts, relative ro water depths, and styles of weathering. The onset interval of the Reingraben turnover in the -p study sections is marked as: ~7 to 10 m in the HWQ and ~42 to 51 m in the HWG sections of re Sichuan (Fig. 4). We employed the combined datasets on lithology, Al/Ti ratios, and Ca and lP TOC contents to demonstrate that the Reingraben turnover in Sichuan is marked by at least four upward trends: (1) There is a generally sharp lithologic change from a reef/bioherm na limestone or other relatively pure carbonate facies into mudstone and shale facies in Sichuan. (2) TOC content increases abruptly from extremely low values (average = 0.06%, n = 53) to ur high values (average = 0.63%, n = 42). (3) As expected from the lithologic change, there is a Jo gradual decrease in Ca content to extremely low values (from 33 %, n = 53, to only 7 %, n = 42). (4) Al/Ti ratios increase from 9.1 (n = 53), to 17.6 (n = 42). We should note that there are distinguishing differences between the Sichuan and Guizhou regions (LC, Sun et al., 2016), which suggest that the expression of the Reingraben turnover was governed by different mechanisms in each region. Journal Pre-proof of Figure 4. Al/Ti ratios, and contents of Al2O3, TiO2, Ca and TOC in the (A) HWQ and (B) HWG sections. ro 4.2. Unitary Association Zone of the entire Carnian and calibration of the CPE -p With the standardization of the taxonomy in the published conodont-stratigraphic studies, re there are a total of 48 valid conodont taxa within the Carnian in our global database. These are grouped into 17 Unitary Association Zones (UAZ-1 to UAZ 17; Fig. 5A). The age model lP for the conodont UAZs and associated magnetostratigraphic zones (Zhang et al., 2020; Ogg na and Chen, 2020) suggest that UAZ-1 represents the basal Julian1 chronostratigraphic unit beginning at 237 Ma, although the resolution of UAZ-1 is relatively low due to the limited ur number of reference sections and the irregular first occurrences (FOs) of P. tadpole and P. Jo intermedius. The Julian1-Julian2 boundary (234.5 Ma) is calibrated within the UAZ-2 and UAZ-3. The UAZ-9 and the UAZ-15 represent the Julian-Tuvalian transition (233.5 Ma) and Carnian-Norian boundary (227.3 Ma), respectively (Fig. 5). Both Ma. carnica and Para. praelindae are typically considered to be marker species for the mid-Carnian (Chen et al., 2016; Rigo et al., 2018), with the Julian1-2 boundary calibrated to the middle part of the Ma. carnica Zone. Thus, both of them are important in recognizing and correlating the CPE. Following the UAZ approach, the FOs of Ma. carnica and of Para. praelindae fall in the UAZ-2 and UAZ-4, respectively (Fig. 5A). 17 16 15 14 13 12 11 10 9 8 7 6 5 4 3 ? 2 1 Carnian-Norian Boundary of Julian-Tuvalian Boundary B -p ro s7 s6 s5 s4 s3 s2 ? s1 na lP re C w13 w12 w11 w10 w9 w8 w7 w6 w5 w4 w3 w2 w1 H. tuvalica C. pseudodiebeli Q.ex gr.carpathica Me. praecommunisti Ne. cavitata C. orchardi C. samueli Me. mersinensis E. heinzi C. pseudoechinata E. vialovi E. miettoi E. quadrata Me. dylani Me. communisti Me. primitius Me. parvus E. rigoi E. triangularis C. gulloate Q. oertlii C. zoae A B. mungoenis Q. shijiangjunensis Q.acuminatus P. inclinata G. malayensis G. tethydis P. intermedius P. foliata P. tadpole Q.cf.maantangensis Q. jiangyouensis Q. polygnathiformis Q. langdaiensis Q. wayaoensis Q. zonneveldi Q. xinpuensis Ma. carnica Q. hanwnagensis Ma. julica P. praelindae Q. uniformis Neo. tatrica Ma. baloghi Q. auriformis Carn. nodosa Q.noah Journal Pre-proof Julian 1-2 Boundary J-TB J1-J2B C-NB J-TB J1-J2B ur Figure 5. Sequence of Unitary Association Zones from statistical analysis of biostratigraphic Jo compilations in different regions, including the UAZ results of (A) all global sections, (B) only the South China sections, and (C) the suite of Western Tethys, southern Tethys and British Columbia sections. Black squares represent discrete (discontinuous) zones, and grey squares represent predicted zones (continuous) defined by the occurrence and/or pairs of characteristic species. For abbreviations of genera, see Fig. 2. To evaluate the uncertainty in applying these UAZs, we ran two tests. In the first test, we computed two regional UAZs by dividing the global database into (1) only the South China sections (Fig. 5B), and (2) only the western and southern Tethys and British Columbia sections (Fig. 5C). There is a perfect match of the sequence of FOs of the zonal conodonts in the global and in both regional versions. Accordingly, we highlight the UAZ results in some Journal Pre-proof key reference sections, including HWG, HWQ, YY and LC sections (eastern Tethys), Aghia Marina section (western Tethys), and Wadi Mayhah section (southern Tethys). All of these key reference sections yield an onset of the CPE excursion within the UAZ-4 and UAZ-8 (Table 1), with the peak in the negative δ13Ccarb excursion in the UAZ-9, and the termination of the CPE excursion around UAZ-10 (Fig. 5). In the second test, the conodont assemblages were examined only in the CPE onset of intervals (2 to 4 m) (Table 1). All the sections yield UAZ-4 to UAZ-8 (Julian2, ~234.5 to 233.5 Ma) as the onset interval for the CPE excursions, which indicates there is no -p ro appreciable discernable offset of the CPE excursion across the Tethys realm. Jo ur na lP re Table 1. Age model calibration points for reference sections based on the pairs of conodont markers or other recognition criteria for the bases of Julian1, Julian2, Tuvalian, and Norian. Notes: ‘<’ or ‘>’ imply that the estimated age should be younger or older than the indicated age-control point; ‘<=’ implies that the estimated age should be younger or equal to the correspond age tied-point; and ‘?’ indicates that the estimated age control was based only on Journal Pre-proof the top of the indicated lithostratigraphic formation. Abbreviations in conodont taxa: 4 = P. inclinata, 8 = P. foliata, 9 = P. tadpole, 12 = Q. polygnathiformis, 17 = Ma. carnica, 20 = Para. praelindae, 23 = Ma. baloghi, 26 = Q. noah. Other abbreviations of genera are given in Fig. 2. 4.3. CPE carbon isotope excursions based on new data from South China The δ13Ccarb values remain relatively stable during UAZ-1 to UAZ-3 (~237 to 234.5 Ma) of in carbonate strata of the studied sections in South China (ca. 3 ± 1‰; n = 62) (Fig. 6). This ro was followed by a single major negative excursion of -3.0‰ (n = 78) in the interval -p equivalent to UAZ-4 to UAZ-9 (~234.5 to 232.5 Ma) in the lower part of the siliciclastic-dominated strata that likely marks the onset of the CPE. Following this excursion, lP the studied intervals (~232.5 to 231 Ma). re δ13Ccarb recovered to relatively higher values (1 ± 1 ‰; n = 44) in UAZ-10 through the top of For organic carbon the δ13Corg values are also stable at around -28 ± 2 ‰ in the UAZ-1 na to UAZ-3 interval. Interestingly the δ13Corg during the CPE in Sichuan sections displays a ur gradual and monotonous positive trend; whereas, we note that the negative excursion in Jo coupled carbon isotopes is clearly recorded in some of the Guizhou sections (e.g. Longchang section; Sun et al., 2016). Considering the pronounced lithologic change of the increased terrigenous clastic influx that marks the CPE in these sections, it is possible that terrestrial organic carbon sources may seriously contaminate signatures of the original marine δ13Corg components thereby producing significantly dissimilar signals of δ13Corg in the different regional depositional settings. According to the lithologic, elemental and isotopic analytical results, we conclude that the Reingraben turnover in lithology and the CPE excursion in carbon isotopes were synchronous events, and that both of them are important features of the CPE. Jo ur na lP re -p ro of Journal Pre-proof Figure 6. Chronology inferred in the Sichuan basins using conodont biostratigraphic constraints on the carbon isotope stratigraphy. (A-B) Lithostratigraphic and coupled δ13Corg and δ13Ccarb stratigraphy for the HWQ and HWG sections. Conodont biostratigraphy is modified after Yuan (2016) and Jiang et al. (2019). Shadings are to highlight the same substage or sub-substage among the sections. (C-D) Cross-plots of δ13Ccarb to δ18O. Abbreviations: Q. = HWQ; G. = HWG; CIE = carbon isotope excursion. For lithologic legend, see Fig. 4. Journal Pre-proof 5. Discussion 5.1. Bio- and magneto-stratigraphic age model In order to resolve common trends in the stable-isotope analyses from the 13 global reference sections, the values with depth in each section need to be transformed to a standard time scale by adjusting for variable sediment accumulation rates. This is particularly of important for the CPE interval, during which the majority of the sections underwent major changes in the relative proportions of carbonate and siliciclastic components. ro The composite UAZ provides partial guidance, but that scale lacks an age model. -p Therefore, we applied four age tie points based on U/Pb zircon geochronology, re biostratigraphy, and magnetostratigraphy for a common age model: (1) The Ladinian-Carnian lP boundary (LCB) was assigned an estimated age of ~237 Ma based on U/Pb zircon dating from the Dolomites (Mietto et al., 2012). The LCB is marked by the FO of conodonts Q. na polygnathiformis or Q. intermedius; (2) The Julian1-2 boundary has an estimated age of ~234.5 Ma based on magneto- and bio-stratigraphic profiles with partial astronomical tuning ur from Zhang et al. (2015, 2020). The Julian1-2 boundary is located in the middle part of the Jo Ma. carnica Zone. In addition, the Julian1-2 boundary corresponds to the onset of the lithologic change and the onset of the major δ13Ccarb excursion; (3) The Julian-Tuvalian boundary (JTB) was assigned an estimated age of ~233.5 Ma (Zhang et al., 2015, 2020), and is marked by the FO of Paragondolella noah or the FO of Quadralella tuvalica; and (4) The Carnian-Norian boundary was assigned as ~227.3 Ma (Ogg and Chen, 2020) based on the conodont marker (FO of Metapolygnathus parvus or of Primatella primitus, Orchard, 2014) with its magnetostratigraphic placement in the Italian GSSP candidate at Pizzo Mondello (Mazza et al., 2018) and correlation to the astronomical-tuned geomagnetic time scale from the Newark Basin. Note that the upper part of the Carnian (Tuvalian substage) in the study Journal Pre-proof sections, especially for South China, generally lacks systematic and detailed biostratigraphic constraints; therefore, the sediment accumulation rates from JTB to the terminal Carnian was commonly assumed to the same as the rates from the Julian1-2 boundary to the JTB. Where possible, the sediment accumulation rates based on these four age-control points were computed for each section (Table. 1) and used to convert the δ13Ccarb data to a common age scale. of For those sections that lacked the reporting of one or more primary conodont zones, then secondary conodont zones were used to constrain the ages (Table 1), including: (1) where the ro Mazzaella carnica Zone was not reported (AM, WM, WY, LSK sections), then FO of Para. -p praelindae was used as the base of Julian2; and (2) for sections that did not have a report of re the Meta. parvus and the Prim. primitus zones (AM, WM, FT sections), then the lowermost lP Norian was recognized by the FOs of Primatella gulloae, Ancyrogondolella quadrata, or Norigondolella nacicula. For the four geochemical reference sections that did not have na conodont zonations (FD, LA, LP, GL sections), the termination of formations that had been dated elsewhere as corresponding to the onset of the CPE was used (tops of the St. Cassian ur Formation for FD and LA, of the Reifing Formation for LP, and of the Chomule Formation Jo for GL); and only the Julian interval (prior to the base of CPE) of the δ13Ccarb data for these four sections was used. 5.2. Global CPE δ13Ccarb correlations There are distinct differences in biostratigraphy and carbon-isotopic signatures between the synthesis of the western sections and the suite of South China sections. The western Tethys, southern Tethys and British Columbia sections provide a more detailed integrated conodont stratigraphic zonation for the upper Carnian (Fig. 5C); and, importantly, the Julian-Tuvalian and the Carnian-Norian boundaries are assigned in this region to the first Journal Pre-proof occurrences of conodonts Q. tuvalica and Meta. parvus, respectively (Orchard, 2014; Mazza et al., 2018). In contrast, the South China sections from Sichuan and Guizhou provide a more detailed chronostratigraphy for the lower Carnian, especially for the CPE interval (Fig. 5B). However, there are slight offsets of the CPE excursion among different regions. For example, the Wadi Mayhah section provides the most complete δ13Ccarb profiles in southern Tethys (Oman, Sun et al., 2019), in which the onset of the CPE excursion occurs in the lower-middle of part of the Ma. carnica Zone. In contrast, the base of the CPE excursion in South China is located at the middle part of the Para. praelindae Zone (Sun et al., 2016; Jiang et al., 2019), ro implying a delayed CPE excursion in South China relative to southern Tethys. However, we -p cannot rule out the possibility that this apparent offset is caused by the incomplete conodont lP re biostratigraphy and/or taxonomic errors in taxa identification. 5.3. Adjustment for potential diagenetic overprints on the δ13Ccarb and δ13Corg data na Considering that an enhanced siliciclastic input is characteristic of the CPE, all clay-rich samples (Al >8 %) were removed from the suites of carbon isotopes values. The Sichuan ur sections record relatively high Mn/Sr and high Al/Ti ratios above the lithologic change, and Jo these are suggestive of meteoric alteration and indicate enhanced detrital contamination, respectively. To further investigate potential relationships of the isotopic shifts to the lithological changes during CPE interval, the HWG and HWQ sections were compared with the Longchang section from Guizhou (Sun et al., 2016). These two regions in South China were suggested to have different triggering mechanisms of the lithological change. When the Ca content of the Sichuan sections decreases to extremely low values (from 33 %, n = 53, to 7 %, n = 42), the Ca content stays at relatively high values in the Guizhou section (average = 27 %). The δ13Corg values of these regions show totally different trends that are mainly controlled by the lithologic changes. However, in contrast, the δ13Ccarb curves have similar Journal Pre-proof characteristics in the two regions and coincide well with the δ13Ccarb curves of Western and Southern Tethys sections. This indicates that some of the chemostratigraphic shifts (e.g. δ13Corg, TOC, and elemental concentrations) could be due to the lithologic change, but this had smaller effect on δ13Ccarb profiles. We should note that the δ13Ccarb values following the lithologic transitions in the Sichuan sections are relatively low. These lower values in the bulk carbonate may be partly caused by of post-burial diagenetic overprints. The post-CPE lithology has high TOC contents, but extremely low Ca content (7%) in Sichuan sections. In some cases, the post-depositional ro oxidation of organic matter releases light carbon into the interstitial waters, which then shifts -p the composition of the forming carbonate cement to lower δ13Ccarb values (Swart, 2015). re However, this scenario of oxidation of organic matter would not apply to many of the lP negative CPE excursion records in the Tethys realm, because the carbonate Ca content remains elevated (> 20 %) in other key sections (e.g., LC and WM sections), which also have na lower TOC but still display a large δ13Ccarb excursion (-2 ± 0.5 ‰). Nevertheless, there are two intervals in Sichuan (HWQ and HWG) in which the δ13Ccarb ur data may have been significantly influenced by diagenetic effects according to comparison of Jo trends in δ13Ccarb and δ18Ocarb and Mn/Sr ratios: (1) A positive correlation was present from 11.6 to 16.7 m in the HWQ section (R2 = 0.46, p-value <0.01, n = 12) above the lithologic change. The Mn/Sr ratios of this interval are relatively high (average = 4.4, n = 5). (2) A positive correlation from 27.8 to 34.4 m in the HWG section (R2 = 0.93, p-value <0.001, n = 13) occurs before the lithologic change. These intervals may have been influenced by meteoric and burial fluids; therefore, the isotopic data from these intervals were excluded from the later comparative compilations. Journal Pre-proof 2 UAZ 11-Top 10 UAZ 11-Top -24 UAZ 1-7 -4 -6 -25 -26 -27 G. Q. -28 -8 -10 UAZ 11-Top 0 1 2 3 δ Ccarb (‰) -30 4 UAZ 8-10 UAZ 1-7 6 4 2 UAZ 1-7 A -1 8 UAZ 8-10 -29 -2 Mn/Sr(ppm/ppm) UAZ 8-10 -2 G. Q. -23 δ13Corg (‰) δ18Ocarb (‰) -22 G. Q. 0 -2 -1 B 0 13 1 2 3 δ Ccarb (‰) 0 4 C -2 -1 0 13 2 25 1 2 δ13Ccarb (‰) 3 4 G. Q. UAZ 11-Top 15 10 G. Q. UAZ 1-7 1 0.5 UAZ 11-Top 5 UAZ 8-10 1.5 UAZ 8-10 -2 -1 D 0 1 2 δ Ccarb (‰) 3 4 13 0 -30 -29 -28 -27 ro UAZ 1-7 0 of TOC (wt.%) Al/Ti (%/%) 20 -26 -25 δ13Corg (‰) E -24 -23 -22 lP re -p Figure 7. Cross-plots of δ13Ccarb and of δ13Corg to Mn/Sr ratio, [Sr], [Mn], Al/Ti ratio, TOC content and δ18Ocarb of the carbon-isotope reference sections of the South China Block (Sichuan). The arrows show the trend of signatures of isotopic or elemental ratio changes from bottom to top for these sections. Cross-plots of δ13Ccarb and δ18Ocarb clearly show that the South China sections generally na have much lower δ18Ocarb values (-2 ‰ to -8 ‰) than the western and southern Tethys ur sections (+2 ‰ to -4 ‰); and they also record relatively steeper slopes of δ18O vs. δ13C (Fig. 8). The Carnian Pluvial Episode may have increased the influx of fresh water into Jo semi-enclosed seas (Hornung et al., 2007a), thereby producing different water-mass salinity conditions; which in turn would lead to variable δ18Ocarb values in the precipitated carbonates (Wei and Algeo 2019). In addition the lower δ18Ocarb values suggest that the South China sections might have been previously subjected to deeper burial diagenetic processes prior to later uplift and/or had undergone significant diagenetic influence by meteoric water. To quantify the CPE excursions worldwide the δ13Ccarb data were normalized using a 5/1 slope burial trend (Hasiuk et al., 2016) (Fig. 8) to eliminate the secondary fractionation of δ13Ccarb during burial processes: (1) Journal Pre-proof A negative δ13Ccarb excursion with magnitude of -2 ± 0.5 ‰ (from an average of ~3 ‰ to ~1 ‰) is evident in these normalized trends thereby representing a moderate CPE excursion in the composite global scale (Fig. 9). Considering the incomplete biostratigraphic study and the limited accuracy of the UAZs, we propose a synchronous CPE excursion age model (Fig. 9C) based on the revised age-tie points (Table 1). This yields an onset of ~234 Ma (Julian2) and a ~1.5 Myr duration for the CPE excursion and sheds light on this global δ 13Ccarb HWQ Longchang Yongyue Wadi Mayhah S1-101 A Primary Marine Signature na 0 Milieres-Dibona South China sections Other sections lP 2 re -p ro HWG of negative excursion as an important feature of the CPE. -2 Pizzo Mondello Aghia Marina Fedares & Lavarella Freygutweg Guling2 Silicka Brezova -6 δ18Ocarb ‰ -4 Jo ur Marine-Meteoric Mixing zone Diagenitic Trend (Meteoric) -8 Burial-Meteoric Mixing zone -10 Burial Trend -12 10 8 6 4 2 0 -2 δ Ccarb ‰ 13 -4 -6 -8 -10 Journal Pre-proof Figure 8. Compilation of cross-plots of δ13Ccarb and δ18Ocarb from the array of global reference sections. (A): Diagenetic trends representing fluid evolution can be elucidated by cross-plots of stable-isotope data from carbonate samples. The trends of the cross-plots of δ13Ccarb and δ18Ocarb plot along widely cited diagenetic trends from the primary marine-carbonate signatures for burial, meteoric, marine-meteoric mixing zone, and burial-meteoric mixing zone processes. 5.4. Wrangellia LIP injection simulation using the C-P cycle model of The CPE excursion occurred after a long-term secular positive δ13Ccarb trend that began ro during the early Anisian (Payne et al., 2004) that culminated in relatively high background δ13Ccarb values (~3 ‰ n = 450). The organic-carbon components in the Julian1 strata yield -p δ13Corg values of ca. -27‰ (n = 167) and a Δδ13C of 30‰, and these values are employed as re the steady-state baseline in the mass balance model. There are significant differences of the lP secular δ13Corg shifts among various sections in western Tethys (e.g., Dal Corso et al., 2018) and in eastern Tethys (Fig. 9D). The trends of δ13Ccarb and δ13Corg can be both coupled (e.g., na the Longchang section of Sun et al., 2016; and the Lunz section of Hornung et al., 2007a and ur Mueller et al., 2016a) or decoupled (e.g., Sichuan sections in this study), with the δ13Corg trends being usually more equivocal than those of δ13Ccarb (Fig. 9). Indeed, there are Jo pronounced positive δ13Corg excursions within the CPE in some sections, which implies that δ13Corg values measured from CPE strata are not only controlled by the δ 13C of the surface ocean dissolved inorganic carbon (DIC), but also involved other sources of organic carbon and/or carbon isotope fractionation. lP re -p ro of Journal Pre-proof Jo ur na Figure 9. Comparison of Carnian δ13Ccarb profiles throughout the Tethys realm and δ13Corg variation among the studied sections. (A-B) The raw δ13Ccarb data from South China sections and other Tethyan sections with comparison to burial-normalized LOWESS fits; (C) The δ13Ccarb data normalized for burial diagenetic trends with LOWESS curves; (D) δ13Corg data from some key sections around the Tethys realm. Conodont UAZ scale (see Fig. 5 for detailed Unitary Association Zones) and geomagnetic polarity (modified after Zhang et al., 2020, and Ogg and Chen, 2020). The distinct increasing trend of TOC values in some sections after the CPE lithologic change (e.g., Sichuan sections in Fig. 4) raises the possibility of facies-dependent temporal and regional δ13Corg variations. The positive excursion of δ13Corg may have been locally attenuated by terrestrial-derived detrital organic carbon, which has relatively low (lighter) δ13Corg signatures. Terrestrial-derived organic material with lighter δ13Corg, especially from wood fragments, probably influenced the δ13Corg trends from the Sichuan Basin (Jin et al., 2018) and from the Heiligkreuz section (Dal Corso et al., 2018). Note that there is no Journal Pre-proof irreconcilable contradiction of the trends of the δ13Corg that include a secular positive shift recorded in this study and the multiple sharp negative shifts observed by Dal Corso et al. (2018). These two features in δ13Corg are superimposed, but the consistent trend in δ13Ccarb suggests that marine carbon-isotope conditions oscillated at the global scale during the CPE interval. The eruption of the Wrangellia Large Igneous Province (W-LIP) has been proposed to be of a trigger for the CPE excursion. Therefore, to evaluate the influence of the possible extra input of LIP-derived light carbon into the oceanic carbon cycle during the CPE interval, we ⁄ ro applied Equation (2) to the burial-trend normalized δ13Ccarb data: -p [ ] (2) re Equation (2) is a modified version of the representation of the oceanic carbon cycle, lP based on mass balance equations for the seawater carbon inventory and its isotopic compositions (Kump and Arthur, 1999; Table 2). A = 3.8 x 10 18 mol ‰ Jo δA Units Reference Size of the modern DIC pool (Kump and Arthur, 1999). ur Abbreviation na Table 2. Abbreviations, units, and references for Equation (2). Isotopic composition of ocean-atmosphere inventory mol/yr Flux of tectonically recycled carbon to the biosphere (Kump and Arthur, 1999). δin = -5 ‰ Isotopic composition of recycled carbon (Kump and Arthur, 1999). δLIP = -5 ‰ Isotopic composition of LIP injected carbon (Kump and Arthur, 1999). δorg = -27 ‰ δ Corg isotopic composition baseline (Julian1) for the reviewed sections. △ = 30 ‰ Carbon isotope fractionation when buried as organic matter. Calculated by the observed data. ƒorg = 0.267 - Forg = F in x ƒorg mol/yr Fin = 5 x 10 13 13 Fraction of carbon buried as organic matter. Value of 0.267 13 gives δ C of DIC of ~3.0 ‰ at steady state, similar to values observed at the start of the excursion. δ 13Corg isotopic composition baseline (Julian1) for the reviewed sections. Journal Pre-proof Throughout the Tethys realm, one major peak of extra light-carbon input (FLIP) occurred during the CPE interval (Fig. 10A). The amount of extra carbon flux into ocean-atmosphere system (given by the integral of FLIP) was calculated by setting the Fin to ~0.5 × 1018 mol/yr, which is 10% of its initial value (Table 2; Kump and Arthur, 1999). For the CPE excursion pair, this calculation yields a total of ~1.4 × 1018 and ~1.8 × 1018 mol of light carbon input to produce these FLIP peaks in South China and in western Tethys, respectively. These amounts of of light carbon would require ca. 300% and 400% of the estimated total light carbon output of Jo ur na lP re -p ro the Wrangellia LIP (Dal Corso et al., 2012; Sun et al., 2019). Journal Pre-proof Figure 10. (A) Response of oceanic carbon input (Fin) and the calculated LIP carbon input (FLIP) required to produce the normalized negative δ13Ccarb excursions during the CPE interval. (B-F) Modeling results of C-P cycle model for cumulative LIP carbon releases of 0.5, 1, 2, 4 × 1018 mol C. (G) Evolutionary trends of selected major Mesozoic biota during the Carnian. See main text for details. Data sources: (1-4) Trotter et al. (2015), Sun et al. (2016), Dal Corso et al. (2012), and Bernardi et al. (2018). These modeling results are consistent with the conclusion of earlier modeling efforts in that the amount of 13 C-depleted carbon released only by the Wrangellia LIP eruption is not of sufficient, by itself, to produce the magnitude of the negative CPE excursion recorded in the additional 13 ro western Tethys (Dal Corso et al., 2012, 2018; Sun et al., 2019). Potential sources of C-depleted carbon to contribute to the negative CPE excursion include release -p from organic-rich sedimentary rocks, oceanic warming that destabilized deep-sea clathrates, re and eruption of other contemporaneous volcanism (Sun et al., 2019) such as the widespread lP magmatic rocks on the Tibetan Plateau region that have yielded dates of ca. 235-190 Ma (Li et al., 2019). na These results from the oceanic carbon cycle model were introduced as disturbances to ur the steady-state conditions of the carbon-phosphorus (C-P) cycle model. The significant predicted environment changes by the model are temperature, atmospheric CO2, oceanic Jo anoxia, and atmospheric oxygen (Clarkson et al., 2018) (Fig. 10B-F). Hence, we simulated the climatic impacts of cumulative LIP carbon releases with a δ13C composition of -5 ‰ in four cases (0.5, 1, 2 and 4 × 1018 molC) starting at 234 Ma and lasting 1 Myr (Fig. 10B). These simulations yielded a range of 0.2 to 1.8 °C temperature increases peaking at ~234.4 Ma (Fig. 10C), which was the peak of the Julian2 climate warming according to published δ18Oconodont studies (e.g. Sun et al., 2016). In these simulations, the carbon releases from LIPs generate a peak CO2 of ~1.05 to 1.5 times PAL, thereby increasing the rate of silicate weathering and associated riverine fluxes of Phosphorus to the ocean that generates moderate deep-sea anoxia (Fig. 10D, 10E). Of course, the degree of oceanic anoxia will be more Journal Pre-proof serious if we concentrated the total LIP carbon release into a shorter time span (e.g. spanning only 0.1 Myr) or increased that total flux. If this oceanic anoxia and oxygen-minimum zone rose in depth to impact the continental shelves, then it may have played a role in the extinction of faunas during the Julian-Tuvalian transition (e.g. turnover in ammonoids and conodonts, Simms and Ruffell et al., 1989; Chen et al., 2016). In contrast, the simulation implies that the resulting enhanced carbon buried in organic form in the anoxic seas would of eventually result in a small net increase in atmosphere O2 in the range of 1.01 to 1.06 PAL during the Tuvalian (Fig. 10E). These simulations reveal interesting potential feedbacks ro between the proposed LIP carbon releases and the evolution of Carnian ecosystems, re -p including the development of dinosaurs (Fig. 10G, e.g., Bernardi et al., 2018). lP 6. Conclusions The Carnian Pluvial Episode (CPE) was an interval of increased global temperatures and na precipitation. A worldwide database of 13 conodont-dated sections with a total of 649 δ13Ccarb and 168 δ13Corg data points enables resolution of the relative timing and modeling of the ur magnitudes of the pronounced negative carbon-isotope excursions during the CPE. The Jo statistical analysis yielded a high-resolution conodont Unitary Assemblage Zonation of 17 UAZs spanning the Carnian. This synthesis indicates that the magnitude of the global carbon-isotope excursion recorded in carbonates was -2 ± 0.5 ‰ (from 3 to 1 ‰). The age model for the UAZs and magnetostratigraphy constraints indicate that the CPE excursion had its onset during UAZ-4 to 8 (~234.5 Ma, Julian2), spanned ~1.0 to 1.5 Myr throughout the Tethys realms, and confirms that it was contemporaneous with the Reingraben turnover in lithologic facies. Applying a combination of an oceanic carbon cycle model and a global carbon and phosphorus model (C-P model) supports a strong relationship between the global CPE excursion and the eruption of the Wrangellia LIP. However, the emissions from the Journal Pre-proof Wrangellia LIP alone were inadequate to account for the required total amount of light carbon required for the magnitude (-2 ± 0.5 ‰) of the global excursion. The C-P model predicts a global temperature rise, a moderate expansion of oceanic anoxia, and a secondary delayed increase in atmospheric oxygen levels; and these feedbacks would have implications for the turnover and evolution of Carnian marine and terrestrial ecosystems. of Acknowledgements Shijie Liu is thanked for assistance in the field and measuring samples in laboratory. We ro are grateful to Thomas Algeo, Jitao Chen and another anonymous reviewer for their critical -p comments and constructive suggestions, which have improved greatly the quality of the paper. re Ziheng Li is funded by the Fundamental Research Funds for National Universities, China lP University of Geosciences（Wuhan). This study is supported by five grants from NSFC (41977264, 41930322, 41673011, 41821001, 41830323), by the 111 project of China na (BP0820004) and by the Fundamental Research Funds for the Central Universities, China University of Geosciences-Wuhan (No. CUGQYZX1728). 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