carbon-isotope stratigraphy across the permian–triassic boundary: a review

Journal of Asian Earth Sciences 39 (2010) 215–235

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Journal of Asian Earth Sciences

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Carbon-isotope stratigraphy across the Permian–Triassic boundary: A review

Christoph Korte a,*, Heinz W. Kozur b

a Institut für Geologische Wissenschaften, Freie Universität Berlin, Malteserstr. 74-100, 12249 Berlin, Germanyb Rézsü u. 83, H-1029 Budapest, Hungary

a r t i c l e i n f o a b s t r a c t

Article history:Received 8 July 2008Received in revised form 15 January 2010Accepted 19 January 2010

Keywords:Permian–Triassic boundaryCarbon-isotopesChemostratigraphyTrap volcanismOcean anoxia

1367-9120/$ - see front matter � 2010 Elsevier Ltd. Adoi:10.1016/j.jseaes.2010.01.005

* Corresponding author.E-mail address: [email protected] (C. Korte).

The Palaeozoic–Mesozoic transition is marked by distinct perturbations in the global carbon cycle result-ing in a prominent negative carbon-isotope excursion at the Permian–Triassic (P–T) boundary, wellknown from a plethora of marine and continental sediments. Potential causes for this negative d13C trend(and their links to the latest Permian mass extinction) have been intensively debated in the literature. Inorder to draw conclusions regarding causation, a general d13C curve was defined after consideration of allavailable datasets and with due reference to the biostratigraphic background. The most important fea-tures of the P–T carbon-isotope trend are the following: the 4–7‰ d13C decline (lasting �500,000 years)is gradual and began in the Changhsingian at the stratigraphic level of the C. bachmanni Zone. Thedecreasing trend is interrupted by a short-term positive event that starts at about the latest Permianlow-latitude marine main extinction event horizon (=EH), indicating that the extinction itself cannot havecaused the negative carbon-isotope excursion. After this short-term positive excursion, the d13C declinecontinues to a first minimum at about the P–T boundary. A subsequent slight increase is followed by asecond (occasionally two-peaked) minimum in the lower (and middle) I. isarcica Zone. The negative car-bon-isotope excursion was most likely a consequence of a combination of different causes that mayinclude: (1) direct and indirect effects of the Siberian Trap and contemporaneous volcanism and (2)anoxic deep waters occasionally reaching very shallow sea levels. A sudden release of isotopically lightmethane from oceanic sediment piles or permafrost soils as a source for the negative carbon-isotopetrend is questionable at least for the time span a little below the EH and somewhat above the P–Tboundary.

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1. Introduction

After one of the most remarkable turning points in Earth’s his-tory, namely the latest Palaeozoic mass extinction, marine and ter-restrial life remained strongly disturbed prior to full recovery, withan unusually long delay of more than five million years (e.g., Erwin,1993, 2006; Retallack, 1995; Eshet et al., 1995; Bowring et al.,1998; Kozur, 1998a,b; Rampino and Adler, 1998; Jin et al., 2000;Benton, 2003; Benton et al., 2004; Peng and Shi, 2009). The as-sumed cause(s) for this biotic crisis are still under discussion andinclude: (1) large-scale volcanic activity of Siberian flood basalt(e.g., Renne and Basu, 1991; Campbell et al., 1992; Conaghanet al., 1994; Renne et al., 1995; Kozur, 1998a,b; Reichow et al.,2002, 2009; Kamo et al., 2003; Visscher et al., 2004; Courtillotand Olson, 2007; Isozaki, 2007; Ganino and Arndt, 2009; Svensenet al., 2009) and contemporaneous volcanism in South China (Yinet al., 1992; Kozur, 1998a,b); (2) ocean anoxia reaching unusuallyshallow depths (Wignall and Hallam, 1992; Isozaki, 1994, 1997;

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Kajiwara et al., 1994; Wignall and Twitchett, 1996; Wignall et al.,1998; Kato et al., 2002; Kidder and Worsley, 2004; Nielsen andShen, 2004; Grice et al., 2005; Kump et al., 2005; Riccardi et al.,2006, 2007; Hays et al., 2007; Xie et al., 2007a; Algeo et al.,2008; Grasby and Beauchamp, 2009), with degassing of CO2 (Knollet al., 1996; Woods et al., 1999), H2S (Kump et al., 2005; Kaihoet al., 2006; Riccardi et al., 2007; Meyer and Kump, 2008; Meyeret al., 2008), methane (Heydari and Hassanzadeh, 2003; Retallacket al., 2003), or a combination of these (Ryskin, 2003); (3) an oce-anic acidification crisis due to increase in atmospheric CO2 concen-trations (Heydari et al., 2003; Fraiser and Bottjer, 2007; Payneet al., 2007); (4) low atmospheric oxygen levels (Berner, 2002,2005; Retallack et al., 2003; Huey and Ward, 2005; Berner et al.,2007); (5) worldwide depletion of stratospheric ozone (e.g., Kozur,1998a,b; Visscher et al., 2004; Kump et al., 2005; Sephton et al.,2005; Beerling et al., 2007); and (6) climate change caused bystrong volcanism (volcanic winter), impact of a celestial body(but see Koeberl et al., 2004), or both (Stanley, 1988; Campbellet al., 1992; Erwin, 1993; Kozur, 1998a,b; Retallack et al., 1998;Jin et al., 2000; Mory et al., 2000; Kaiho et al., 2001; Becker et al.,2001, 2004). Reviews dealing with the latest Permian extinction

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216 C. Korte, H.W. Kozur / Journal of Asian Earth Sciences 39 (2010) 215–235

were recently published by Erwin et al. (2002), Racki and Wignall(2005), Isozaki (2007, 2009a,b), Knoll et al. (2007), Twitchett(2007a,b), Wignall (2007), Bottjer et al. (2008), Krassilov and Kara-sev (2009) and Posenato (2009).

The mass extinction and environmental changes close to thePermian–Triassic (P–T) boundary were accompanied by major per-turbations in the global carbon cycle, as marked by a pronouncednegative carbon-isotope excursion. Unusually high 13C-enrich-ments in Permian carbonates were reported in the 1960s and1970s (e.g., Compston, 1960; Osaki, 1973), marking the highestvalues of the Phanerozoic (Veizer et al., 1980). Wilgus (1981) pro-posed a global negative carbon-isotope trend across the P–T tran-sition from her own data (western USA) and from compilation ofliterature (Compston, 1960; Osaki, 1973; Galimov et al., 1975;Kirkland and Evans, 1976; Veizer and Hoefs, 1976; Magaritz andSchulze, 1980; Veizer et al., 1980), and this was later confirmedby Clemmensen et al. (1985), but only from a few data points inthe E-Greenland successions.

Chen et al. (1984) presented the first detailed d13C trend for theP–T boundary section at Meishan (South China), the Global Strato-type Section and Point (GSSP) of the Permian–Triassic boundary,showing a decline in excess of 6‰ in the latest Permian, reachingthe lowest values at �3‰. Subsequent higher resolution carbon-isotope datasets, published by Xu et al. (1986) for Meishan, Shangsi(both South China) and Sovetashen (Armenia) and by Holser andMagaritz (1987) for the Kuh-e-Ali Bashi section near Jolfa (NWIran), proposed that the negative peak is situated close to the P–Tboundary. Further results from a shallow-marine section that hashigher sedimentation rates than Meishan, the Tesero in SouthernAlps (Italy), subsequently indicated that the negative carbon-iso-tope shift is gradual (Magaritz et al., 1988). Still later, Baud et al.(1989) showed that this d13C excursion is present in several classicP–T boundary successions of the Tethys Realm as well as in themarginal seas, such as the sections at Jolfa (Holser and Magaritz,1987), Vedi (Armenia), Sovetashen (Armenia), Emarat (N-Iran),Idrijca (W-Slovenia), Çürük Dag (SW-Turkey), Kemer Gorge (Antal-ya), Nammal Gorge (Salt Range, Pakistan), Thongde (Zanskar,Himalaya, India), Shangsi (S-China), Meishan (S-China) and KakiVigla (Greece). Its amplitude was generally in the 4–7‰ range(e.g., +4 to 0‰ for Çürük Dag; +4 to �3‰ for Shangsi). This featureargues strongly that the P–T boundary negative shift is global innature. As the same time, Oberhänsli et al. (1989), while reportinga similar gradual d13C decline for the San Antonio section of South-ern Alps, also published the data for the Schuchert Dal section ofJameson Land at East Greenland that shows Late Permian valuesof about �3‰ followed by sudden two-peak negative shifts in ex-cess of 20‰ near the P–T boundary, a feature contrasting with allinvestigated marine low- and mid-latitude boundary sections.

The pilot study of Holser and Magaritz (1985) for the Reppwandsection of Carnic Alps (Austria) that did not show any discerniblechanges in carbonate d13C at the P–T boundary was followed bythe Holser et al. (1989) study of the adjacent Gartnerkofel core(see also Magaritz and Holser, 1991) that presents one of the mostimpressive carbon-isotope curves for the P–T boundary interval.The Gartnerkofel succession was deposited in a shallow-marineenvironment of the western Tethys more than 150 km away fromthe shoreline (Buggisch, 1974, 1978; Holser et al., 1989). Its mod-erately high sedimentation rates enabled generation of a high-res-olution carbon-isotope record that shows a gradual negative trend,similar to that in Tesero (Magaritz et al., 1988), with a minimumvery close to the biostratigraphic P–T boundary. A second two-peak minimum is reported for the I. isarcica Zone. The Holseret al. (1989) paper was followed by a plethora of publications thatreport both the carbonate and the organic d13C trends for the aboveand additional successions. These were reviewed for example inScholle (1995) and Corsetti et al. (2005), and they are listed in

Fig. 1 and Table 1. The reported amplitudes, shapes, durationsand particularly the assigned exact stratigraphic positions of thenegative d13C peaks appear to have been somewhat variable (e.g.,Twitchett, 2007a; Metcalfe et al., 2008), but nevertheless recordedin a wide range of marine and continental deposits. If global, due toshort residence time of carbon in the ocean, the peaks and troughsof d13C excursions should represent coeval time markers and en-able trans-continental stratigraphic correlations.

Here we review details of several P–T boundary carbon-isotopetrends in the context to their biostratigraphic backgrounds, withupdated biostratigraphy that may differ somewhat from that pre-sented in the original publications. This review may then enableconstruction of a general P–T boundary d13Ccarb curve, and recogni-tion of geological factors that may impacted the ocean/atmospheresystem during the latest Permian.

2. Biostratigraphic background and numerical ages

In order to follow the discussion, in particular by readers notfamiliar with the details of P–T biostratigraphy, we will briefly ex-plain the highly detailed Late Permian to earliest Triassic conodontzonation (Fig. 2). We use the international stages Wuchiapingianand Changhsingian, defined at the Chinese stratotype sections, inaddition to regional stage names of Dzhulfian (=Wuchiapingian)and Dorashamian (=Changhsingian) for the Tethyan successions.

The conodont zonation for Late Permian earliest Triassic ofSouth China was developed by Wang and Wang (1981a,b) and laterrefined by Wang (1996), Mei et al. (1998), Jiang et al. (2007) andothers and recently compiled (Fig. 2) by Metcalfe and Isozaki(2009) based on Jin et al. (1997, 2006) and Zhao et al. (2007,2008). Where necessary, we also employ the conodont zonationof Iran and Transcaucasia (Fig. 2) because it shows the highest res-olution of all conodont zonations during this time span (Kozur(2005, 2007); see also Metcalfe and Isozaki (2009)). The lower res-olution of the South Chinese zonation results mainly in the LatePermian–Early Triassic intraplatform nature of this region thatwas generally prone to endemism indicated by the absence ofwidespread open marine Tethyan ammonoid guide forms, suchas Paratirolites that occur in Central and NW Iran, Transcaucasia,Southern Alps and Madagascar. South Chinese conodonts are onlyslightly affected by endemism, but the Tethyan guide forms C. ira-nica (Kozur) – dominating the C. iranica Zone in Transcaucasia,western Sicily, central and NW Iran – is absent at Meishan, andC. nodosa (Kozur) – dominating the C. nodosa Zone in central Iran,NW Iran and Transcaucasia – is absent in the entire intraplatformbasins of South China. Other Changhsingian open sea index species,however, are present in South China, but they were not yet used forthe South Chinese conodont zonation. In order to compare the car-bon-isotope trends of South China, Iran and Transcaucasia we haveinterpreted the Chinese conodont zonation by the review of litera-ture (e.g., Wang and Wang, 1981a,b; Wang, 1996; Jin et al., 1997,2006; Mei et al., 1998; Lai et al., 1999; Chen, 2007; Henderson,2007; Ji et al., 2007; Jiang et al., 2007; Metcalfe and Nicoll, 2007;Zhao et al., 2007, 2008; Chen et al., 2009), by our material and byconodonts of Prof. Wang Cheng-yuan (Nanjing). The following dis-coveries are most important (Fig. 2). C. hauschkei (Kozur) shows noprovincialism and occurs from Iran to Greenland. It was reportedfrom South China by B. Wardlaw (Reston, USA) in several sessionsof the Permian Subcommission, but he has not yet published his re-sults. C. bachmanni (Kozur) was recognised in South China in theM.Sc. thesis by Jun Chen (Chen, 2007) under Shuzhong Shen’sand C.M. Henderson’s supervision (Prof. C.M. Henderson, Calgary,pers. comm.). As C. bachmanni is the index species of the C. bach-manni Zone, this zone is obviously also present in South China.Metcalfe and Nicoll (2007) reported Merrillina ultima (Kozur) in

Fig. 1. Localities-map of the discussed P–T boundaries (red: marine; green: non-marine). Regions and/or countries (localities in parentheses): 1: Zhejiang, China (Meishan);2: Anhui, China (Chaohu); 3: Shaanxi, China (Wujiaping) and Hubei, China (Daxiakou); 4: Guizhou, China (Dajiang, Dawen, Gaoqao, Heping, Guizhou, Lekang, Shitouzhai,Zhongliangshang, Zhuzang, Chahe, Zhejue); 5: Guangxi, China (Dongpan, Taiping, Yalang, Zuodeng) and northern Vietnam (Lung Cam, Nhi Tao); 6: Sichuan, China (Huayun,Shangsi); 7: S-Tibet, China (Qubu, Selong, Tulong) and Central Nepal (Thini Shu); 8: Spiti, India (Losar); 9: Kashmir, India (Guryul Ravine, Palgham) and Salt Range, Pakistan(Landu Nala, Nammal Gorge); 10: NE–India (Banspetali, Damodar Valley, Mahanadi Valley, Son Valley, Warda coalfield); 11: Xinjiang, China (Jimsar: Dalongkou, Lucaogou);12: Oman (Buday’ah, Wadi Alwa, Wadi Maqam, Wadi Musjah, Wadi Sahtan, Wadi Sawat, Wadi Wasit); 13: Central Iran (Abadeh, Kuh-d-Yagma, Shahreza); 14: NE–Iran (Amol,Emarat); 15: NW Iran (Jolfa, Zal) and Armenia (Sovetachen, Vedi); 16: Turkey (Çekiç Dag, Çürük Dag, Demirtas�-Kus�davut, Karabayır Yayla, Kemer Gorge, Kopuk Dag, Tas�kent);17: Salamis, Greece (Kaki Vigla); 18: Sosio, Sicily, Italy (Palazzo Adriano); 19: Bükk Mts, Hungary (Bálvány, Gerennavár); 20: Southern Alps, Italy and Austria (Gartnerkofel,Peitlerkofel/Sas de Pütia/Sass de Putia, Pufels/Bula/Bulla, San Antonio, Seis/Siusi, Seres, Tesero, Uomo) and Slovenia (Brsnina, Idrijca, Košutnik Creek, Masore, Trzic); 21:Germanic Basin (Nelben, Thale); 22: Spitsbergen (Festningen, western Dicksonland); 23: Jameson Land, Greenland (Fiskegrav, Schuchert Dal); 24: British Columbia, Canada(Williston Lake); 25: NE–Siberia (Okhotsk); 26: NE–Siberia (Omolon); 27: Hyogo, Japan (Sasayama); 28: Kyushu, Japan (Kamura, Takachiho) and Shikoku, Japan (Taho,Tenjin’maru); 29: New Zealand (Wairoa Gorge); 30: Sydney Basin, New South Wales, Australia (Murrays Run, Wybung Head); 31: Bowen Basin, Australia (Denison,Eddystone, Newlands Coal Mine); 32: Bonaparte Basin, Australia (Fishburn, Petrel, Tern); 33: Canning Basin, Australia (Paradise); 34: Perth Basin, Australia (Hovea,Woodada); 35: Southern Madagascar (Morondava); 36: Karoo Basin, South Africa (Bethulie, Carlton Heights, Commando Drift Dam, Injusiti, Lootsberg, Wapadsburg); 37:Central Transantarctic Mts (Allan Hills, Shapeless Mountain, Mount Crean, Portal Mountain); 38: Central Transantarctic Mts (Coalsack Bluff, Graphite Peak); 39: Persian Gulf,Iran (South Pars Gas Field); 40: Sverdrup Basin, Canada (Borup Fiord Pass, Buchanan Lake, Confederation Point, Drake Point, East van Hauen Pass, Lindström Creek, South OttoFiord Mouth, West Blind Fiord); Chongqing, China (Liangfengya), is situated between 3, 4 and 6. (For interpretation of the references to colour in this figure legend, the readeris referred to the web version of this article.)

C. Korte, H.W. Kozur / Journal of Asian Earth Sciences 39 (2010) 215–235 217

the uppermost Changhsingian of South China. This species is pres-ent in all pelagic sections of Transcaucasia and central and NWIran, occasionally in co-occurrence with Stepanovites ? mostleri.The latter conodont co-occurs in topmost Permian shallow-watersediments in the Southern Alps with Hindeodus praeparvus. Weconclude therefore that the M. ultima(–S. mostleri) Zone of theuppermost Changhsingian is also present in South China.

As shown below, minor differences in taxonomy result only insmall differences in zonation (Fig. 2). These differences are limitedto (1) the base of the Changhsingian at the base of the C. hambast-ensis Zone in the Tethys, and (2) to the base of the C. wangi Zone inthe South Chinese intraplatform basins. C. hambastensis is part of C.wangi sensu lato; therefore, the base of the C. hambastensis Zone iscorrelative with the base of the C. wangi Zone in the stratotype sec-tion. The problem is that the holotype of C. wangi was sampledfrom a bed much higher than the base of the Changhsingian. A C.wangi morphotype is neither present at the base of the Changhsin-gian in the stratotype nor in Iran, Transcaucasia or elsewhere in theTethys. The appearance of C. hambastensis, however, begins exactlyat the defined base of the Changhsingian at the Meishan GSSP. Itthus appears suitable for precise correlation of the South Chineseintraplatform basins with the Tethys. Using C. hambastensis, the ex-act level of the base of the Changhsingian at Meishan can also beadopted for the Tethys, where the C. wangi morphotype s.s. ismostly absent.

We suggest not using the C. zheijiangensis Zone at Meishan (thiszone is not used in most of the Chinese conodont zonations of theChanghsignian GSSP) because its acceptance would abandon the H.parvus Zone and the first appearance datum of H. parvus is useful

because it defines the base of the Triassic. If, on the other hand,the definition of the C. zheijiangensis Zone would be changed withthe result that it is restricted to its uppermost Permian part abovethe C. meishanensis–H. praeparvus Zone, then the C. zheijiangensisZone would be a synonym of the M. ultima range zone which is alsopresent in South China (see Metcalfe and Nicoll, 2007). We prefertherefore to assign the uppermost Permian and the Triassic por-tions of the C. zheijiangensis Zone to the M. ultima and H. parvuszones, respectively (Fig. 2).

In shallow-marine successions, as in the Southern Alps (e.g., Pu-fels/Bula/Bulla and Tesero in the Dolomites, and Gartnerkofel inthe Carnic Alps), Outer Dinarides, Bükk Mountains of Hungary,and the Tauride Platform of Turkey, conodonts are rare, but theoccurrence of Hindeodus and Isaricella enables subdivision (Fig. 2)into the H. praeparvus Zone, the H. parvus Zone and the I. isarcicaZone (see Schönlaub (1991), for Gartnerkofel; Perri (1991), Farab-egoli and Perri (1998), Korte and Kozur (2005a), Korte et al.(2010), for Pufels/Bula/Bulla; and Perri (1991), Nicora and Perri(1999), for Tesero). The H. praeparvus Zone of shallow-water, gon-dolellid-free facies corresponds to the pelagic, gondolellid-bearingC. meishanensis–H. praeparvus Zone and the M. ultima–S. ? mostleriZone of Iran and South China.

In pelagic, gondolellid-bearing facies, the hindeodid conodontsmay be rare above the I. isarcica Zone. In these sediments, a gondo-lellid zonation may be applied. In the I. isarcica Zone, Clarkina car-inata (Clark) dominates the gondolellid facies. In the high-latitudesC. carinata begins in the upper Changhsingian. In the H. postparvus–H. sosioensis Zone, Neoclarkina discreta (Orchard and Krystyn)(=Neogondolella krystyni Orchard, see Henderson and Mei, 2007)

Table 1References for the P–T boundary sections investigated for bulk carbonate and organic-carbon carbon-isotopes. The classification, whether marine or non-marine, of the Australiansections is taken from Retallack and Krull (2006). In addition, carbon-isotope trends for non-marine P–T boundary successions have been published for the Karoo Supergroup(South Africa) on therapsid tooth apatite (Thackeray et al., 1990) and for organic matter of Russia (Foster et al., 1997).

Section or core Region Country Publication and material analyzed

Marine

Abadeh Central Iran Iran Heydari et al. (2000, carb, org, 2001, carb), Korte et al. (2004a, carb, org),Korte et al. (2010, carb), Richoz (2006, carb), Horacek et al. (2007b, carb)

Amol NE-Iran Iran Horacek et al. (2007b, carb)Bálvány Bükk Mts Hungary Haas et al. (2006, carb)Borup Fiord Pass Sverdrup Basin Canada Grasby and Beauchamp (2008, org)Brsnina Karavanke Mts Slovenia Dolenec et al. (1998, carb)Buchanan Lake Sverdrup Basin Canada Grasby and Beauchamp (2008, org)Buday’ah Hawasina nappes Oman Richoz (2006, carb)Çekiç Dag Central Taurus Turkey Richoz (2006, carb, org)Chaohu Anhui China Hansen (2006, org)Confederation Point Sverdrup Basin Canada Grasby and Beauchamp (2008, org)Çürük Dag Antalya nappes Turkey Baud et al. (1989, carb), Richoz (2006, carb)Dajiang Luodian, Guizhou China Payne et al. (2004, carb), Tong et al. (2007a, carb)Dawen Guizhou China Payne et al. (2004, carb)Daxiakou Xingshan, Hubei China Tong et al. (2007a, carb)Demirtas�-Kus�davut Alanya, S-Turkey Turkey Richoz (2006, carb)Dongpan Liuqiao, Guangxi China Zhang et al. (2006, org), Yin et al. (2007a, org)Drake Point Sverdrup Basin Canada Grasby and Beauchamp (2008, org)East van Hauen Pass Sverdrup Basin Canada Grasby and Beauchamp (2008, org)Emarat N-Iran Iran Baud et al. (1989, carb)Festningen (Kapp Starostin) Spitsbergen Norway Gruszczynski et al. (1989, carb), Mii et al. (1997, carb), Wignall et al. (1998,

org)Fishburn-1 Bonaparte Basin Australia Morante (1996, org)Fiskegrav Jameson Land Greenland Stemmerik et al. (2001, carb, org)Gaoqao Zunyi, Guizhou China Chen et al. (1991, carb)Gardony Transdanubian Range Hungary Hass et al. (2006, carb)Gartnerkofel (Reppwand) Carnic Alps, Southern Alps Austria Holser and Magaritz (1985, carb), Holser et al. (1989, carb), Magaritz and

Holser (1991, carb), Magaritz et al. (1992, org), Wolbach et al. (1994, org)Gerennavár Bükk Mts Hungary Korte and Kozur (2005a, carb)Guryul Ravine Kashmir India Baud et al. (1996, carb), Atudorei (1999, carb), Algeo et al. (2007b, org),

Korte et al. (2010, carb)Heping Pinggou, Guizhou China Krull et al. (2004, carb, org)Hovea 3 Perth Basin Australia Thomas et al. (2004, org)Huayun Lingshui, Sichuan China Yan et al. (1989, carb)Idrijca W-Slovenia Slovenia Baud et al. (1989, carb), Dolenec et al. (1999a, carb, org; Dolenec et al.

(2001, carb, org), Hansen et al. (2000, org), Schwab and Spangenberg (2004,carb, org)

Jolfa (Kuh-e-Ali Bashi) NW Iran Iran Holser and Magaritz (1987, carb), Korte and Kozur (2005a, carb), Korte et al.(2005a, carb), Kakuwa and Matsumoto (2006, carb), Richoz (2006, carb)

Kaki Vigla Salamis Greece Baud et al. (1989, carb)Kamura Kyushu Japan Musashi et al. (2001, carb, org)Kangjiaping Cili, Hunan China Wang et al. (2009, carb)Karabayır Yayla Central Taurus Turkey Richoz (2006, carb)Kemer Gorge Antalya nappes Turkey Baud et al. (1989, carb)Kopuk Dag Antalya nappes Turkey Richoz (2006, carb)Košutnik Creek Karavanke Mts Slovenia Dolenec et al. (1998, carb)Kuh-d-Yagma Central Zagros Mts Iran Wang et al. (2007, carb)Landu Nala Salt Range Pakistan Baud et al. (1996, carb)Lekang Guizhou China Peng et al. (2006, carb)Liangfengya Chongqing China Yin et al. (2007a, carb)Lindström Creek Sverdrup Basin Canada Grasby and Beauchamp (2008, org)Losar Spiti (Himachal Pradesh) India Atudorei (1999, carb, org), Galfetti et al. (2007, carb)Lung Cam Northern Vietnam Vietnam Son et al. (2007, carb)Masore Western Slovenia Slovenia Dolenec et al. (2004, carb, org)Meishan Changxing, Zhejiang China Chen et al. (1984, carb, 1991, carb), Xu et al. (1986, carb), Baud et al. (1989,

carb), Yan et al. (1989, carb), Xu and Yan (1993, carb), Wang et al. (1996, org),Li (1999, carb), Hansen et al. (2000, org), Jin et al. (2000, carb), Cao et al. (2002,carb, org, 2009, carb, org), Gruszczynski et al. (2003, carb), Nan and Liu (2004,carb, org), Korte et al. (2005a, carb), Zuo et al. (2006, carb), Riccardi et al.(2007, carb, org), Tong et al. (2007a, carb), Xie et al. (2007b, carb)

Nammal Gorge Salt Range Pakistan Baud et al. (1989, 1996, carb), Atudorei (1999, carb)Nhi Tao Northern Vietnam Vietnam Algeo et al. (2007a, carb, 2008, carb), Son et al. (2007, carb)Okhotsk NE Siberia Russia Zakharov et al. (2005, carb)Omolon NE Siberia Russia Zakharov et al. (2005, carb)Palazzo Adriano Sosio, Sicily Italy Korte et al. (2005b, carb)Palgham Kashmir India Baud et al. (1996, carb), Atudorei (1999, carb)Paradise 1–6 Canning Basin Australia Morante (1996, org)Peitlerkofel/Sas de Pütia/Sass de Putia Dolomites Italy Korte and Kozur (2005a, carb)Petrel-4 Bonaparte Basin Australia Morante (1995, org)Pufels/Bula/Bulla Dolomites, Southern Alps Italy Korte and Kozur (2005a, carb), Horacek et al. (2007a, carb), Gorjan et al.

(2008, carb, org)Qubu S-Tibet China Shen et al. (2006, carb, org)


Table 1 (continued)


San Antonio Auronzo, Southern Alps Italy Oberhänsli et al. (1989, carb)Sasayama Hyogo Japan Ishiga et al. (1993, org)Schuchert Dal Jameson Land Greenland Oberhänsli et al. (1989, carb), Twitchett et al. (2001, carb, org), Visscher

et al. (2004, carb), Fenton et al. (2007, carb, org)Seis/Siusi Dolomites, Southern Alps Italy Newton et al. (2004, carb) Kraus et al. (2009, carb)Selong-Xishan S-Tibet China Jin et al. (1996, carb), Shen et al. (2006, carb)Seres (Campill/Lungiarü/Longiarù) Dolomites, Southern Alps Italy Sephton et al. (2002, carb, org, Sephton et al. (2005, carb, org)Shahreza Central Iran Iran Korte et al. (2004b, carb), Richoz (2006, carb.)Shangsi Guangyuan, Sichuan China Xu et al. (1986, carb), Baud et al. (1989, carb), Yan et al. (1989, carb), Huang

(1994, carb), Hansen et al. (2000, org), Riccardi et al. (2007, carb, org)Shitouzhai Ziyun, Guizhou China Nan et al. (1998, carb)South Otto Fiord Mouth Sverdrup Basin Canada Grasby and Beauchamp (2008, org)South Pars Gas Field Persian Gulf Iran Insalaco et al. (2006, carb), Rahimpour-Bonab et al. (2009, carb), Esrafili-

Dizaji and Rahimpour-Bonab (2009, carb)Sovetachen Transcaucasia Armenia Xu et al. (1986, carb), Baud et al. (1989, carb), Zakharov et al. (2001, carb)Taho Shikoku Japan Musashi et al. (2001, carb, org)Takachiho Kyushu Japan Horacek et al. (2009, carb)Taiping Pingguo, Guangxi China Wang et al. (2001, carb), Krull et al. (2004, carb, org), Zhang et al. (2005,

carb)Tas�kent Central Taurus Turkey Richoz (2006, carb, org)Tenjin’maru Shikoku Japan Ishiga et al. (1993, org)Tern-3 Bonaparte Basin Australia Morante et al. (1994, org)Tesero Dolomites Italy Magaritz et al. (1988, carb), Korte and Kozur (2005a, carb)Thini Shu Central Nepal Nepal Baud et al. (1996, carb), Atudorei (1999, carb)Thongde Zanskar Himalaya India Baud et al. (1989, carb)Trzic Karavanke Mts Slovenia Dolenec et al. (1999b, carb, org), Hansen et al. (2000, org)Tulong S-Tibet China Yan et al. (1989, carb)Uomo Dolomites, Southern Alps Italy Horacek et al. (2007a, carb)Vedi Transcaucasia Armenia Baud et al. (1989, carb), Zakharov et al. (2001, carb)Wadi Alwa Ba’id Oman Atudorei (1999, carb), Richoz (2006, carb)Wadi Maqam Sumeini Group Oman Atudorei (1999, carb), Richoz (2006, carb)Wadi Musjah Ba’id Oman Atudorei (1999, carb)Wadi Sahtan Jabal Akhdar Mts Oman Atudorei (1999, carb), Richoz (2006, carb)Wadi Sawat Saih Hatat Oman Atudorei (1999, carb), Richoz (2006, carb)Wadi Wasit Ba’id Oman Atudorei (1999, carb), Krystyn et al. (2003, carb), Richoz (2006, carb, org)Wairoa Gorge South Island New Zealand Krull et al. (2000, org)West Blind Fiord Sverdrup Basin Canada Grasby and Beauchamp (2008, org)Western Dicksonland Spitsbergen Norway Wignall et al. (1998, org)Williston Lake British Columbia Canada Wang et al. (1994, org)Woodada-2 Perth Basin Australia Foster et al. (1997, org)Wujiaping Hanzhong, Shaanxi China Chen et al. (1991, carb)Yalang Pingguo, Guangxi China Yan et al. (1989, carb)Zal NW Iran Iran Korte et al. (2004c, carb), Richoz (2006, carb), Horacek et al. (2007b, carb)Zhongliangshang Luodian, Guizhou China Huang (1994, carb)Zhuzang Zijin, Guizhou China Yan et al. (1989, carb)Zuodeng Tiandong, Guangxi China Zhang et al. (2005, carb), Tong et al. (2007a, carb)

Non-marine

Allan Hills Central Transantarctic Mts Antarctica Retallack et al. (2005, org)Banspetali Raniganj Basin, NE India India Hansen et al. (2000, org), Sarkar et al. (2003, org)Bethulie Karoo basin South Africa MacLeod et al. (2000, carb, org), Ward et al. (2005, carb, org)Carlton Heights Karoo basin South Africa Ward et al. (2005, carb, org), Tabor et al. (2007, carb)Chahe Weining, Guizhou China Peng et al. (2005, org),Yu et al. (2005a, org), Yin et al. (2007b, org)Coalsack Bluff Central Transantarctic Mts Antarctica Retallack et al. (2005, org)Commando Drift Dam Karoo basin South Africa Coney et al. (2007, carb, org)Dalongkou Xinjiang China Metcalfe et al. (2009, org)Damodar valley NE India India de Wit et al. (2002, org)Denison NS-20 Bowen Basin Australia Morante (1996, org)Eddystone-1 Bowen Basin Australia Morante (1996, org)Graphite Peak Central Transantarctic Mts Antarctica Krull and Retallack (2000, org), Retallack et al. (2005, org)Injusiti Karoo basin South Africa Coney et al. (2007, carb, org)Jimsar Xinjiang China Hansen et al. (2000, org)Lootsberg Karoo basin South Africa de Wit et al. (2002, org), Ward et al. (2005, carb, org)Lucaogou Xinjiang China Metcalfe et al. (2009, org)Mahanadi valley NE India India de Wit et al. (2002, org)Morondava Southern Madagascar Madagascar de Wit et al. (2002, org)Mount Crean Central Transantarctic Mts Antarctica Retallack et al. (2005, org)Murrays Run-1 Sydney Basin Australia Morante (1996, org)Nelben Germanic Basin Sachsen-Anhalt Germany Korte and Kozur (2005b, carb), Hansen (2006, org)Newlands Coal Mine Bowen Basin Australia Hansen et al. (2000, org)Portal Mountain Central Transantarctic Mts Antarctica Retallack et al. (2005, org)Pranhita–Godavari Northeast India India de Wit et al. (2002, org)Shapeless Mountain Central Transantarctic Mts Antarctica Retallack et al. (2005, org)Son Valley NE India India de Wit et al. (2002, org)Thale Germanic Basin Sachsen-Anhalt Germany Korte and Kozur (2005b, carb)

(continued on next page)


Fig. 2. Conodont zonations and numeric ages for the Wuchiapingian (=Dzhulfian), Changhsingian (=Dorashamian) and Induan (=Brahmanian). Wuchia.: Wuchiapingian. Theboundaries of the ammonoid zones are commonly not correlated in detail with the boundaries of the conodont zones. In these cases the position of the boundary betweenadjacent ammonoid zones is shown as a dashed line.

Table 1 (continued)


Wapadsburg Karoo basin South Africa Ward et al. (2005, carb, org)Warda coalfield NE India India de Wit et al. (2002, org)Wybung Head New South Wales Australia Retallack and Jahren (2008, org)Zhejue Weining, Guizhou China Wang and Yin (2001, org)


dominates, and ranges a little above the top of the H. postparvus–H.sosioensis Zone. Thus, the H. postparvus–H. sosioensis Zone com-prises the largest part of the Neoclarkina discreta Zone. NeoclarkinaHenderson (in Henderson and Mei, 2007) is an important genusand forerunner of the Sweetospathodus and Neospathodus clines,from which the Triassic gondolellids derived: Neogondolella (Bend-er and Stoppel) through the transitional genus Chiosella (Kozur)and Paragondolella (Mosher) through the transitional genus Chen-gyuania (Kozur).

The continental lake deposits from the Buntsandstein of theGermanic Basin are subdivided and correlated by conchostracans(e.g., Kozur and Seidel, 1983a,b; Kozur, 1993, 1999; Bachmannand Kozur, 2004; Kozur and Weems, 2007, 2010). All conchostr-acan zones from the uppermost Permian to early Anisian are wellcorrelated with the international marine scale with magnetostra-tigraphy, Milankovitch cyclicity and, for the uppermost Permianto early Smithian interval, also with the carbon-isotope curve (Ko-zur and Mock, 1993; Kozur, 1999; Bachmann and Kozur, 2004;Korte and Kozur, 2005a; Korte et al., 2007, 2010). The P–T bound-ary lies between the Falsisca postera and the F. verchojanica Zones.

The low-latitude marine main extinction event occurred at theso-called event horizon (=EH = latest Permian Event Hori-zon = event boundary; green line in Fig. 2), and this is at the baseof the C. meishanensis–H. praeparvus Zone and at the base of theBoundary Clay (BC). Note that the BC is developed only in sectionsthat were deposited below the storm wave base, such as Abadeh,Jolfa, Zal, Shahreza and Meishan. Shallow-marine successions, suchas Pufels, Tesero, Gartnerkofel, Idrijca or Heping lack the BC. Anexception is the shallow-marine Gerennavár section (Bükk Moun-tains, Hungary) that is insignificantly deeper than the SouthernAlps successions. This is the only shallow-marine section with welldeveloped BC.

To discuss the possible causes for the isotope variations acrossthe P–T boundary it is necessary to take into consideration radio-metric ages published by Claoué-Long et al. (1991), Bowringet al. (1998) and Mundil et al. (2001, 2004). Furthermore, chrono-

logical estimates by Kozur ((2003); see also Bachmann and Kozur(2004), Kozur and Bachmann (2008)), who applied Milankovitchcycle theorem and used 252.5 Ma (Mundil et al., 2001) for theBed 28 of Meishan as reference value, are also taken into account.Combining all these geochronological data yields absolute ages of252.7 Ma for the base of the EH, 252.6 Ma for the P–T boundaryand 252.5 Ma for for the lower I. isarcica zone (Fig. 2).

3. Carbon-isotope values across the P–T boundary

3.1. Carbon-isotope trends in marine carbonates

3.1.1. Long-term trendThe long-term carbon-isotope trend across the P–T boundary

can be excellently demonstrated at the Meishan section (e.g., Caoet al., 2002) or at the Shahreza section, Central Iran (Korte et al.,2004b) because of the outstandingly calibrated conodont biostra-tigraphy (e.g., Wang and Wang, 1981a,b; Mei et al., 1998; Kozur,2004, 2005, 2007). Bulk carbonate carbon-isotopes from Shahreza(Fig. 3) show values between +3.0‰ and +4.5‰ for the Wuchiapin-gian (=Dzhulfian) Clarkina leveni, Clarkina transcaucasica, Clarkinaorientalis, Clarkina inflecta and Clarkina longicuspidata Zones andthe lower Changhsingian (=lower Dorashamian) C. hambastensisand C. subcarinata Zones, and these results are consistent withd13C data for well-preserved brachiopods from North Caucasus(Zakharov et al., 2000), eastern Greenland (Stemmerik et al.,2001), Jolfa and Meishan (Korte et al., 2005a). Note that brachio-pods are relatively resistant to diagenetic alteration and can be uti-lized to obtain the carbon (and oxygen and strontium) isotopevalues of past seawater (e.g., Veizer et al., 1997a,b, 1999). Unfortu-nately, because of the biotic crisis these fossils are not availableimmediately across the P–T boundary. The d13C-decrease for thebulk carbonates at Shahreza (Fig. 3) starts in the C. bachmanniZone, a first minimum is situated around the P–T boundary, anda second minimum occurs in the lower I. isarcica Zone.

Fig. 3. Wuchiapingian (=Dzhulfian) to Induan (=Brahmanian) carbon-isotope trendat Shahreza (d13C values from Korte et al. (2004b), and conodont zonation afterKozur (2005, 2007); 1: Clarkina nodosa Zone, 2: Clarkina zhangi Zone, 3: Clarkinairanica Zone, 4: Clarkina hauschkei Zone).


The Late Permian carbon-isotope trend from the Gartnerkofel(Holser et al., 1989) is similar to that from Shahreza. The biostrati-graphic level of the onset of the d13C decline in the Gartnerkofelcore cannot be exactly defined because conodonts are missing atthis stratigraphic level. The first and the second minima, on theother hand, are well defined by conodonts (Schönlaub, 1991) andsituated around the P–T boundary and in the I. isarcica Zone, withthe latter is subdivided into two separate events. The lower TriassicGartnerkofel carbon-isotope values remain about 1.5‰ lower,compared to the Late Permian data. This is consistent with bulkcarbonate d13C records of coeval other strata followed by distinctpositive and negative fluctuations (e.g., Baud et al., 1989; Atudorei,1999; Tong et al., 2002, 2005a,b,c, 2007a; Payne et al., 2004; Korteet al., 2005b; Zuo et al., 2006; Galfetti et al., 2007; Horacek et al.,2007a,b).

3.1.2. Short-term trends and their stratigraphic significance at the P–Tboundary

Carbon-isotope excursions are recognised in a wide range ofmarine and continental deposits, such as platform carbonates, pe-lagic carbonates, organic-rich shales, terrestrial palaeosols andlacustrine deposits. They represent major perturbations of the glo-bal carbon cycle and are often global events. Consequently, the on-set, duration and the termination of the P–T boundary negatived13C excursion should occur at the same time-stratigraphic level.If so, the carbon-isotope curve, especially its characteristic minimaand maxima, can be used for stratigraphic correlation.

In order to define a general carbon-isotope trend for the P–Tboundary interval it is essential to compare the negative d13Cexcursions of different sections (Fig. 4). For this purpose it is essen-tial to utilize preferentially biostratigraphically well-defined sec-tions, such as the Chinese and Iranian profiles (Fig. 2). Thesections with less biostratigraphic data are here also considered,but only for comparison. Carbon-isotope trends for the Yangtzeplatform (S-China), Perigondwanan margin of the Tethys (Kashmirand Oman), northern shelf of the Neotethys (Central Iran), south-ern shelf of Palaeotethys or shelf sea between Palaeo- and Neote-thys (NW Iran), Bükk Mountains (Hungary), Southern Alps(Austria and Italy), Western Slovenia and the Germanic Basin arecompiled here (Fig. 4). We utilize two time lines (1) the EH (greenline) and (2) the base of the Triassic (=first appearance datum(FAD) of H. parvus) (red line). Lithological logs and their respectivecarbon-isotope curves were expanded or reduced as required fordirect comparison.

At the Southern Chinese locality of Meishan, the P–T boundaryinterval was investigated in great detail (e.g., Sheng et al., 1984;Yin et al., 1992, 1996a,b, 2001; Yin and Zhang, 1996; Jin et al.,2000) and adjacent sections show identical successions (Fig. 4A;see also Table 1). The most spectacular d13C trend was reportedby Xu and Yan ((1993); but see Cao and Shang (1998)) with twoshort �8‰ negative shifts, one in the upper Boundary Clay (C. mei-shanensis–H. praeparvus Zone; lowest value: ��5.5‰) and one incarbonates of the lowermost part of the M. ultima–S. ? mostleriZone (lowest value: ��6.5‰). It is this strongly fluctuating car-bon-isotope curve that is presently utilized in discussions andmodels dealing with the P–T mass extinction (e.g., Bowring et al.,1998; Mundil et al., 2001; Peng et al., 2001; Twitchett et al.,2001; Berner, 2002; Lo et al., 2002; Newton et al., 2004; Algeoet al., 2007a; Farabegoli et al., 2007). Neither the minimum inthe carbonates of the lowermost part of the M. ultima–S. ? mostleriZone nor the strongly negative d13C values for the BC have beenreproduced in subsequently studies (Fig. 4A). Jin et al. (2000),Cao et al. (2002) and Xie et al. (2007b) reported data of only about�1.3‰ for the lower BC, these represent the lowest values for theentire P–T transition. Nan and Liu (2004), on the other hand, re-ported a decrease from �+3‰ to �+1‰ in the C. zhangi Zone(Fig. 4A), an �1‰ rise (starting somewhat below the event horizonin the higher bed 24e) and continuing in the lower BC (C. meishan-ensis–H. praeparvus Zone) and a further decline to values of �2‰ inthe lowermost Triassic (base of H. parvus Zone). A similar trendwas obtained for core material by Cao et al. (2009). Their low val-ues in the higher bed 24e (�3.2‰) represent the lowest for the en-tire succession (>150 m).

For the BC at Abadeh (Fig. 4F) and Jolfa (Fig. 4I) one (�0.6‰;empty circle) and two (�1.5‰ and �1.2‰; empty circles) low car-bon-isotope values were obtained from marly shales and siltstoneswith low carbonate contents. Coeval samples from Shahreza(Fig. 4G) and Zal (Fig. 4H) sections show distinctly higher valuesbetween +1.3 and +2.1‰ and between +0.7 and +1.1‰, respec-tively. Compared to Abadeh and Jolfa, the siltstones, clays and mar-ly clays of the BC at Shahreza and Zal are much less affected bymeteoric-water diagenesis. Therefore, the carbon-isotope records

Fig. 4. Lithology and d13Ccarb data for P–T boundary successions. Fifteen sections are correlated. 1 = C. changxingensis–C. deflecta Zone; 2 = C. zhangi Zone; 3 = C. iranica Zone;4 = C. hauschkei Zone; 5 = C. meishanensis–H. praeparvus Zone; 6 = M. ultima–S. ? mostleri Zone (= Permian part of C. zhejiangensis Zone for South Chinese intraplatform basins);5/6 = H. praeparvus Zone (=equivalent of zones 5 and 6 for shallow-water deposits without Clarkina); 7 = H. parvus Zone; 8 = I. isarcica Zone; green line: event horizon (EH) =base of the C. meishanensis–H. praeparvus Zone = base of the Boundary Clay (BC); red line: base of the Triassic (=FAD of H. parvus). (A) Meishan (South China; GSSP, Yin et al.(2001)). Dashed line after Cao et al. (2002; see also Jin et al. (2000)) and grey envelope-curve from Xie et al. (2007b), solid line and circles from Nan and Liu (2004). The often-cited carbon-isotope curve from Xu and Yan (1993) is incorporated for comparison (dotted line). (Figure modified after Tong et al. (2007b)). (B) Shangsi (South China), carbon-isotope data from Yan et al. (1989). The figure is adapted after Peng et al. (2001) and the biostratigraphy is modified as follows: Nicoll et al. (2002) reported the first H. parvusin Bed 30. However, this is not the FAD of H. parvus, but obviously the first occurrence datum (FOD) of this species because in the same level Isarcicella turgida s.s. also appears.Typical I. turgida s.s. occurs in the I. isarcica Zone and therefore Bed 30 belongs to the I. isarcica Zone, also indicated by the earlier FAD of Hindeodus eurypyge (Nicoll et al.,2002). This species appears at Meishan in bed 26, about 12 cm below the PTB. In Shangsi it appears in Bed 28 b (Nicoll et al., 2002). According to pers. comm. of Prof. Dr.Charles Henderson (Calgary) the PTB ‘‘is basically at the same position or just slightly above the FO of eurypyge at the Shangsi section.” His statement is based on graphiccorrelation and unpublished new data. The PTB therefore lies only insignificantly higher than shown in Riccardi et al. (2007), around the first minimum in delta C. Thedistinctly stronger minimum in Bed 30 lies not in the H. parvus Zone, but in the I. isarcica Zone, as indicated by I. turgida s.s. (C) Heping (S. China), carbon-isotope data fromKrull et al. (2004), lithology and carbon-isotope curve (blue) adapted from Lehrmann et al. (2005). The lowest data (yellow), missing in Lehrmann et al. (2005), were takenfrom the original Fig. 8 of Krull et al. (2004) and may be therefore somewhat inaccurate because of the poor resolution of the original figure. The base of the Triassic wasshifted to the FAD of H. parvus given in Lehrmann et al. (2005). (D) Guryul Ravine (Kashmir) modified after Korte et al. (2010). Most of the data by Atudorei (1999) werepublished in advance by Baud et al. (1996); in this early publication, however, the most important low value at the P–T boundary is missing. (E) Wadi Maqam (Oman)modified after Richoz (2006). (F) Abadeh (central Iran) modified after Korte et al. (2010). (G) Shahreza (central Iran) modified after Korte et al. (2004b) (H) Zal (NW Iran)modified after Korte et al. (2004c) (I) Jolfa (=Kuh-e-Ali Bashi, NW Iran) modified after Korte and Kozur (2005a). Carbon-isotope data for bulk carbonates for the Iraniansections Abadeh, Shahreza, Zal and Jolfa also were published by Richoz (2006) and for Jolfa by Kakuwa and Matsumoto (2006). Although the d13C curves are similar, thebiostratigraphic control in Richoz (2006) and Kakuwa and Matsumoto (2006) is less detailed compared with that of Korte et al. (2004b,c, 2010) and Korte and Kozur (2005a)and their data are therefore not incorporated in the present figure. (J) Gerennavár (Bükk Mountains, Hungary) modified after Korte and Kozur (2005a). (K) Gartnerkofel(Carnic Alps, Austria), modified after Holser et al. (1989) and Magaritz and Holser (1991), conodont ranges are taken from Schönlaub (1991). (L) Pufels (=Bula/Bulla; SouthernAlps, Italy): compiled carbon-isotope curve after Korte and Kozur (2005a) and Horacek et al. (2007b) discussed in Korte et al. (2010). I. isarcica staeschi, defining the lower I.isarcica Zone is, in general, also in Pufels, much less common compared with H. parvus (ratio 1:10 to 1:100) in the lowermost I. isarcica Zone. The beginning of I. isarcicareported by Farabegoli and Perri (1998) is most probably not the FAD, but the FOD of I. isarcica. Therefore the base of the I. isarcica Zone is drawn below the longer-lastingsecond carbon-isotope minimum which is somewhat below the FOD of I. isarcica according to Farabegoli and Perri (1998). The base of the I. isarcica Zone is assumed to lielower in the section than the level of the second carbon-isotope minimum. (M) Tesero (Southern Alps, Italy) modified after Korte and Kozur (2005a). The EH lies within theTesero Oolite Horizon. The highest occurrence of those Permian brachiopods that disappear below the Boundary Clay lies within the lower Tesero Oolite Horizon, indicatingthat this part of the section lies below the EH. The boundary between the Bellerophon Limestone and the Tesero Oolite Horizon is a slightly diachronous facies boundary. Thelower Tesero Oolite Horizon of the Tesero and other sections in the Dolomites changes in the eastern part of the Southern Alps into the uppermost, somewhat shallower-water Bellerophon Limestone, whereas the upper Tesero Oolite Horizon changes into the Mazzin Member. Therefore, in the eastern part of Southern Alps, the Mazzin Memberlies directly on the Bellerophon Limestone, and the boundary between the two lithostratigraphic units corresponds there to the EH (Kozur, 1994, 1998a,b; Kraus et al., 2009).(N) Idrijca (Slovenia) modified after Dolenec et al. (2001). (O) Könnern and Thale (Sachsen-Anhalt, Germany) Germanic Basin: combined continental Nelben-Thale sections(Korte and Kozur, 2005b). The carbon-isotope trend is plotted against the �100,000 years short eccentricity cycles, but the P–T boundary is correlated with the marine realmusing biostratrigraphy and other methods. As the carbon-isotope trends of the other sections are plotted against the conodont zonation, a double line separates thecontinental uppermost Permian–lowermost Triassic of the Germanic Basin from the marine sections. The second minimum in the Germanic Basin lies at the levelrepresenting a time interval �100,000 years later then that of the first minimum at the P–T boundary and corresponds to the minimum in the lower I. isarcica Zone. Thiscorrelation fits well with the approximate 100,000-year duration of the H. parvus Zone. The EH corresponds, in the basinal facies of the continental deposits of the GermanicBasin, to the base of the Buntsandstein above the Fulda Fm of the uppermost Zechstein. At this level, an abrupt climatic change can be observed from arid to humid conditions(Kozur, 1998a,b). The base of the Triassic lies at the base of the Oolite Bank a 2, just above the Falsica postera Zone. Only in a few places does the F. verchojanica (Molin) beginin this level. Therefore, in general, the boundary must be defined by the last occurrence datum (LOD) of the common F. postera (Kozur and Seidel). In Dalongkou (Sinkiang), thedisappearance of F. postera coincides with the LOD of Dicynodon, and therefore, this boundary can be well correlated with this datum level in South Africa. (For interpretationof the references to colour in this figure legend, the reader is referred to the web version of this article.)


of the BC at Shahreza and Zal should represent seawater signals,whereas the low values of Abadeh and Jolfa (empty circles in Fig4F and I) likely reflect diagenetically altered signals (Korte et al.,2004a,b; Korte and Kozur, 2005a). These considerations imply thatthe low d13C values in the BC at Meishan (Xu and Yan, 1993; Jin

et al., 2000; Cao et al., 2002; Xie et al., 2007b) are also the resultof diagenesis (Korte et al., 2004b; Tong et al., 2007b; Cao et al.,2009; see also Jin et al., 2000), an interpretation supported bythe fact that the Meishan-BC is deeply and strongly weathered,resulting in illite–smectite content of 95% (He, 1989; see also Yin

Fig. 4 (continued)


et al., 1992; Yang et al., 1993; Bowring et al., 1998; Yu et al., 2005b;Tong et al., 2007b).

As for Meishan (Nan and Liu, 2004, Fig. 4A; Cao et al., 2009), thecarbon-isotope values at Shahreza (Fig. 4G) and Abadeh (Fig. 4F)show an increasing trend in the BC. A similar feature at the samestratigraphic level is present also for Shangsi (Fig. 4B) and for shal-low-water sections that lack the BC, such as Heping (Fig. 4C), Gart-nerkofel (Fig. 4K), Pufels (Fig. 4L), Tesero (Fig. 4M) and Idrijca(Fig. 4N). Such short-term positive trend was reported also forother localities in Iran (Richoz, 2006; Kakuwa and Matsumoto,2006), the Southern Alps (Newton et al., 2004; Kraus et al., 2009)and Turkey (Richoz, 2006), strongly suggesting its global nature.We therefore suggest that there is no marked ‘primary’ d13C-min-imum at the BC stratigraphic level. The general carbon-isotopetrend is therefore as follows (Figs. 3–5): (a) the decline commencedin the C. bachmanni Zone and continued gradually throughout theC. nodosa, C. changxingensis–C. deflecta, C. zhangi, C. iranica and C.hauschkei Zones, (b) a short rise in the C. meishanensis–H. praepar-vus Zone interrupting the gradual decline (at Meishan and Abadehit commenced already in the latest C. hauschkei Zone), (c) the con-

tinuation of the decline in the higher C. meishanensis–H. praeparvusand the M. ultima–Stepanovites ? mostleri Zones (with occasionallya second order positive shift in the M. ultima–S. ? mostleri Zone)leading to a minimum very close to the P–T boundary. This mini-mum represents an important stratigraphic marker (Korte et al.,2005b). For Shahreza and Meishan (Nan and Liu, 2004), and formost of the other sections (e.g. Heping, Guryul Ravine, Pufels, Tese-ro, Germanic Basin) (Fig. 4), the values increase subsequently in thehigher H. parvus Zone and a second minimum occurs in the I. isar-cica Zone, a feature that confirms the carbon-isotope characteris-tics of the South Alpine Gartnerkofel succession of Holser et al.(1989; Fig. 4K). The second minimum can be subdivided into atwo-peaked event (Guryul Ravine, Gartnerkofel, Pufels, Tesero;see also Korte et al., 2010). The amplitude of the P–T boundaryd13C negative excursion varies between 4‰ and 7‰, with loweramplitudes characteristic of low-latitude pelagic sections (Abadeh,Zal, Sahreza) and higher fluctuations in the shallow-marine (Pufels,Gartnerkofel, Tesero) low-latitude and some mid-latitude (GuryulRavine) sections (Krull et al., 2000; Twitchett et al., 2001; Korteet al., 2010).


3.1.3. Deviating d13Ccarb-trends and stratigraphic difficultiesStratigraphic difficulties or deviations from the general d13C

trend exist for other sections and some of these, including mostof the Oman sections, Nammal Gorge (Pakistan), Bálvány-North(Bükk Mountains, Hungary) and Schuchert Dal (East Greenland),are reviewed below.

3.1.3.1. Oman and Pakistan. For most of the Oman sections (WadiMaqam is an exception; Fig. 4E), no concordant P–T boundary ex-ists because H. parvus and I. isarcica are not documented. In addi-tion, the latest Permian contains breccias, a feature at variancewith P–T boundary sections in all other investigated regions. InWadi Wasit (Krystyn et al., 2003), H. parvus occurs within a largebreccia block.

At the Nammal Gorge, a gap exists within the uppermost Perm-ian section, between the White Sandstone Member of the ChhidruFm and the Kathwai Member of the Mianwali Fm. The P–T bound-ary lies within the lower Dolomite Unit of the Kathwai Memberand not at its base. A d13C-minimum within the lower Kathwai(Atudorei, 1999) may indicate the P–T boundary, but its proximityto the basal Kathwai hiatus indicates that a freshwater overprintcannot be ruled out.

3.1.3.2. Bálvány-North (Bükk Mountains, Hungary). Haas et al.(2007) pointed out that the negative d13C-shift in the Bálvány-North section (Bükk Mountains, Hungary; Haas et al., 2006) is a‘‘quasi-symmetric negative peak” with a minimum about 30 cmbelow the top of the ‘‘Boundary Shale”. This negative peak is notassociated with a lithological change and was considered to be aprimary signal (Haas et al., 2007). However, this ‘‘Boundary Shale”is unlikely to be the latest Permian BC because (1) it is about 1 mthick, much thicker than the BC of the adjacent Gerennavár section(Fig. 4J); (2) it contains rich fossil fauna (brachiopods, bivalves andcrinoids), significantly more diverse than within the BC (and itstime equivalents) worldwide, and contains brachiopod species thatelsewhere became extinct prior to BC; and (3) its carbon-isotopecurve differs from other trends worldwide (Fig. 4 and chapter3.1). In the lower, approx 0.4 m-thick portion of the ‘‘BoundaryShale” a gradual drop, from about +1‰ to about 0‰, was reportedwhich conforms to the general upper Changhsingian trend belowthe BC. This is followed by about 0.4 m-thick interval with theabove-cited ‘‘quasi-symmetric negative peak”, having a minimumof <�4‰ and this feature is dissimilar to any other P–T boundarytrend. The upper 0.15 m of the ‘‘Boundary Shale” are characterizedby d13C values of about 0‰. The uppermost, approx. 5 cm thick sec-tion, consists of a sandy marlstone and siltstone. It is this level thatmay correspond to the lithologically similar BC at Gerennavár, buta short gap cannot be excluded.

In the Bálvány-North section, Haas et al. (2007) illustratedHindeodus praeparvus (Kozur) from Bed 3, about 0.8 m below the‘‘Boundary Shale”. However, the illustrated specimens belong toHindeodus typicalis (Sweet) which is common in low-latitude sec-tions for the entire Lopingian below the BC, but rare or absent fromthe BC up to the I. isarcica Zone. In mid- and high-latitudes, thisspecies is more frequent and occurs from the C. meishanensis–H.praeparvus Zone up to the lowermost Triassic. Hindeodus parvus(Kozur and Pjatakova) was reported from the Bálvány-East sectionwhich is younger than the ‘‘Boundary Shale”. The precise positionof the base of the Triassic in the Bálvány-North section and is notyet known.

3.1.3.3. East Greenland. Stratigraphic correlation of the Boreal fau-na close to the P–T boundary is problematic and discussed in theSupplementary Material. Regardless of the exact definition of theP–T boundary, the negative (�4‰) Late Permian d13C values ofbulk carbonates and the marked negative excursions about

20 m below the proposed first occurrence of H. parvus, as nega-tive as �10‰ (Twitchett et al., 2001; Visscher et al., 2004) and�21‰ (Oberhänsli et al., 1989) for the Schuchert Dal section ofJameson Land at East Greenland, require explanations. Suchlow values for the entire Late Permian were not detected else-where (Fig. 4). The reported carbon-isotope values for Late Perm-ian brachiopods from the Fiskegrav section in southern JamesonLand (Stemmerik et al., 2001) are at about +3‰ or +4‰, about7‰ heavier than bulk carbonate data. These brachiopod-datamatch well the low- and high-latitude Late Permian brachiopoddata of Korte et al. (2005a, 2008) as well as the bulk carbonatedata prior to Late Permian decline (Fig. 3; see also e.g., Baudet al., 1989; Holser et al., 1989; Cao et al., 2002, 2009; Korteet al., 2004a, 2010). We therefore suggest that the depleted bulkcarbonate d13C values for the Upper Permian reflect diageneticalteration; note that the Wordie Creek sediments are muddy silt-stones (Twitchett et al., 2001) and are not carbonate-bufferedlimestones.

3.2. Carbon-isotope trends in marine organic matter

Carbon-isotope fluctuations, if global, should be reflected inboth marine carbonates and marine phytoplankton because theyoriginate from the same carbon pool, the dissolved inorganic car-bon (DIC). The carbon-isotope composition of the organic matter(=d13Corg) is strongly depleted in 13C because plants discriminateagainst 13C during the production of their tissues. The fraction-ation for marine phytoplankton with C3-pathway for carbon fixa-tion is complex and depends on several factors, including pCO2

and pO2 (Hayes et al., 1989, 1999). The d13C of Permian organ-ic-matter carbon is about 28–32‰ more negative than that ofthe carbonate carbon (Hayes et al., 1999). The d13Corg curve acrossthe P–T boundary of the Festningen section at Spitsbergen (Wign-all et al., 1998) shows very similar features to the general carbon-ate carbon-isotope trend (Fig. 4). The biostratigraphic backgroundat Festningen, however, is too poor for precise correlation. Parallelcarbon-isotope trends for bulk organic matter and carbonateswere reported for marine sections at Gartnerkofel, Carnic Alps(Magaritz et al., 1992; Wolbach et al., 1994), at Seres near Cam-pill/Lungiarü/Longiarù, Dolomites (Sephton et al., 2002), at GuryulRavine, Kashmir (Algeo et al., 2007b) and at Kamura and Taho,southwest Japan (Musashi et al., 2001). However, not all sectionsexhibit such parallel trends (e.g., Korte et al., 2004a; Schwab andSpangenberg, 2004; Zhang et al., 2006; Riccardi et al., 2007; Yinet al., 2007a). Differing trends may result from low TOC contentsin the marine sediments, such as the Abadeh succession (Korteet al., 2004a), particularly for TOC-concentrations below 0.02%(Magaritz et al., 1992), although this cut-off limit for preservationof the primary d13Corg signal is debatable and might be higher orlower for certain deposits depending on the sediment type. Car-bon-isotope values of successions with low TOC concentrationsare therefore difficult to evaluate. Bulk organic carbon-isotopevalues of marine sediments can be influenced by a multitude ofadditional factors, such as heterogeneous biological origin ofTOC or variable marine–terrigeneous–bacterial mixtures (Whiti-car, 1996). Permian wood, for example, is characterized by higherd13C values than the coeval marine sourced organic matter (Faureet al., 1995; Foster et al., 1997; Korte et al., 2001). For instance,the negative carbon-isotope excursion in the marine successionof borehole Woodada 2 in Australia (Gorter et al., 1995; Fosteret al., 1997) was caused by a change from a predominance ofwoody tissues (�24‰) to predominance of acritarchs (up to�30‰). Similar explanations were advanced by Foster et al.(1997) also for the Tern-3 (Morante et al., 1994) and DenisonNS-20 (Morante, 1996; regarded as non-marine in Table 1) cores.These negative shifts therefore do not represent the negative P–T


boundary excursion. Note that the P–T boundary for non-marineAustralian sections (Morante, 1996; and compilation of Retallacket al., 2005: Fig. 13) was recently re-dated by Peter Jones (in Kor-te et al. (2008); see also Metcalfe et al., 2008) and situated some-what higher stratigraphically. A change from terrestrial woody toalgal-dominated sediments occurs also in the sediments of theHovea 3 core in the Perth Basin of W-Australia (Thomas et al.,2004) and an opposite trend is reported for Meishan (Cao et al.,2002).

While the factors that influence the d13C values of sedimentarymarine organic matter are complex (e.g., Tyson, 1995), the mostimportant for the chemostratigraphy at the P–T boundary is therole of bacteria which can modify considerably the primaryd13C. Methanotrophic bacteria prefer light 12C from the already13C-depleted organic matter for assimilatory uptake and conse-quently their biomass is even more severely depleted. Organicmatter from methanotrophs may thus shift the sediment-TOC tolighter carbon-isotope values. Moreover, in present-day marineenvironments, particularly in restricted basins with anoxic condi-tions (e.g., Black Sea), bacteria reduce sulphate to sulphide (Trullet al., 2001), with green photosynthetic bacteria or green sulphurbacteria (=Chlorobiaceae) utilizing H2S and CO2 in the photic zonefor their anaerobic photosynthesis. During this photosyntheticreaction the bacteria do not discriminate as strongly against 13Cas in the C3-pathway (Sirevag et al., 1977). As a result, the Chlo-robiaceae biomass is much less depleted in 13C compared to aer-obic photosynthesisers (Summons and Powell, 1986). Theresultant biomass is characterized by (relatively 13C-‘enriched’)compound-specific d13C values of about �12‰ to �14‰ (Koop-mans et al., 1996; Grice et al., 1996), but this fractionation takesplace in anoxic waters in which the CO2 might be already de-pleted in 13C by microbial activity and this leads to further 12C-enrichment in the produced biomass (e.g., Struck et al., 2001).The d13Corg values in such environments are difficult to evaluate,particularly when no organic geochemical biomarker informationis available.

For the P–T boundary, widespread anoxic conditions occurredduring the deposition of several successions (e.g., Wignall and Hal-lam, 1992; Isozaki, 1994, 1997; Kajiwara et al., 1994; Knoll et al.,1996; Wignall and Twitchett, 1996; Wignall et al., 1998; Katoet al., 2002; Kidder and Worsley, 2004; Nielsen and Shen, 2004;Grice et al., 2005; Kump et al., 2005; Riccardi et al., 2006, 2007;Hays et al., 2007; Xie et al., 2007a; Algeo et al., 2008; Grasby andBeauchamp, 2009; Takahashi et al., 2009). Anoxic conditions, interms analogous to Tyson and Pearson (1991), are indicated bythe presence of compounds or alteration products of green sulphurbacteria from the shallow-marine photic zone in the Hovea 3 core(W-Australia) (Grice et al., 2005). For the special situations of thewidespread anoxic conditions at the P–T boundary, the signal de-pends on the proportion of organic matter produced by aerobic/anaerobic photosynthesisers and on the number of isotopic fracti-onations steps. In addition, the TOC-d13C of the sediments dependsalso on the degree of preservation of the bacterially-derived organ-ic matter.

The above considerations suggest that the d13Corg of bulk organ-ics at the P–T boundary may have been, at least temporarily,decoupled from the oceanic surface production of carbonates gen-erating difficulties for utilization of d13Corg from bulk organics forstratigraphic purposes (Riccardi et al., 2007). Biomarker analysesmay be a useful tool for identification of organisms participatingin the production of the organic matter (Sephton et al., 2002; Tho-mas et al., 2004; Korte et al., 2004a; Schwab and Spangenberg,2004; Grice et al., 2005; Xie et al., 2005, 2007a; Fenton et al.,2007; Hays et al., 2007; Wang and Visscher, 2007; Luo et al., inpress) and enable better understanding of the organic-mattercarbon-isotope values.

3.3. Carbon-isotope trends from continental successions

The utilization of bulk carbon-isotope values from terrestrial or-ganic matter as a stratigraphic tool is possible, but several factorsmust be taken into account. Compared with marine deposits, it ismore difficult to delineate the secular d13C trend from continentalsections across the Palaeozoic–Mesozoic transition because, formost levels, the biostratigraphic subdivision and the correlationwith the marine scale are not very detailed. However, the conti-nental P–T boundary itself can be stratigraphically identified (seereview in Korte et al. (2010)) in several successions (e.g. Dalongkouof Xinjiang, Carlton Heights of South Africa, Nelben section of theGermanic Basin) by the disappearance of the conchostracan Falsicapostera (Kozur and Seidel), the FAD of the conchostracan F. ver-chojanica (Molin) or/and the LOD of the vertebrate Dicynodon, aswell as by the d13C-minimum coinciding in South Africa with theLOD of Dicynodon (Ward et al., 2005). Further away from the P–Tboundary it is impossible to correlate continental sediments insuch detail because the biostratigraphic framework of continentalsections is generally poorer than in marine successions.

The carbon-isotope values from bulk terrestrial organic matterin palaeosols are controlled by a plethora of factors such as origin,redox conditions, and/or degree of microbial decay. These isotopicvalues depend therefore on palaeoenvironmental conditions thatmight change in very short time, even seasonally. It is thereforeexceedingly difficult to recognise secular d13C changes. Krull andRetallack (2000) reported carbon-isotope variations of several per-mil in palaeosols of Late Permian and Early Triassic ages in Antarc-tica that show similarities to patterns of modern soil profiles. TheEarly Triassic palaeosols show strong 13C-depletion interpreted asmicrobial methane oxidation, reflecting elevated methane concen-trations in the coeval soil–atmosphere system. Such interpretationis in accord with the extremely 13C-depleted rocks of Early Triassicreported from other Antarctic sections (Retallack et al., 2005). Inter-estingly, a distinct negative excursion at the P–T boundary is miss-ing. The bulk carbon-isotope trends illustrated by de Wit et al.(2002) for Madagascar, South Africa and India have insufficient fos-sil data to evaluate the curves stratigraphically, and Sarkar et al.(2003) do not have d13C values across the P–T boundary in India.

Bulk d13Corg values were reported for four biostratigraphicallywell defined non-marine P–T boundaries (Smith, 1995; MacLeodet al., 2000) of the Karoo basin in South Africa (Ward et al.,2005). No distinct negative carbon-isotope excursion was reportedfor the Wapadsburg and Bethulie sections, but a carbon-isotopetrend similar to the marine standard (Fig. 4) exists for the CarltonHeights and the Lootsberg sections. For the two latter localities,d13Ccarb values were obtained also from carbonate nodules andfrom palaeosol carbonates that have been utilized for model esti-mates of ancient atmospheric CO2 levels (Cerling, 1991, 1999;Ekart et al., 1999; but see Quast et al., 2006). In this scenario, thelow pCO2 is reflected in low d13C values, and vice versa, becauseisotopic fractionation of soil plant material depends on the atmo-spheric CO2 level (Retallack, 2004). The d13Ccarb values of carbonatenodules from the Carlton Heights and Lootsberg sections (Wardet al., 2005) show, however, the same trend as their d13Corg dataand even the marine carbonate standard (Fig. 4; see also the Bethu-lie section; MacLeod et al., 2000). Tabor et al. (2007) argued thatthe Carlton Heights carbonate nodules do not reflect atmosphericconditions due to their precipitation under ‘‘poorly drained swam-py conditions”. The carbon-isotope trends in MacLeod et al. (2000)and Ward et al. (2005), however, do show a minimum at the P–Tboundary and a second minimum corresponding to the isarcica-minimum. Despite similar trends, the carbonate-nodules basedtrend shows greater amplitudes than marine carbonates, inter-preted as anaerobic overprint of the global carbon-isotope signalin South African palaeosols (Tabor et al., 2007).


The carbon-isotope values of continental lake carbonates of theLower Buntsandstein (Korte and Kozur, 2005b) are in good agree-ment with the marine d13Ccarb standard trend. Note, nevertheless,carbon-isotope values in terrestrial samples are generally lowerthan those of marine carbonates (Talbot, 1990). A first minimumoccurs at the P–T boundary followed by an increasing trend anda second minimum that corresponds to the minimum in theisarcica Zone (Fig. 4 O, see also Fig. 2). These results support ateleconnection between ocean and continental lakes throughatmospheric CO2-exchange, arguing that ocean surface and atmo-sphere behaved effectively as one system during the Paleozoic–Mesozoic transition.

4. Causes of the negative carbon-isotope excursion

It is generally accepted that the carbon-isotope ratio of theocean/atmosphere is principally controlled by the burial andre-oxidation of 13C-depleted organic matter in the oceans and oncontinents (Scholle and Arthur, 1980; Kump and Arthur, 1999).However, changes in the amount of buried/reoxidized organic car-bon are usually effective on longer time scales and additional fac-tors have therefore been proposed to have been responsible for thenegative d13C excursion around the P–T boundary. The proposedalternatives include: (1) erosion of organic matter or soil; (2) re-duced primary productivity in oceanic surface layers due to bioticcrisis that resulted in a diminished transport of organic matter todeeper water layers (‘Strangelove oceans’); (3) the influence ofthe Siberian Trap and contemporaneous volcanism by (i) exhala-tion of 13C-depleted CO2, (ii) destabilisation of methane from per-mafrost soils and/or pre-Trap coals, or (iii) thermal metamorphismof pre-Trap organic-rich sediments; (4) anoxia reaching shallow-marine levels by an upward rise in the chemocline, by ocean over-turn(s) or by a diminution or absence of oxygen-rich cold bottomcurrents; and (5) dissociation of isotopically light methane hy-drates from oceans. Several authors (e.g., Berner, 2002; Korteet al., 2004a; Sephton et al., 2005; Erwin, 2006; Algeo et al.,2008) pointed out that more than one source may have been ulti-mately responsible.

To evaluate the causes of the negative d13C excursion, an accu-rate estimate of its duration is necessary. Bowring et al. (1998) pro-posed <165,000 years for the entire carbon-isotope excursion, butthis estimate was not confirmed by subsequent studies (Metcalfeet al., 2001). Ages calculations, based on Milankovitch cyclicityand sedimentation rates (Rampino et al., 2000; Mundil et al.,2001, 2004; Kamo et al., 2003; Kozur, 2003, 2005, 2007; Bachmannand Kozur, 2004; Kakuwa and Matsumoto, 2006; see also Korteet al., 2010) suggest that the decline in d13C values (C. bachmanniZone to H. parvus Zone) lasted for about 500,000 years (Figs. 2and 3) and the interval from EH to the base of the Triassic com-prises about 120,000 years (Figs. 2 and 4). Note that the sudden�2.3‰ d13C decline in the Gartnerkofel core between 228 and224 m (Magaritz and Holser, 1991) calculated by Rampino et al.(2000) to represent �30,000 years accounts for about one fourthof the time span between the EH and the P–T boundary (Fig. 4K).

4.1. Erosion of organic matter and reduced primary productivity

13C-depleted organic matter may be oxidized at exposed conti-nental shelves, due to a global sea level fall, and transported intothe oceans (Holser and Magaritz, 1987, 1992; Baud et al., 1989).Yet, recent geological observations do not support the idea thatglobal sea level fall straddled the P–T boundary. There was only ashort-term sea level drop that started in the C. zhangi Zone and cul-minated in the C. iranica Zone (Fig. 2 and 3; chapter 2). However,during the C. hauschkei Zone (and the equivalent Boreal Otoceras

concavum and lower O. boreale Zones) a distinct sea-level rise com-menced, and continued across the P–T boundary into the earliestTriassic. This is reported for several sections (e.g., Wignall and Hal-lam, 1992; Yin et al., 1996b; Hallam and Wignall, 1999; andnumerous papers which show the beginning of the transgressionat the base of the Boreal Otoceras faunas and in Bed 24e immedi-ately below the Boundary Clay at Meishan; e.g., Yin et al., 2001).

A strong decline in marine and terrestrial productivity mighthave also produced the negative carbon-isotope shift by a varietyof ways in the oceans or on land (Magaritz, 1989; Wang et al.,1994; Visscher et al., 1996; Broecker and Peacock, 1999; Wardet al., 2000; Twitchett et al., 2001; Grard et al., 2005; Rampinoand Caldeira, 2005), such as the availability of light carbon fromthe main extinction event itself (e.g., Jin et al., 2000). Alternatively,a shutdown of the biological pump in the oceans may have re-sulted in the negative d13C excursion. Sea-surface water is enrichedin 13C compared to deeper water (Kroopnick et al., 1972) due tophotosynthetic withdrawal of 12C in the photic zone and its trans-port to deeper water and seafloor. A sudden reduction or collapsein primary productivity in oceanic surface layers, may result in adiminished transport of 13C-depleted organic matter to deeperwater layers (‘strangelove oceans’; Hsü and McKenzie, 1985;Kump, 1991). Simulations by Rampino and Caldeira (2005) suggestthat such a scenario may have produced a 3‰ negative shift at theextinction event. In another proposition, Ward et al. (2000) andSephton et al. (2005) argued that soils may have been destabilizedby the destruction of land plants, leading to enhanced erosion ofterrestrial organic matter, thus contributing to the negative d13Cexcursion.

For all these interpretations it is important to note that themain extinction event was much later than the start of the negativecarbon-isotope excursion (e.g., Korte et al., 2010). Moreover, thegeneral carbon-isotope trend of carbonates is characterized by aninterim increase within the Boundary Clay and its time-equiva-lents, starting somewhat earlier than the main extinction event(Figs. 3 and 4; see also Newton et al., 2004; Kakuwa and Matsum-oto, 2006; Richoz, 2006; Kraus et al., 2009); no decline in d13C wasobserved at that time. This short-term positive excursion in the C.meishanensis–H. praeparvus Zone (and M. ultima–S. ? mostleri Zone;Fig. 4) may have been caused by a positive feedback of ocean anox-ia on nutrient availability due the enhanced regeneration of phos-phate from the sediments that in turn triggered higherbioproductivity (Payne and Kump, 2007; Korte et al., 2010; butsee Saltzman, 2005). Demise of land plants, on the other hand,lasted several million years into the Early Triassic (Looy et al.,1999) and may have resulted in advanced continental weathering,even in wet regions, that produced widespread clastic successionsof the Alps, Western Carpathians, Germanic Basin, Russian Plat-form or western North America (Korte et al., 2003). Such scenariomay have maintained low sedimentation rates of organic matteron land (Retallack et al., 1996) resulting in the persistence of lowd13C values in the earliest Triassic. The severe changes in Earth’secosystem, due to the demise of vascular land plants, formerly rep-resented a huge storage reservoir, may also have resulted in lesscarbon sequestered in the biosphere (Broecker and Peacock, 1999).

4.2. Volcanism

The Siberian Trap volcanism, which straddled the P–T boundaryand was much larger in size than previously thought (e.g., Kozur,1998a,b; Reichow et al., 2002; Ivanov et al. 2005, 2009; Vyssotskiet al., 2006; Saunders and Reichow, 2007), might have directlyand indirectly influenced the seawater/atmosphere d13C value.The plateau basalts, representing the main phase of this volcanism,lasted about 600,000 years (recently accurately dated by Kamoet al. (2003)). This phase was preceded by explosive volcanism


(very thick tuffs) that commenced somewhat less than400,000 years earlier (Korte et al., 2010).

At first, 13C-depleted volcanic CO2 might have influenced theocean/atmosphere d13C value (Berner, 2002; Grard et al., 2005;Hansen, 2006). Volcanic CO2 (d13C: ��5‰) acting on seawaterd13C (that was relatively heavy in the Late Permian) at the P–Tboundary is unlikely to have been a main cause of the negativeexcursion because this would only have produced a small d13C-shift (Kump and Arthur, 1999; Wignall, 2001; Berner, 2002). Han-sen (2006), however, argued that carbon-isotope values of volcanicCO2 may have been considerably lower than �5‰, a debatableproposition.

On the other hand, the release of methane from older coals and/or thermal metamorphism of such coals, organic-rich sedimentsand petroleum-bearing strata by the Siberian Trap magmatism(lava flows, sills, dykes) may have donated significant amounts of13C-depleted CO2 to the atmosphere/hydrosphere system (Svensenet al., 2004, 2009; Payne and Kump, 2007; Retallack and Jahren,2008; Korte et al., 2010). The influence of the Siberian Traps onocean/atmosphere d13C values is strongly supported by the shapeof the carbon-isotope curve: the more intensive the volcanic activ-ity, the lower the carbon-isotope values (Korte et al., 2010).

4.3. Ocean anoxia

Ocean anoxia was widespread in the latest Permian and theseconditions likely reached even shallow-marine areas due to up-ward rise/s of the chemocline, oceanic overturn/s, or via weakcold oxygen-rich bottom-water currents (e.g., Kajiwara et al.,1994; Knoll et al., 1996; Wignall and Twitchett, 1996, 2002; Iso-zaki, 1997; Kozur, 1998a,b; Hotinski et al., 2001; Kump et al.,2005; Algeo et al., 2007a, 2008). In stagnant oceans, organic mat-ter becomes oxidized under aerobic conditions only present insea-surface waters. After oxygen (and subsequently nitrate) is ex-hausted in the subsurface, sulphate reduction takes over underanoxic conditions and H2S is generated (and iron sulphides pre-cipitate in the presence of dissolved iron), permitting the pCO2

to continuously increase, ultimately reaching high concentrationsin these water bodies (Knoll et al., 1996). This CO2 has low d13Cvalues and represents a large source of isotopically light carbon.When and if these deeper waters return to shallow-water depthsclose to the sea surface, a fall in shallow-water carbonate d13C re-sults (Küspert, 1982). For the P–T boundary, it has been proposedthat massive ocean mixing temporarily occurred (Kajiwara et al.,1994). An overturn or movement of chemocline upwards led thento a fusion of CO2-rich and H2S-rich bottom and surface waters(Knoll et al., 1996; Kump et al., 2005; Riccardi et al., 2006,2007). Such anoxia may have produced toxic conditions in partsof the oceans (Knoll et al., 1996) or even in the atmosphere(Kump et al., 2005), and may have contributed to the negatived13C excursion (Malkowski et al., 1989; Hoffman et al., 1991;Wignall and Hallam, 1992; Korte et al., 2004a; Wignall et al.,2005; Riccardi et al., 2007; Algeo et al., 2007a, 2008; Hayset al., 2007). Dysoxic or anoxic conditions have been reportedfor several regions, such as Spitsbergen, China, W-Australia, Viet-nam, Kashmir, Western Canada (e.g., Wignall et al., 1998, 2005;Grice et al., 2005; Riccardi et al., 2006, 2007; Hays et al., 2007;Algeo et al., 2008), but a long-lasting anoxia in the entire low-lat-itude deep-sea Panthalassa (Isozaki, 1994, 1997; Kato et al., 2002)may be an exaggeration (Kakuwa, 2008; see also Hotinski et al.,2001; Zhang et al., 2001). Moreover, oxic conditions across theP–T boundary apparently existed in the Iranian sections. The lat-ter is indicated by synsedimentary carbonate cements at Abadeh(Heydari et al., 2003), by red colour of sediments across the P–Tboundary at Shahreza and Zal, and by a diverse ostracod faunarequiring high oxygen levels at Abadeh, Jolfa, Shahreza and Zal

(Kozur, 2005, 2007; Richoz, 2006). At Jolfa, anoxic conditionswere proposed for deposition of the strata about two metres be-low the BC (within the red Paratirolites Limestone) by Kakuwaand Matsumoto (2006) because of a negative cerium anomaly.In the horizon with the cerium anomaly filter-feeding ostracods,at that time Cavellinidae and Hollinacea which could live inlow-oxygen water, are absent and only species are present thatrequired high oxic conditions, as reported by Belousova (1965)for the adjacent Dorasham succession. The same fauna occurs inthe entire Iran and Transcaucasia realms. While, dysoxic – partlyanoxic – conditions are indicated by the black marker horizon alittle above the P–T boundary, at this level the cerium valuesare ‘‘normal”. These results show that Kakuwa and Matsumoto’s(2006) negative Ce-anomaly about two metres below the BCcould not have been caused by suboxic/anoxic conditions. Inaddition, Kakuwa and Matsumoto (2006) did not analyse the crit-ical element praseodymium (Pr) and this is necessary to identifyfairly a subtle Ce-anomaly (Kato and Isozaki, 2009).

Nevertheless, upward rises of the chemocline and/or oceanoverturns in regions with anoxic water conditions have most likelycontributed to the negative carbon-isotope excursions becauseseveral such events in the latest Permian and earliest Triassic werereported by Algeo et al. (2008) for Nhi Tao. In this case short-termnegative carbon-isotope excursions of <1‰ correspond to enrich-ments in sulphide–sulphur in the sediments indicating an upwell-ing of deep sulphidic watermasses. These short-term eventsstarted somewhat later than the EH and persisted into the EarlyTriassic. Taking into account that anoxia was widespread in the lat-est Permian and earliest Triassic, regional influence on seawaterd13C by recurrent upwelling of euxinic waters is likely and may ex-plain deviations of some sections from the general carbon-isotopetrends above the EH. These results also show that the anoxia mostprobably impacted the ocean/atmospheric d13C only in the latestphase of the negative excursion.

4.4. Hypothesis of methane hydrate dissociation

Methane hydrates, stored in sea-floor sediments and perma-frost soils as frozen clathrates, will dissociate when warming takesplace (Kvenvolden, 1988, 1993; MacDonald, 1990; Nisbet, 1990;Paull et al., 1991; Dickens et al., 1995; Dorritie, 2002). CH4, whichoxidizes quickly to CO2 in the ocean and atmosphere, is extremelydepleted in 13C (d13C: <�60‰) and therefore represents anotherpotential source of the isotopically light carbon. Such release ofseafloor and permafrost methane has been repeatedly proposedto have produced the negative carbon-isotope shift at the P–Tboundary (e.g., Erwin, 1993, 1994; Morante et al., 1994; Krulland Retallack, 2000; Krull et al., 2000, 2004; Twitchett et al.,2001; Gruszczynski et al., 2003; Berner, 2002; de Wit et al.,2002; Sarkar et al., 2003). A plethora of factors was involved to ex-plain the warming required for this scenario. A decline in silicateweathering and nutrient supply or CO2 release from Siberian Trapvolcanism were involved as sources for significantly elevatedatmospheric carbon dioxide content that then triggered the globalwarming (Wignall and Hallam, 1993; Kidder and Worsley, 2004;Twitchett, 2007a,b) causing methane release that further amplifiedthe strong greenhouse conditions (Erwin, 1993; Kidder and Wors-ley, 2004; Racki and Wignall, 2005; Winguth and Maier-Reimer,2005). Some authors (Benton and Twitchett, 2003) even claimedthat it resulted in a ‘runaway greenhouse’.

A Late Permian warming trend is supported by model consider-ations (Kiehl and Shields, 2005; Rampino and Caldeira, 2005).Conodont d18O values also indicate a 4–6 �C warming in low-lati-tudes during transition from the late Permian to the early Triassic(Korte et al., 2010), but note that the decline in oxygen-isotopetrend may result from, or be amplified by, palaeobathymetric

Fig. 5. General carbon-isotope trend across the P–T boundary constructed from thestratigraphically well defined d13C data from Meishan (Nan and Liu, 2004), GuryulRavine (Korte et al., 2010), Abadeh (Korte et al., 2004a, 2010), Shahreza (Korte et al.,2004b), Pufels/Bula/Bulla (Korte and Kozur, 2005a; Horacek et al., 2007b; Korteet al., 2010) and the Gartnerkofel core (Holser et al., 1989; Magaritz and Holser,1991). 1: C. changxingensis–C. deflecta Zone, 2: C. zhangi Zone, 3: C. iranica Zone, 4: C.hauschkei Zone, 5: C. meishanensis–H. praeparvus Zone (5 + 6: H. praeparvus Zone forshallow-water without Clarkina such as the Southern Alps), 6: M. ultima–S. ?mostleri Zone = Permian part of C. zhejiangensis (for South Chinese intraplatformbasins), 7: H. parvus Zone = Triassic part of C. zhejiangensis (for South Chineseintraplatform basins), 8 = I. isarcica Zone; carbon-isotope data below the H.praeparvus Zone at Pufels/Bula/Bulla are not incorporated.


and palaeolatitudinal effects. The best evidence for a dramaticchange from cold to warm temperature prior to the P–T boundarywas obtained from the high-latitude E-Australian palaeosols. Forthe Sydney Basin, Retallack (1999a) reported that the palaeosolsat the top of the Illawarra Coal Measures Group and the overlyingNarrabeen Group show characteristics that are typical of present-day latitudes of 68–70 �S and 40–58 �S, respectively. He alsoshowed that the climate change took place at the same time asthe extinction event (Conaghan et al., 1994; Retallack et al.,1996). The basal part of the Narrabeen Group (and Rewan Groupin the Bowen Basin), however, contains the Australian Proto-haploxpinus microcorpus palynofloral Zone thought to be the latestPermian (Helby et al., 1987; Price, 1997), an assignment recentlyconfirmed by Metcalfe et al. (2008) by conodont and palynologicalevidence from Western Australia, where the P. microcorpus Zone isof late (but not latest) Changhsingian age (cf. Jones in Korte et al.(2008)). The abrupt extinction of plants in the continental bedsof Australia took place during the late Changhsingian climatechange and thus likely does not coincide with the main marineTethyan extinction event. This pattern of terrestrial disaster pre-ceding the marine low-latitude extinction event may apply alsoto other regions (Wignall, 2007). Warm oceanic temperatures overwide latitudinal range for parts of the Early Triassic are indicatedby a low sea-surface pole-to-equator temperature gradient, re-flected in latitudinal distribution of ammonoids (Brayard et al.,2006; Galfetti et al., 2007) and by high-latitude Antarctic deeplyweathered palaeosols (Retallack and Krull, 1999). A considerablewarming of high-latitudes is also supported by the appearance oflow-latitude, warm-water biota in Spitsbergen in the early Gandar-ian (Wignall et al., 1998).

These above-mentioned geological and palaeontological obser-vations indicate that the warming took place at the Palaeozonic–Mesozoic transition. If this warming was uninterrupted, the meth-ane-release hypothesis would be an attractive possibility to ex-plain the carbon-isotope anomaly and the mass extinction. Butwas there an uninterrupted warming trend throughout the mainextinction and across the P–T boundary? We do not have a contin-uous oxygen-isotope dataset for fossils with constant palaeobathy-metry (see Korte et al., 2010), but bulk carbonate d18O values showa negative shift for the entire time period across the P–T boundary(e.g., Holser et al., 1989). Bulk carbonate d18O values, however, rep-resent (1) at least in part the temperature of rock lithification andnot just that of the past seawater, and (2) are prone to diageneticalteration.

Reviews of palaeontological, geochemical and sedimentologicaldata indicate that cooling periods took place in (1) the lower C.zhangi Zone and (2) over the time span of the C. meishanensis–H.praeparvus Zone to the top of the M. ultima–S. mostleri Zone (Kozur,1998a,b; Krassilov and Karasev, 2009). We argue that these coolingperiods were triggered by aerosols of the Siberian Traps and con-temporeanous volcanism in South China (Campbell et al., 1992;Renne et al., 1995; Kozur, 1998a,b, 2007; Isozaki, 2007,2009b). Itis therefore open to debate whether such data provide the best pal-aeoclimate information. To obtain global temperature informationfor the critical time span additional studies are necessary. Thus, itis difficult to state how much of the negative carbon-isotope excur-sion was caused by an oceanic methane event. Additional informa-tion for the evaluation of the methane hypothesis exists from thed13C trend itself. If a methane event caused the mass extinction(Ryskin, 2003) then a sudden sharp negative excursion in d13C(as proposed for the Paleocene–Eocene Thermal Maximum, Dick-ens et al., 1995) would be expected to be coeval with the mainextinction event (e.g., Benton and Twitchett, 2003): but this is atodds with the general P–T boundary carbon-isotope curve (Figs. 4and 5 and chapter 3.1.2). In addition, a release of methane hydratesis a sudden event and would produce a rapid d13C-decrease (see

Katz et al., 1999; Norris and Röhl, 1999). The carbon-isotope de-crease, however, is gradual and had already started several100,000 years before the EH (Fig. 3). Such a long time span forthe negative trend is difficult to explain by a dissociation of meth-ane hydrates.

It has been further suggested that the Siberian Trap volcanismmay have destabilized isotopically light methane hydrates thatwere stored in surrounding permafrost soils (Dorritie, 2002;Racki and Wignall, 2005). For present-day high-latitude perma-frost soils, methane hydrates have been reported up to depthsof several 100 m (Collett, 1993; Collett et al., 2010). Becauseferns, other higher plants, a very rich conchostracan fauna, andeven the branchiosaurid Temnospondyli (Werneburg, 2009) oc-curred before and between the volcanic rocks associated withthe Siberian Trap, the theory of deeply frozen soils at that timemust be regarded with caution (Korte et al., 2010). On the otherhand, in southern high-latitudes Permian permafrost palaeosolsare proven (e.g., Retallack, 1999b) and a distinct warming wouldhave resulted in methane release that could potentially explainthe extremely low carbon-isotope values in Triassic palaeosols(e.g., Krull et al., 2000; Retallack et al., 2005; Retallack andJahren, 2008).


5. Summary

The negative d13C excursion that commences about500,000 years prior to the P–T boundary is interrupted by twoshort-term positive events and the most pronounced of these startsnear the latest Permian main extinction event. A first carbon-iso-tope minimum occurred at about the P–T boundary followed bysubsequent slight increase and a second (occasionally two-peaked), minimum in the lower (and middle) I. isarcica Zone. Thenegative d13C excursion is most probably due to a combination ofcauses and was most likely triggered by a combined effect of theSiberian Trap volcanism and the flow of anoxic deep waters to veryshallow sea levels. Short-term events, such as release of isotopi-cally light methane from the ocean or permafrost soils, or the massextinction itself, are questionable as causes for the carbon-isotopeexcursions.

Acknowledgements

We acknowledge Ethan Grossman (College Station, Texas),Charles Henderson (Calgary) and Yukio Isozaki (Tokyo) for the re-views and pertinent comments. We thank Christoph Heubeck (Ber-lin), Hugh Jenkyns (Oxford), Graham Shields (London) and JánVeizer (Ottawa) for discussion and English corrections, Peter Jones(Canberra) and Charles Henderson (Calgary) for stratigraphic infor-mation for Australian successions and Shangsi (South China),respectively, Gerhard Stampfli and Caroline Wilhem (both Lau-sanne) for palaeogeographic information, Guo Qingjun (Münster)for translation/transcription of details for publications in Chineselanguage and the Freie Universität Berlin for contributions tofinancing this project and for providing the necessary facilities.

Appendix A. Supplementary material

Supplementary information associated with this article can befound, in the online version, at doi:10.1016/j.jseaes.2010.01.005.

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carbon-isotope stratigraphy across the permian–triassic boundary: a review

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