massive volcanism at the permian–triassic boundary and its impact on the isotopic composition of...

Journal of Asian Earth Sciences 37 (2010) 293–311

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Journal of Asian Earth Sciences

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Massive volcanism at the Permian–Triassic boundary and its impacton the isotopic composition of the ocean and atmosphere

Christoph Korte a,*, Prabhas Pande b, P. Kalia b, Heinz W. Kozur c, Michael M. Joachimski d, Hedi Oberhänsli e

a Institut für Geologische Wissenschaften, Freie Universität Berlin, Malteserstr. 74-100, 12249 Berlin, Germanyb Department of Geology, University of Delhi, Delhi 110 007, Indiac Rézsü u. 83, H-1029 Budapest, Hungaryd Institut für Geologie und Mineralogie, Universität Erlangen, Schlossgarten 5, 91054 Erlangen, Germanye Geo Forschungs Zentrum Potsdam, Section 3.3, Telegraphenberg, 14473 Potsdam, Germany

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

Article history:Received 24 October 2007Received in revised form 31 July 2009Accepted 2 August 2009

Keywords:Carbon isotopesOxygen isotopesPermian–Triassic boundaryGuryul RavineAbadehPufels/Bula/BullaVolcanism

1367-9120/$ - see front matter � 2009 Elsevier Ltd. Adoi:10.1016/j.jseaes.2009.08.012

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

Bulk carbonate and conodonts from three Permian–Triassic (P–T) boundary sections at Guryul Ravine(Kashmir), Abadeh (central Iran) and Pufels/Bula/Bulla (Italy) were investigated for d13C and d18O. Car-bon isotope data highlight environmental changes across the P–T boundary and show the followingfeatures: (1) a gradual decrease of �4‰ to more than 7‰ starting in the Late Permian (Changhsingian)C. bachmanni Zone, with two superimposed transient positive excursions in the C. meishanensis–H.praeparvus and the M. ultima–S. ? mostleri Zones; (2) two d13C minima, the first at the P–T boundaryand a higher, occasionally double-minimum in the lower I. isarcica Zone. It is unlikely that the short-lived phenomena, such as a breakdown in biological productivity due to catastrophic mass extinction,a sudden release of oceanic methane hydrates or meteorite impact(s), could have been the main con-trol on the latest Permian carbon isotope curve because of its prolonged (0.5 Ma) duration, gradualdecrease and the existence of a >1‰ positive shift at the main extinction horizon. The P–T boundaryd13C trend matches in time and magnitude the eruption of the Siberian Traps and other contempora-neous volcanism, suggesting that volcanogenic effects, such as outgassed CO2 from volcanismand, even more, thermal metamorphism of organic-rich sediments, as the likely cause of the negativetrend.

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1. Introduction

The most severe extinction of the Phanerozoic took place nearthe Palaeozoic–Mesozoic boundary and affected both marine andcontinental biota (e.g. Schindewolf, 1953; Sepkoski, 1989; Raup,1991; Erwin, 1993; Kozur, 1998a,b, 2007). The crisis was a conse-quence of dramatic environmental changes during the latest Perm-ian, including strong perturbations of the Earth’s carbon cycle, asreflected in major negative carbon-isotope excursions reportedfrom marine and lacustrine carbonates (e.g. Chen et al., 1984,1991; Holser and Magaritz, 1987; Magaritz et al., 1988; Baudet al., 1989; Holser et al., 1989; Magaritz, 1989; Oberhänsli et al.,1989; Jin et al., 2000; Xu and Yan, 1993; Dolenec et al., 2001;Heydari et al., 2000, 2001; Korte et al., 2004a,b,c, 2005a; Payneet al., 2004; Korte and Kozur, 2005a,b, Submitted for publication;Algeo et al., 2007a,b; Horacek et al., 2007a,b) and from marineand terrestrial organic material (e.g. Wang et al., 1994; Morante,

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1996; Wignall et al., 1998; Krull and Retallack, 2000; Krull et al.,2000; Twitchett et al., 2001; Musashi et al., 2001; de Wit et al.,2002; Sephton et al., 2002; Thomas et al., 2004; Hansen, 2006;Coney et al., 2007; Riccardi et al., 2007; Grasby and Beauchamp,2008). The causes of the pronounced negative P–T boundarycarbon isotope excursion are still under discussion. The coevalseawater strontium (e.g. Veizer and Compston, 1974; Burke et al.,1982; Koepnick et al., 1990; Martin and Macdougall, 1995; Korteet al., 2003, 2004a) and sulphur isotopic ratios (e.g. Holserand Kaplan, 1966; Claypool et al., 1980; Cortecci et al., 1981;Kampschulte and Strauss, 2004; Korte et al., 2004a; Newtonet al., 2004) also show significant changes, rising rapidly acrossthe P–T boundary.

Here we present new carbon isotope values for bulk carbonateacross the P–T boundary sections at Guryul Ravine (Kashmir) andAbadeh (central Iran) and new oxygen isotope data for conodontsfor the Pufels/Bula/Bulla section in the Southern Alps, Italy. Thisenables us to propose a detailed carbon-isotope trend for the LatePermian–earliest Triassic interval and to discuss the causes for thenegative carbon isotope excursion.

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294 C. Korte et al. / Journal of Asian Earth Sciences 37 (2010) 293–311

2. Geological setting and stratigraphic background

2.1. Guryul Ravine (Kashmir, India)

The Guryul Ravine section (Kashmir, India; Fig. 1), located16 km southeast of Srinagar near the Khunamuh village, comprisesmarine strata above the Gangamopteris shales underlain by Panjalvolcanics. The succession is stratigraphically well constrained byearlier palaeontological and sedimentological studies (Nakazawaet al., 1975; Sweet, 1970, 1973; Teichert et al., 1970; Furnishet al., 1973; Matsuda, 1981, 1982, 1983, 1984; Kalia and Pande,1987; Pande and Kalia, 1994; Kapoor, 1996). The Zewan Formation(Fig. 2), Late Permian in age, consists of marine calcareous sand-stone, sandy claystone, and sandy limestone. The succession be-comes increasingly calcareous up-section and the succeedingKhunamuh Formation consists of alternations of grey to blacklimestone intercalated with black shale layers. In the present paperwe adopt the lithostratigraphic concept proposed by Nakazawaet al. (1975) subdividing Zewan Formation into four (A–D) andKhunamuh Formation into six (E–J) lithounits. Fig. 2 summarizesthe litho- and biostratigraphy of the upper part of the Zewan For-mation and lower part of the Khunamuh Formation.

Four faunal divisions have been recognized in the Permian Ze-wan Formation and basal 2.6 m of the Khunamuh Formation (Nak-azawa et al., 1975). Faunal divisions I and II correspond tomembers A and B of the Zewan Formation, respectively. They arecharacterized by the common occurrence of bryozoans, foramini-fers and brachiopods. According to Nakazawa et al. (1975) divisionI is comparable on palaeontological basis to the Kalabagh Memberof Wargal Formation, Salt Range.

The fossil assemblage of Division II (Fig. 2) though less diversethan in I, is marked by the presence of brachiopods that suggestcorrelation with the Middle and Upper Productus Limestone ofthe Salt Range. The most important fossils are the brachiopods Clei-othyridina cf. subexpansa (Waagen), Neospirifer musakheyleylensis(Davidson), Spiriferella rajah (Saltar), Leptodus nobilis (Waagen),Waagenoconcha cf. abichi (Waagen), Lissochonetes cf. morahensis(Waagen), Linoproductus cf. lineatus (Waagen), Costiferina indica(Waagen), Marginifera himalayensis (Diener), the bivalves Aviculo-pecten (Etheripecten) cf. hiemalis (Salter), and the coral Euryphyllumcainodon (Koker).

Fig. 1. Palaeogeographic map of the Late Permian at 260 Ma (modified afterStampfli and Borel, 2002; Korte et al., 2008). Sampled regions: (1) Guryul Ravine(Kashmir), (2) Abadeh (central Iran), (3) Pufels/Bula/Bulla (Italy). The palaeogeog-raphy at the P–T boundary is similar to that at 260 Ma (see Stampfli and Kozur,2006).

Division III encompasses Member C and D and is characterizedby the predominance of gastropods and bivalves over brachiopods.Cyclolobus walkeri (Diener) was found at the bottom and top ofMember C in association with Hindeodus typicalis (Sweet) andClarkina ex gr. C. carinata (Clark) strongly suggesting a ‘‘Chhidruan”(= middle Dzhulfian, middle Wuchiapingian) age for Member C.Member D can be correlated with the uppermost Dzhulfian (upper-most Wuchipingian) and Dorashamian (= Changhsingian) stages(Kalia and Pande, 1987). Division III is characterized by a diversefossil assemblage represented by the foraminifers Globivalvulinacyprica (Reichel), Lunucammina postcarbonica (Spandel), Glomospirasimplex (Harlton), Pyrulinoides zewanensis (Kalia and Sharma), thebryozoans Dyscritella tenuirama (Crockford), Stenodiscus cf. chaetet-iformis (Waagen and Wentzel), the brachiopods Waagenoconchapurdoni (Waagen), Lissochonetes bipartita (Waagen), L. cf. morahen-sis (Waagen), Linoproductus lineatus (Waagen), M. himalayensis(Diener), the gastropods Bellerophon (B.) branfordianus (Waagen),Retispira ornatissima (Waagen), R. kattaensis (Waagen), and the bi-valves Permophorus ‘‘cf. subovalis‘‘ (Waagen), Etheripecten aff. hie-malis (Salter), Cyrtorostra aff. lunwalensis (Reed).

The faunal Divison IV comprises the basal 2.6 m of the Khuna-muh Formation (Unit E1). This faunal zone occurs in an alternationof black shales and limestone that suggest a change to deeperwater. Transiently less oxygenated conditions are also reflectedin the foraminiferal population. These beds overlie shallow marine,thick-bedded calcareous and muddy sandstones characterized byslump, nodular and convolute bedding. The following fossils werefound in Division IV by Nakazawa et al. (1975): The foraminiferNodosinella longissima (Miklucho-Maklay), the brachiopods Derbyiasp., Dielasma ? sp., Linoproductus cf. lineatus (Waagen), Lissochone-tes morahensis (Waagen), M. himalayensis (Diener), Neospirifer sp.,Pustula sp., Schellwienella sp., W. purdoni (Waagen), the bivalvesClaraia bioni (Nakazawa), Cyrtorostra aff. lunwalensis (Reed), Etheri-pecten haydeni (Nakazawa), ‘‘Palaeolima” middlemissi (Nakazawa).Of special stratigraphic importance are the conodonts with H. typ-icalis (Sweet), H. praeparvus Kozur and Clarkina carinata (Clark), atypical fauna of the latest Changhsingian C. meishanensis–H. prae-parvus Zone of middle and high-latitudes. C. carinata is not yetpresent at this level at low-latitude sites.

Beginning with Member E2, ammonoids and conodonts arecommon in the Khunamuh Formation, indicating a greater waterdepth. The conodont faunas (Matsuda, 1981, 1982, 1983, 1984) en-able detailed correlation with the Tethys Realm (Fig. 2).

The conodont-association with H. typicalis and C. carinata in theupper third of Member D and the common presence of Otoceraswoodwardi (Griesbach) in Member E2 indicate that Kashmir duringthe latest Permian and earliest Triassic belonged to the mid-lati-tude cool-water Perigondwanan faunal province, rather than tothe warm-water Tethyan Province.

2.2. Abadeh (central Iran)

During the Late Permian, Abadeh was located on the northernNeotethyan shelf at the southern margin of the Sanandaj-Sirjanblock, which was a part of the Cimmerian microcontinent(Fig. 1). The P–T boundary succession in this locality has beenintensively investigated biostratigraphically (Taraz, 1971, 1973;Taraz et al., 1981; Kozur et al., 1975, 1978) and recently preciselydefined by conodonts (Kozur, 2004a, 2005, 2007). In summary,the complete biostratigraphic subdivision of the discussed timespan consists of the late Dzhulfian (late Wuchiapingian) Clarkinatranscaucasica Zone, C. orientalis–C. mediconstricta Zone, C. inflectaZone, C. longicuspidata Zone, followed in ascending order by theDorashamian (Changhsingian) C. hambastensis Zone, C. subcarinataZone, C. bachmanni Zone, C. nodosa Zone, C. changxingensis–C.deflecta Zone, C. zhangi Zone, C. iranica Zone, C. hauschkei Zone,

Fig. 2. Guryul Ravine section, showing biostratigraphy, sample locations and carbon isotope values for bulk carbonates. P–V: Paranorites–Vishunites Zone; O–G: Otoceras–Glyptophiceras Zone; C–E: Claraia cf. griesbachi–Eumorphotis multiformis Zone; Ev–Eb: Eumorphotis venetiana–E. aff. bokharica Zone; Cb–Eh: Claraia bioni–Etheripecten haydeniZone; Stars: occurrence of Cyclolobus walkeri: (1) Furnish et al. (1973), (2) Kalia and Pande (1987). The conodont genus Neoclarkina was introduced by Henderson and Mei(2007).

C. Korte et al. / Journal of Asian Earth Sciences 37 (2010) 293–311 295

C. meishanensis–Hindeodus praeparvus Zone, Merrillina ultima–Ste-panovites ? mostleri Zone and the Early Triassic H. parvus Zone,Isarcicella isarcica Zone and H. postparvus Zone of the Brahmanian(Induan). This biostratigraphic subdivision is utilised here to corre-late the carbon-isotope trend (Fig. 3). Most of the zonal index formswere described earlier in Kozur (2004a) and this remains the onlydetailed Tethyan conodont zonation of the interval. The coevalSouth Chinese successions, sometimes erroneously assigned tothe Tethys, were deposited on intraplatform basins that were proneto endemism characterized by the absence of nearly all openmarine ammonoid guide forms from the Tethys. South Chineseconodonts are only slightly affected by endemism. C. iranica (Kozur)and C. nodosa (Kozur) are absent here, but common throughoutthe entire central Tethys over a distance of about 1500 km fromcentral Iran to Transcaucasia. However, most of the open marineChanghsingian (Dorashamian) conodont guide forms (Fig. 4), eventhe index species C. bachmanni (Kozur) and M. ultima (Kozur), arepresent in South China which facilitates the correlation of theChanghsingian GSSP at Meishan with the Tethyan sequence.

Three sections have been investigated at Abadeh (Fig. 3). Car-bon, oxygen and sulphur isotope data from one locality (sectionI) were reported in Korte et al. (2004a) and 87Sr/86Sr ratios of con-odonts are available in Korte et al. (2003) and Korte et al. (2004a).Because the biostratigraphic studies were completed only subse-quently, we incorporate here those isotope data into the newly

developed biostratigraphy (Fig. 3). The new carbon isotope datafor the neighbouring sections, Abadeh II-a and Abadeh II-c, withslightly different lithology, are also presented.

2.3. Pufels/Bula/Bulla (Southern Alps, Italy)

The Southern Alpine P–T boundary section at Pufels/Bula/Bullawas palaeogeographically situated at the outermost western mar-gin of the Palaeotethys near the equator (Fig. 1). The uppermostpart of the Late Permian Bellerophon Limestone Formation of theSouthern Alps is overlain by the Werfen Formation, beginning withoolitic limestones and marls of the Tesero Oolite Horizon (TOH),followed by the Mazzin Member, the Andraz Member, the SiusiMember and the Campil Member. A transitional facies is presentbetween Siusi and Campil Members (Fig. 5).

The Pufels section was biostratigraphically investigated byHuckriede (1958), Staesche (1964) and Kozur and Mostler (in Mos-tler, 1982) using conodonts. A further refinement was proposed byPerri (1991), Farabegoli and Perri (1998) and Farabegoli et al.(2007). The H. praeparvus, H. parvus and I. isarcica Zones are presentin the sequence (Figs. 4 and 5). Sample-numbers follow Perri(1991) and Farabegoli and Perri (1998). Some additional samplesare labeled with new numbers. The base of the Triassic suggestedby Perri (1991) and Farabegoli and Perri (1998) has been slightlymodified by Korte and Kozur (2005a) who propose that the

Fig. 3. Stratigraphic sections for Abadeh I and Abadeh II, showing sample positions and carbon, oxygen and strontium-isotope data from this study and from Korte et al.(2003, 2004a).

Fig. 4. Biostratigraphy of marine and continental sequences in the Changhsingian and in the lowermost Triassic of Iran, South China and Southern Alps (conodonts),Transcaucasia and Iran (ammonoids, the two Otoceras zones are from Perigondwana), Germanic Basin, Tunguska Basin + Taimyr and Dalongkou (all conchostracans). Not toscale. Metres in the Dalongkou column are metres above the base of the Guodikeng Formation in the southern limb of the anticline, measured by Dr. Spencer Lucas(Albuquerque) during the American–Chinese project in Dalongkou, sponsored by the NGS, USA. *Suchonella is a typical Middle and Upper Permian ostracod genus. Tut.H. = Tutonchana Horizon/svita. Gagary-ostrov H./s. = Gagary-ostrov Horizon/svita. Further data for Dalongkou: 171.2 m above the base of Guodikeng Fm.: FAD of Falsiscapostera. 111–132 m above the base of Guodikeng Fm.: level of lower and upper boundary of the Falsisca turaica–F. zavjalovi Zone. The Tethyan Changhsingian conodontzonation was developed in ocean-facing settings of central and northwestern Iran and Transcaucasia by Kozur (2005), and can be applied from central Iran throughnorthwestern Iran, Transcaucasia to Sicily (there not all levels are present in the blocks).


H. parvus Zone begins somewhat below sample BU-17 (Korte andKozur, 2005a), considerably later than assumed by Farabegoliand Perri (1998).

We have defined the base of the Olenekian by the first appear-ance datum (FAD) of Pachycladina obliqua and by magnetostrati-graphic data from Scholger et al. (2000) that were re-dated byKorte and Kozur (2005a) and Korte et al. (2007). The base is pro-posed to be near sample BU 45 (Fig. 5) by Perri (1991) and Farab-

egoli and Perri (1998), while Korte and Kozur (2005a) have placedit somewhat lower, based on a few broken Sc elements of Ellisoniaagordina (Perri and Andraghetti) that were earlier erroneously re-garded as Sc elements of P. obliqua (Staesche). Later studies haveshown that only E. agordina is present below BU 45 and that P. obli-qua occurs in sample BU 45, as already shown by Perri (1991 andsubsequent papers). In Korte et al. (2007) we regarded sampleBU 45 to be at the base of the Olenekian. Orchard (2007), on the

Fig. 5. P–T boundary section at Pufels/Bula/Bulla, d13C values for bulk carbonates and oxygen isotope values for conodont phosphates (lithology slightly modified after Perri,1991; Farabegoli and Perri, 1998; Scholger et al., 2000; (1) FOD of I. isarcica in the Pufels sections according to Farabegoli and Perri, 1998; (2) assumed base of the I. isarcicaZone based on the carbon isotope data (see also Korte and Kozur, Submitted for publication).



other hand, referring to Perri and Andraghetti (1987), assumed thatP. obliqua occurs in Pufels/Bula/Bulla in the upper Gandarian (Die-nerian) and Olenekian. This view cannot be confirmed because P.obliqua occurs, according to Perri and Andraghetti (1987) and Perri(1991), only in sample BU 45. The conodont fauna of one samplecannot belong to both the upper Dienerian and Olenekian withoutfaunal mixing.

3. Material and methods

Carbon and oxygen isotope analyses on finely ground bulk car-bonates from the Guryul Ravine section were performed at theInstitute of Earth Sciences, University of Lausanne, Switzerland.Aliquots of 3–6 mg of 39 samples were reacted offline overnightwith 100% H3PO4 at 50 �C and the generated CO2 was analyzedfor d18O and d13C (Table 1) on a Finnigan MAT 251 mass-spectrom-eter, calibrated against PDB. External precision was 0.08‰ and0.15‰ for d13C and d18O values, respectively.

Sixteen of these samples were re-analyzed for their carbon andoxygen isotope values at the Department of Earth Sciences at theUniversity of Oxford. All data are shown in Fig. 2 and are listedin Table 1. At Oxford about 2 mg of fine-grained whole rocks car-bonate were analyzed isotopically for d18O and d13C using a VG Iso-gas Prism II mass spectrometer with an on-line VG Isocarbcommon acid-bath preparation system. In the instrument the pow-dered sample is reacted with purified phosphoric acid (H3PO4) at90 �C. Calibration to V-PDB standard via NBS-19 is made dailyusing the Oxford in-house (NOCZ) Carrara Marble standard. Repro-

Table 1Analytical data.

Samples (Guryul Ravine) d13C (V-PDB) Lab d13C (V-PDB) Lausanne

B 02-1 1.64 LB 02-2 �0.23 LB 03-3 1.87 LC 01 �2.12 LC 02 2.85 LC 03 0.20 LC 04 1.84 LC 04 1.60 LC 05 2.66 LC 06 �0.17 LC 07 3.41 LC 08 3.53 LC 09 3.47 LD 03 3.00 LD 04 2.27 LD 05 2.90 O 2.76D 07 1.06 O 0.53D 08 0.37 O 1.23D 09 1.24 LD 10 0.60 O 0.59D 11 1.42 LD 12 1.86 O 0.93D 13 �0.21 O �0.42D 14 1.39 LD 15 1.77 LD 16-1 1.79 LD 16-2 1.13 O 0.74D 17 1.12 O 0.88E 01-2 �0.83 O �0.12E 01-4 �0.19 O �0.30E 02-2 �1.29 O �1.40E 02-4 �2.57 LE 02-5 �1.83 O �1.87E 02-7 �2.80 LE 02-9 �3.24 O �3.37E 02-12 �2.30 O �2.37E 02-13 �3.17 O �3.38E 02-16 �2.54 O �2.54F 3 �0.46 L

ducibility of replicated standards is typically better than 0.1‰ ford13C and d18O.

Samples of about 0.15–0.45 mg size, drilled from fresh surfacesof bulk rock carbonates from the sections Abadeh II-a and II-c, wereflushed with helium, reacted with H3PO4 in 10 mL borosilicate Exe-tainers and analyzed isotopically for d13C and d18O at the Institutfür Geologie und Paläontologie of the Universität Innsbruck usinga GasBench II linked to a ThermoFinnigan DeltaplusXL mass spec-trometer. For further details see Spötl and Vennemann (2003). Iso-topic data (Table 1) are reported in delta notation relative to theVienna Pee Dee Belemnite (V-PDB) international scale. Externalprecision was better than 0.1‰ for d13C and d18O.

Conodont-samples (about 1 mg) from Pufels/Bula/Bulla weredissolved in nitric acid and chemically converted to Ag3PO4 usinga method slightly modified from that of O’Neil et al. (1994). Oxygenisotope ratios were analyzed on CO generated by reducing trisilver-phosphate using a high-temperature conversion-elemental ana-lyser (TC-EA) connected on-line to a ThermoFinnigan Delta plusmass spectrometer at the Institut für Geologie und Mineralogie,Universität Erlangen (Joachimski et al., 2004, 2006). Samples andstandards were run in triplicate. All phosphate d18O values are re-ported in permil relative to V-SMOW (Table 1). Accuracy andreproducibility were monitored by multiple analyses of trisilver-phosphate prepared from NBS120c and several trisilverphosphatereference samples received from other isotope laboratories. Theoverall reproducibility of the apatite oxygen isotope analysis wasbetter than ±0.25‰ (1r). The mean d18O value of NBS120c was22.4‰ V-SMOW which is relatively close to the value of 22.58‰

Samples (Abadeh) d13C (V-PDB) Samples (Pfufels) d18O (V-SMOW)

Aba 120 3.61 BU 38 18.24Aba 121 3.62 BU 39 18.25Aba 122 3.68 BU 40NA 17.61Aba 123 3.84 BU 40NB 18.39Aba 124 3.84 BU 41 18.47Aba 125 3.79 BU 42 18.45Aba 126 3.88 BU 48 17.76Aba 127 3.84Aba 129 3.58Aba 130 3.57Aba 131 3.37Aba 133 3.70Aba 132 3.19Aba 134 3.58Aba 135 3.73Aba 136 3.40Aba 137 3.15Aba 138 2.63Aba 139 2.64Aba 140 2.41Aba 141 2.27Aba 142 2.16Aba 143 2.15Aba 144 1.99Aba 145 1.74Aba 146 1.82Aba 147 1.80Aba 148 1.63Aba 149 1.89Aba 150 1.73Aba 151 1.35Aba 153 1.34Aba 154 0.35Aba 154 a 0.57Aba 155 0.45Aba 156 �0.06

Fig. 6. Compiled oxygen isotope trend for the P–T transition for equatorial sections. Brachiopods are from Jolfa (NW Iran) and Peitlerkofel/Sas de Pütia/Sass de Putia(Southern Alps, Italy) and conodonts from Abadeh (central Iran) and Pufels/Bula/Bulla (Southern Alps, Italy).


V-SMOW measured by conventional fluorination with BrF5

(Vennemann et al., 2002).In order to compare phosphate (SMOW) and carbonate (PDB)

d18O, different oxygen isotope fractionation factors for carbonatesand phosphates must be taken into account whereby the offset be-tween these two phases is approximately 8.5‰ at temperaturesbetween 20 and 30 �C (Longinelli and Nuti, 1973; Bryant et al.,1996; Iacumin et al., 1996; Joachimski et al., 2004). These differ-ences have been incorporated in Fig. 6.

4. Results

4.1. Guryul Ravine

Carbon isotope values from member B-2 and B-3 and from thelower member C range between �3 and +2‰ (Fig. 2). Because ofthe extreme variations and the differences with the general Dzhul-fian d13C-curve (Fig. 3; see also Korte and Kozur, Submitted forpublication), we believe that these carbon isotope values represent,at least in part, diagenetically altered signals, an assumption that isstrongly supported by the petrology. The samples consist of calcar-eous sandstone, sandy claystones, and sandy limestone (see chap-ter 2) which make isotope values more prone to diageneticalteration and weathering processes (e.g. Veizer, 1983; Marshall,1992). Alternatively, the 13C-depleted carbon isotope values mayreflect the burial diagenetic environment during cementation.The succession becomes increasingly calcareous up-section andcarbon isotope values are therefore more reliable in the upper thirdof member C and in member D and become most pristine in mem-bers E-1, E-2, E-3 and F. The d13C values are about 3.5‰ in theupper third of member C and in the lower part of member D,decreasing in the middle part of member D to values of about1‰ and decreasing more-or-less gradually from the upper thirdof member D – interrupted by a short-term positive excursionstarting near the Late Permian event horizon (EH) – to a first min-imum of �2.6‰ at the base of the H. parvus Zone (P–T boundary).The values increase somewhat to �1.9‰ followed by a second,more prolonged, two-peaked minimum below �3‰ in the I. isarci-ca Zone (Fig. 2).

The bulk carbonate carbon isotope values of Baud et al. (1996)and Atudorei (1999), recently re-published with conodont rangesof Algeo et al. (2007b), show the same decreasing trend with theshort positive excursion mentioned above starting near the EH, afirst minimum at the P–T boundary (no data for the lower I. isarcicaZone are reported there), a minimum in the middle/late I. isarcicaZone, and are similar to our data (Fig. 2). We note, however, thatthe previously published data are more depleted in 13C in the high-er member D and in the kummeli and dieneri Zones compared toother P–T boundary sections (Korte and Kozur, Submitted forpublication).

4.2. Abadeh

The carbon isotope curve of the Abadeh section II (a and c) –similar to that of the main section (Abadeh I) – starts at 3.5‰ inthe late C. transcaucasica Zone (Dzhulfian), increases to about 4‰

for the most of the C. orientalis–C. mediconstricta Zone and variesbetween 3.3‰ and 3.9‰ in the C. inflecta, C. longicuspidata, C. ham-bastensis and C. subcarinata zones (Fig. 3). The gradual d13C declinestarts in the early C. bachmanni Zone, continues throughout the en-tire upper Dorashamian (upper Changhsingian) and culminateswith the lowest value (�0.1‰) at the P–T boundary.

Korte et al. (2004a) reported data for the upper Dorashamianfrom the main section (Abadeh I) showing two short-term rever-sals that interrupt the main negative trend (Fig. 3), with a posi-tive-excursion of 1.5‰ within the Boundary Clay (C.meishanensis–H. praeparvus Zone, starting near the EH) and anotherpositive shift in the M. ultima–S. ? mostleri Zone.

Conodont oxygen-and strontium-isotope trends in Fig. 3 arefrom Korte et al. (2003,2004a) and will be discussed later.

4.3. Pufels/Bula/Bulla

Carbon isotope data for the P–T boundary section at Pufels havebeen published already by Korte and Kozur (2005a) for the P–Ttransition interval and by Korte et al. (2005a) for the upper partof the section (Fig. 5). Recently Horacek et al. (2007a) reportedd13C data for the same section and time interval and these have


been incorporated into our compilation (Fig. 5). Note that themember-subdivided data of Horacek et al. (2007a) have been takenfrom their Fig. 2. Isotope values and trends from the two datasetsare similar, though a bed-by-bed correlation is not possible, illus-trating the potential for chemostratigraphic correlation (Fig. 5).

Carbon isotope values in the uppermost Bellerophon LimestoneFm are about 3.5‰. The curve descends gradually – with two �1‰

positive excursions (one in the H. praeparvus fauna of the lowerTesero Horizon and one somewhat above the top of the TeseroHorizon, the latter in a conodont-free interval between the cono-dont-proven H. praeparvus Zone and the H. parvus Zone) – to low-est values of about �2.8‰ in the lower Mazzin Member. This d13Cminimum is situated about 12 m above the base of the Tesero Oo-lite Horizon and about 1.3 m below sample BU-17. This level wasproposed by Korte and Kozur (2005a) to represent the P–T bound-ary at the Pufels-section (see also Korte and Kozur, Submitted forpublication). Such assignment differs from Farabegoli and Perri(1998) and Horacek et al. (2007a) who placed the P–T boundaryin the lower third of the TOH at sample BU 12B, about 1.4 m abovethe base of the Tesero Oolite Horizon. In this regard, we note thatthe P–T boundary level of Horacek et al. (2007a) lies at the top of ashort, reversed magnetostratigraphic interval (Scholger et al.,2000) whereas the base of the Triassic (base of H. parvus Zone) issituated all over the world within the lower third of a longer nor-mal palaeomagnetic zone (Kozur, 2004b; Yin et al., 2005). For fur-ther discussion see Korte and Kozur (2005a).

The compiled carbon isotope curve (Fig. 5) increases after thefirst (P–T boundary) minimum and returns subsequently to lowvalues in the middle Mazzin Mb, somewhat below sample BU 22and distinctly below the base of the I. isarcica Zone sensu Farabego-li and Perri (1998). We believe that this second d13C-minimum maybe the start of a second, and more prolonged, two-peak-minimumwithin the I. isarcica Zone (see Korte and Kozur, Submitted for pub-lication). At Pufels the first occurrence datum (FOD) of I. isarcica,according to Farabegoli and Perri (1998), is in sample BU-25A,but no sample has been investigated for conodonts in the underly-ing 10 m of the section (Fig. 5). Furthermore, I. isarcica staeschei,which defines the lower I. isarcica Zone, is at this stratigraphic levelmuch less common than H. parvus (ratio 1:10 to 1:100, respec-tively). Therefore, in poor conodont faunas (such as in Pufels) thelower I. isarcica Zone is frequently assigned to the H. parvus Zoneand the correct assignment to the lower I. isarcica Zone can onlybe recognized if very large samples are available (>20 kg). With re-gard to the carbon isotope curve, the FAD of I. isarcica in Pufels(Fig. 5) is distinctly lower than reported by Farabegoli and Perri(1998). This is in agreement with the estimated much longer dura-tion of the I. isarcica Zone compared to that of the H. parvus Zone(Kozur, 2005, 2007).

Oxygen isotope data are from conodonts of the Pufels sectionassigned to the late Brahmanian (Induan) and early Olenekian(Fig. 5). Only two conodont d18O values are known from previousstudies (Early Triassic; Korte et al., 2004a). Together with the cono-dont-data of the present study we now can delineate an oxygenisotope trend across the P–T boundary (Fig. 6). In Fig. 6 we displayall available low-latitude data from our working group, brachio-pods (Korte et al., 2005b) and conodonts (Korte et al., 2004a andpresent study).

Late Permian brachiopods (Jolfa, NW Iran) and conodonts (Aba-deh, central Iran) show similar d18O values and ranges. The d18O-mean-values vary by about 1‰ in the Late Permian and decrease(including the overlapping data) by about 1.5‰ across the P–Tboundary. The conodont-data are slightly higher compared to coe-val brachiopods. Note that the Abadeh succession (Figs. 3 and 6)was deposited close to the Tropic of Capricorn, and Jolfa was situ-ated about 1000 km N of Abadeh (Fig. 1, see also Korte and Kozur,submitted for publication). The Jolfa brachiopods show somewhat

warmer temperatures compared to the Abadeh conodonts (Fig. 6)which is to be expected because of their 1000 km closer proximityto the equator. The sediments from the Southern Alps (Figs. 5 and6) represent shallow marine facies deposited near the equator.

5. Discussion

5.1. The d13C fluctuations at the P–T boundary

High amplitude d13C-fluctuations are generally caused by majorperturbations in the Earth’s carbon cycle (Kump and Arthur, 1999).It is believed that ocean/atmosphere carbon isotope values areprincipally controlled by the burial and re-oxidation of 12C-en-riched organic matter and influenced by a plethora of additionalfactors, such as atmospheric CO2 levels, nutrient supply, sedimen-tation rates, net primary productivity, biological isotope fraction-ation, or sea-level changes (Scholle and Arthur, 1980; Jenkyns,1996; Holser, 1997; Hayes et al., 1999; Kump and Arthur, 1999;Jarvis et al., 2006). For the P–T boundary a major perturbation inthe global carbon cycle involving the atmosphere is indicated bythe fact that the d13C negative excursion is not restricted to seawa-ter, but is reported also for land plant material (e.g. de Wit et al.,2002; Coney et al., 2007) and for lacustrine carbonates of the Ger-manic Basin (Korte and Kozur, 2005b).

Different causes, such as Siberian Trap volcanism (e.g. Renneet al., 1995; Hansen, 2006), re-oxidation of previously stored 12C-enriched organic material due to eustatic sea-level fall (e.g. Baudet al., 1989; Holser and Magaritz, 1992), changes in the proportionof buried organic carbon (e.g. Payne et al., 2004), release of isotopi-cally light methane (e.g. Erwin, 1994; Krull and Retallack, 2000;Krull et al., 2000; Twitchett et al., 2001; de Wit et al., 2002; Sarkaret al., 2003; see also Dickens et al., 1995), a breakdown in oceanicprimary productivity (e.g. Visscher et al., 1996; Rampino and Cal-deira, 2005; see also Hsü and McKenzie, 1985; Kump, 1991), andshallow marine anoxia resulting from an upward rise in thechemocline or ocean overturn (Malkowski et al., 1989; Korteet al., 2004a; Algeo et al., 2007a; see also Kuespert, 1982; Knollet al., 1996; Kump et al., 2005) have all been proposed as potentialcauses of the negative carbon isotope excursion. It has also beenpointed out that multiple causes may have been responsible (e.g.Renne et al., 1995; Berner, 2002; Korte et al., 2004a; Sephtonet al., 2005) for this isotope pattern. Additional discussion is avail-able in Erwin et al. (2002), Kump (2003), Corsetti et al. (2005),Racki and Wignall (2005), Erwin (2006), Isozaki (2007, 2009),Twitchett (2007) and Korte and Kozur (Submitted for publication).

The slope and shape of the negative carbon isotope excursion isimportant for understanding its origin. It is therefore necessary topoint out the d13C-characteristics of the three investigated P–Tboundary sections. The general carbon-isotope trend is marked bythe following features: (1) the Late Permian decrease of between�4‰ and �7‰ is gradual, starting in the C. bachmanni Zone; (2)the decline is interrupted by a transient positive excursion in theC. meishanensis–H. praeparvus Zone and occasionally one small,short-term, positive shift in the M. ultima–S. ? mostleri Zone; (3)the first minimum occurs around the P–T boundary; (4) the secondmore prolonged minimum is situated in the lower I. isarcica Zoneand may be subdivided into two separate events (see also Korteand Kozur, Submitted for publication). Note that lower amplitudesare characteristic of low-latitude pelagic sections (e.g. Abadeh) andlarger shifts more commonly occur in shallow marine successions(e.g. Pufels), higher latitudes (e.g. Guryul Ravine) and intraplatformbasins (Korte and Kozur, Submitted for publication), an observationpointed out already by Krull et al. (2000) and Twitchett et al. (2001).

The gradual decline in the carbon isotope curve, reported al-ready by Magaritz et al. (1988) and Holser et al. (1989), indicates


that the d13C-negative excursion covered an extended time intervalof approximately 0.5 Ma, taking into account the durations of thelate Permian and Early Triassic conodont zones estimated by Kozur(2005). Short-lived events, such as impact related causes, suddenbreakdowns in oceanic primary productivity at the level of massextinction, and sudden releases of oceanic methane hydrates, cantherefore be excluded as causative factors. A worldwide eustaticregression (e.g. Baud et al., 1989; Holser and Magaritz, 1992) thatcould result in enhanced re-oxidation of previously buried 12C-en-riched organic carbon is also an unlikely explanation, despite thefact that some evidence for a wide-spread sea-level fall duringthe late Permian (in the C. iranica Zone) can be found in the shal-lowing horizon of the uppermost Bellerophon Limestone Forma-tion in the eastern Southern Alps. The Tesero Oolite Horizon ofthe basal Werfen Formation in the western parts of the SouthernAlps is a coeval sequence. Contemporary regression took place alsoin S-China, as indicated by the fact that in situ deposits of C. iranicaZone sediments are absent there; this species occurs in S-China inthe Xifanli section where reworked C. iranica are common in the C.meishanensis–H. praeparvus Zone (Kozur, 2005). However, the latePermian sea-level drop had already ceased by the beginning ofthe C. hauschkei Zone (see Yin et al., 2001).

The main extinction event occurred at the ‘event-horizon’, cor-relative to the base of the C. meishanensis–H. praeparvus Zone(= base of the Boundary Clay), but is not associated with a d13C de-crease. Instead, a short-term positive excursion at this stratigraphiclevel occurs at the Guryul Ravine (Fig. 2), Abadeh (Fig. 3) and Pufels(Fig. 5; not expressed by Gorjan et al., 2008, d13Ccarb), and is re-ported also in other sections, such as the Meishan B (China; Nanand Liu, 2004; not expressed in other datasets), Shahreza (Iran;Korte et al., 2004b), Zal (Iran; Korte et al., 2004c), Gartnerkofel core(Austria; Holser et al., 1989), Tesero (Italy; Korte and Kozur, 2005a)and Seis (Italy; Newton et al., 2004; Kraus et al., 2009), suggestingthat it is a general feature. It can be therefore ruled out that a sud-den breakdown of biological productivity, due to mass extinction(e.g. Visscher et al., 1996; Rampino and Caldeira, 2005), causedthe negative carbon isotope excursion.

5.2. Conchostracan stratigraphy of the continental P–T boundarysuccessions and the Siberian traps

The most obvious geological process that operated during thetime span under consideration is the Siberian Trap volcanism,marked by more than 4000 m of volcaniclastic sediments overlainby predominantly flood basalts (Fedorenko et al., 1996; Kamo et al.,2003) and intercalated locally with lacustrine deposits that containfossil conchostracans and plant material.

Geochronological data for Siberian Trap volcanism yield agesthat may implicate it in the mass extinction event (e.g. Renne andBasu, 1991; Campbell et al., 1992; Renne et al., 1995; Kozur,1998a,b; Reichow et al., 2002; Kamo et al., 2003; Courtillot and Ol-son, 2007; Ganino and Arndt, 2009; Svensen et al., 2009). Its surfaceexposure in the Tungusska Basin of around 2.5 million km2, andprobably much larger originally (Kozur, 1998a,b; Vyssotski et al.,2006; Saunders and Reichow, 2007) makes it the largest continentalflood basalt province on the planet (Racki and Wignall, 2005; Saun-ders and Reichow, 2007). A growing consensus argues that volcanicactivity started early, already in the middle Changhsingian, with itsexplosive phase (tuffs) (e.g. Kozur, 1998a). The Siberian Trap mayhave extended all the way to the Urals (Tuzhikova, 1985), were of-ten overlooked because it was erroneously regarded to be of Trias-sic age; for example, in the Timan-Petchora Basin (Tuzhikova,1985). However, Kozur (1998a,b) pointed out that the ‘‘Triassic”sporomorphs of Tuzhikova (1985) are in reality a more-or-less com-plete Lueckisporites palynodeme succession of Late Permian agethat is time equivalent with the upper Zechstein cycles 3–7.

In order to evaluate the influence of Siberian Trap volcanism onthe isotope curves, the stratigraphy of continental deposits byconchostracan zonation (Fig. 4) from the Germanic Basin, Siberia(Tunguska Basin and Taimyr) and China (Dalongkou, Sinki-ang = Xinjiang Province) will be explained and reviewed here.

The conchostracan zonation of the uppermost Permian andLower Triassic of the Germanic Basin was first elaborated by Kozurand Seidel (1983a,b) and later refined and correlated with the mar-ine realm by Kozur (1993, 1999), Kozur and Mock (1993), Bach-mann and Kozur (2004), Kozur and Bachmann (2005) and Kozurand Hauschke (2008), among others. The detailed studies were fo-cused particularly on the P–T boundary. Because conchostracansoccur also in the brackish to brachyhaline marine sediments thatare interbedded with the conodont-bearing Werfen Beds in Hun-gary (Kozur and Mock, 1993; Kozur, 1999), in the Southern Alps,and in brackish intercalations of ammonoid-bearing sediments inSiberia, the base of the Triassic (per definition FAD of H. parvus)can be correlated with the base of the Falsisca verchojanica Zone.F. verchojanica (Molin) occurs also in NE Siberia, in brackish orfresh-water beds that intercalate with marine beds at the base ofthe Triassic, the latter, while devoid of conodonts, contain ammo-noids (Molin in Molin and Novozhilov, 1965; Novozhilov, 1970). Inmarginal parts of the Germanic Basin, both F. postera (Kozur andSeidel) and F. verchojanica (Molin) are common. Ptaszynski andNiedzwiedzki (2004a,b, 2005) documented that the F. verchojanicaZone overlies the F. postera Zone (a very short overlap of the twoindex species occurs at Dalongkou) in the Holy Cross Mountainsof Poland. F. postera is common also towards the top of theF. postera Zone in the conchostracan-rich basinal facies (e.g. thetype locality of F. postera at Nelben section near Könnern, NWof Halle, Germany) but F. verchojanica is absent. Mostly only aFalsisca-free interval is present above the Falsisca-rich topmostF. postera Zone. In somewhat younger beds, F. cf. verchojanica ispresent. As a result, the P–T boundary in the basinal facies mustbe defined by the LAD of F. postera.

The distinct P–T boundary d13C minimum in marine successionsthat is at the base of the H. parvus Zone can also be found in lacus-trine limestones of the Germanic Basin. In these lacustrine depos-its, the P–T boundary is defined by the FAD of F. verchojanica or theLAD of F. postera (Korte and Kozur, 2005b). A revision of the upper-most Permian to Triassic conchostracan faunas by Kozur andWeems (in press) shows that Falsisca podrabineki (Novozhilov) ispresent in and restricted to the F. postera Zone of the Germanic Ba-sin. Here, the Falsisca zavjalovi (Novozhilov) is present in the upperpart of the F. eotriassica Zone. These new results enable correlationof the Germanic Triassic with the Tunguska Basin and the SiberianTraps. In the Germanic Basin the base of the F. eotriassica Zone is animmigration horizon for conchostracans (after a facies controlledtime interval without conchostracans). In the Tunguska Basin, F.zavjalovi is common in the Khungtukun tuffs but no longer presentin sediments between the Putarana flood basalts (Fig. 4). Thismakes it clear that the F. eotriassica Zone of the Germanic Basincannot be any younger than the tuffs below the Puturana flood bas-alts, as confirmed by the presence of F. podrabineki in the sedi-ments intercalated in the Puturana flood basalts. The distinctclimate change at the Zechstein–Buntsandstein boundary, coinci-dent with the base of the Boundary Clay at the base of the C. mei-shanensis–H. praeparvus Zone, is also coincident with the beginningof the wide-spread flood basalt volcanism of the Siberian Traps(Fig. 4).

During the Sino-American NGS project in 1996, palaeomagneticand conchostracan samples were collected in Dalongkou (H.W.Kozur was involved in these studies). A detailed section atthe southern limb of the anticline was measured by S.G. Lucas(Albuquerque). These measurements are used here (differs slightlyfrom Metcalfe et al., 2009). Important bioevents are given in


metres above the base of the Guodikeng Formation (Fig. 4). Despitethe fact that all collected material was confiscated by local author-ities, Kozur managed to identify large or distinctly sculpturedconchostracans (Bipemphigus, Falsisca, Megasitum, Trimpemphigus)in the field. Smaller conchostracan genera, and some pieces ofBipemphigus and Tripemphigus, could only be identified at thegenus level or species groups. The conchostracan succession at Dal-ongkou is evidently identical with the Germanic Basin sections up-wards from the F. eotriassica Zone (Fig. 4). The upper Zechstein ofthe Germanic Basin below the F. eotriassica Zone, contains noconchostracans, but the succession at Dalongkou coincides wellwith the sections of the Tunguska Basin and the base of the Triassicat Dalongkou is therefore precisely defined (Kozur, 1998a,b). Thebase of the Triassic is not, as assumed earlier, situated close tothe FAD of Lystrosaurus (only King and Jenkins, 1997 reportedLystrosaurus from the Permian of Zambia), but close to the LAD ofDicynodon. This latter result was confirmed by Ward et al. (2005)who reported for the Karroo Basin in South Africa that the carbonisotope minimum at the P–T boundary is situated at the LAD ofDicynodon, distinctly above the FAD of Lystrosaurus (see alsoMacLeod et al., 2000; Smith and Ward, 2001). In addition, theFAD of Lystrosaurus lies in the upper part of a palaeomagnetic hori-zon marked by a short reversal interval. This horizon correspondsto the lower Fulda Formation and the lowermost upper Fulda Fmof the upper Zechstein in the Germanic Basin. This means thatthe interval of the overlap of Lystrosaurus and Dicynodon belongsto the latest Permian as shown by Kozur (1998a,b) by conchostr-acan correlation in Dalongkou. In the present publication (Fig. 4),as in Kozur (1998a,b), we place the base of the Triassic in continen-tal successions at the base of the F. verchojanica Zone (in Dalongkou210 m above the base of the Guodikeng Fm and 24.7 m below itstop). This is because the boundary between the F. postera and F.verchojanica Zones is correlated (1) with conodonts in the Werfenbeds (P–T boundary is conodont-defined) and (2) with the mini-mum in d13Ccarb (Korte and Kozur, 2005b) in the Germanic Basin.The LAD of Dicynodon (219 m above the base of the GuodikengFm) lies at Dalongkou, only insignificantly above the FAD of F. ver-chojanica. Considering the extremely high sedimentation rates, the9 m difference represents only a very short time interval. Whenconchostracans are not present, the LAD of Dicynodon can be usedto define the P–T boundary, as in South Africa.

The first occurrence of Lystrosaurus at Dalongkou (Metcalfeet al., 2009) is not the FAD of this genus, but the FOD because it oc-curs within the F. eotriassica Zone in the long palaeomagnetic nor-mal interval that straddles the P–T boundary. As mentioned above,in South Africa, Lystrosaurus begins in the short reversal intervalimmediately below this long palaeomagnetic normal interval (seeWard et al., 2005). The sediments of the short reversed intervalat Dalongkou were either not sampled, not deposited, removedby erosion or any combination of the above. In the critical interval,mudstones occur almost exclusively and were not sampled byMetcalfe et al. (2009). Their view that the palaeomagnetic horizonwith a short reversal in the lowermost Guodikeng Fm and upperWutonggou Fm corresponds to the short reversal immediately be-low the long normal palaeomagnetic interval straddling the P–Tboundary cannot be confirmed, because it contradicts the biostrati-graphic data. The basal Guodikeng and upper Wutonggou forma-tions have conchostracan faunas well below the characteristicBipemphigus–Megasitum–Trimpemphigus conchostracan fauna. Theaforementioned short reversal horizon lies above the Bipemphi-gus–Megasitum–Trimpemphigus fauna and, in this stratigraphic le-vel, Falsisca species are different. In the Germanic Basin, the topof this reversal interval lies in the lowermost F. eotriassica Zone.At Dalongkou, this zone begins more than 100 m above the topof the short reversed interval. Within this short reversed palaeo-magnetic zone, tuff fallouts occur in the Nedubrovo Formation

NE of Moscow and volcanic microspherules are common in thelower Fulda Formation of Germany. This interval contains spo-romorphs of the Triquitrites proratus Zone and the megaspore Oty-nisporites eotriassicus (Fuglewicz). O. eotriassicus is common in theupper Guodikeng Formation, but absent in its entire lower half(Metcalfe et al., 2009). In addition, the top of the short reversalinterval lies about 100 m below the stratigraphically lowest occur-rence of O. eotriassicus. Thus, the palynological data and the conc-hostracan data show the same: the reversed horizon in thelowermost Guodikeng Fm and in the Wutonggou Fm cannot becorrelated with the short reversed horizon immediately belowthe long normal palaeomagnetic zone that straddles the P–Tboundary. The reversed palaeomagnetic horizon of Dalongkouwhich ends within the lowermost Guodikeng Fm corresponds tothe reversed interval (Szurlies, 2007) in the uppermost Zechstein3 and Zechstein 4. This stratigraphic correlation is clearly con-firmed by the Milankovitch cyclicity. The Guodikeng Fm containsfour easily recognizable, short-eccentricity cycles with mainly five,just as easily recognizable, precession cycles, as recognized by Ko-zur and Lucas during the field work. Thus, the duration of the Guo-dikeng Fm is about 400,000 years and the upper limit of thereversed horizon in the Wuttongou and basal Guodikeng forma-tions lies about 400,000 years below the top of the GuodikengFm and more than 350,000 years below the base of the F. verchoja-nica Zone (= base of the Triassic). In the Germanic Basin, the top ofthe short palaeomagnetic reversal lies about 200,000 years belowthe base of the F. verchojanica Zone (Bachmann and Kozur, 2004),and in the conodont-dated beds of central and NW Iran and South-ern Alps it lies about 200,000 years below the base of the Triassic(Kozur, 2007).

Conchostracans described from Dalongkou (Liu, 1989) arerather difficult to evaluate because the taxonomy is partly out-dated. Important genera are assigned to other, unrelated generasuch as the Permian genus Megasitum (Novozhilov), which wasdetermined as the Gandarian (Dienerian) genus Cornia (Ljutke-vich). For other genera, such as Bipemphigus (Novozhilov, 1965)and Tripemphigus (Novozhilov, 1965) the junior synonym Trinodus(Liu, 1982) was established. Falsisca cf. kanandaensis Novozhilovsensu Liu (1989) belongs to Falsisca turaica. Falsisca beijiangensis(Liu, 1982) is a junior synonym of F. zavjalovi (Novozhilov, 1970).

In the Tunguska Basin and on Taimyr Peninsula, the SiberianTrap volcanics are well dated by a rich flora and conchostracan fau-na (e.g. Novozhilov, 1970; Sadovnikov and Orlova, 1993; Sadovni-kov, 1997). Radiometric dating (e.g. Renne and Basu, 1991;Campbell et al., 1992; Renne et al., 1995; Kamo et al., 2003; Sado-vnikov, 2008) of trap volcanics in the Norilsk and the Maymetscha-Kotuy areas suggests a 600,000 years (Kamo et al., 2003) interval ofvolcanic activity. At that time, the single-grain U–Pb zircon analy-sis involving chemical abrasion was not yet refined and the resultsby Kamo et al. (2003) tend to yield slightly younger ages (see Mun-dil et al., 2004). However, this does not influence the duration ofthe main volcanic activity, only the absolute ages. The Kamoet al. (2003) data from the Norilsk area were apparently from thelowermost Triassic part of the Siberian Trap plateau basalts. TheMaymetscha-Kotuy area, which yields younger radiometric ages,has volcanic successions with several larger gaps, rendering diffi-cult any correlation with the volcanic deposits of the Tunguska Ba-sin and Taimyr. Older radiometric data in the Norilsk area are fromsub-volcanic bodies (Sadovnikov, 2008) and therefore ambiguous.

Important are the five fossil-rich horizons (Gagary-ostrov,Tutontchana, Lebedeva, Khungtukun, Putorana; Fig. 4) establishedin the Tunguska Basin (e.g. Sadovnikov and Orlova, 1993; Sadovni-kov, 1997), and the next younger ‘‘horizon” (Marininskii Horizon)in East Taimyr. These ‘‘horizons” or ‘‘svita” are units in rank be-tween formations and groups. The Marininskii Horizon, absent inthe Tunguska Basin but present in East Taimyr, is more than


500 m thick and consists of thick basaltic flows and thinner sedi-mentary or tuffitic intercalations. Falsisca cf. verchojanica and otherconchostracans indicate an earliest Triassic age. The underlyingPutoranian Horizon is the youngest Permian lithostratigraphic Unitof the Siberian Traps in the Tunguska Basin, comprised of lavaflows and some tuffs and shales. The typical (Middle and) UpperPermian fresh-water ostracod genus Suchonella, as well as theconchostracan fauna with Falsisca turaica (Novozhilov) and F. pod-rabineki (Novozhilov), indicate a Permian age. The latter species oc-curs in the Germanic Basin within the F. postera Zone. Therefore,the Putoranian Horizon can also be assigned to the latest PermianF. postera Zone and the succession is similar to Germanic Basin, be-cause both faunas are situated immediately below the F. verchoja-nica Zone. Thus, the P–T boundary evidently lies within the plateaubasalts. The Puturana Horizon is at times used in a wider sense(Orlova, 1999) and includes in its lower part the upper portion ofthe thick tuffs which are otherwise assigned to the upper Khungt-ukun Horizon.

The underlying Khungtukun Horizon consists of several 100 mof tuffs and interbedded sediments that are rich in conchostracans(e.g. Novozhilov, 1970; Sadovnikov and Orlova, 1990, 1993; Sado-vnikov, 1997). As pointed out above, its upper part is sometime as-signed to the lower Puturana Horizon. The uppermost occurrence ofFalsisca zavjalovi is reported in the upper Khungtukun Horizon,where the stratigraphically older Bipemphigus, Megasitum and Trip-emphigus are no longer present. It is quite possible that F. eotriassicais present in the upper or uppermost Khungtukun Horizon, becausethis species is very similar to F. zavljalovi. Its distinguishing featureis only a rounded upper half of the posterior margin against astraight or slightly back-curved upper posterior margin in F. zavjal-ovi. As F. zavjalovi has its LAD at the top of the F. eotriassica Zone andat the top of the Khungtukun Horizon, this unit should reach the F.eotriassica Zone. The occurrence of the sporomorph species Triqui-trites proratus (Balme) in the upper Khungtukun Horizon (Sadovni-kov, 2008) indicates its correlation with the Fulda Fm because T.proratus in the Germanic Basin is restricted to the Fulda Formation.The mafic tuffs in the Nedubrovo Formation NE of Moscov may alsoindicate that the Nedubrovo Formation is certainly no younger thanthe upper Khungtukun Horizon with its thick tuffs, providing theserocks are not assigned to the lower Puturana Horizon. The LowerFulda Fm and Nedubrovo Horizon can be easily correlated by mag-netostratigraphy with the upper Clarkina changxingensis–C. deflectaand most of the overlying C. zhangi conodont zones of Iran (Kozur,2007). The Nedubrovo Formation contains the megaspore assem-blage with Otynisporits eotriassicus (Fuglewicz) which in the South-ern Alps begins close to the base of the Upper Permian Hindeoduspraeparvus Zone and straddles the P–T boundary.

A distinct change in the conchostracan fauna occurs within theKhungtukun Horizon. The lower Khungtukun Horizon has a muchmore diverse fauna with the characteristic Bipemphigus gennisi(Novozhilov), Megasitum harmonicum (Novozhilov), M. vanum(Novozhilov) and Tripemphigus sibiricus (Novozhilov). Falsisca isrepresented by F. turaica (Novozhilov) and F. zavjalovi (Novozhi-lov). The changes in the conchostracan faunas at 107 m abovethe base of the Godikeng Fm in Dalongkou are abrupt and arethe strongest conchostracan turnover in Permian layers. This conc-hostracan turnover is much older than the P–T boundary, mucholder than the event horizon with the strongest extinction eventin marine faunas, and even distinctly older than the FAD of Lystro-saurus. It is not contemporaneous with any distinct faunal changesin the marine realm, not related to the beginning of the tuffitic vol-canism of the Siberian Traps which precede it, and also not coevalwith the beginning of the wide-spread plateau basalt volcanismthat is younger (base of the Puturana Horizon). Any associationof conchostracan turnover can therefore be only with the strongexplosive volcanism (before the Trap basalts). The unusual diverse

Bipemphigus–Megasitum–Tripemphigus fauna and the diverse mac-roflora even in the high-latitude northern Siberia indicate a warmclimate, at least during the summer time. The change to the low-diversity Falsisca–Euestheria fauna in northern Siberia, in the low-latitude Dalongkou fauna, and in the Germanic Basin (Megasitumis present in brackish equivalents of the lower Zechstein, materialHammerich seen by Kozur by courtesy of Mrs. B. Hammerich),would require a distinct drop or rise (or strong fluctuation) in tem-perature across all latitudes, possibly triggered by very strongexplosive volcanic eruptions. During one short interval, theChanghsingian warm-water low-latitude conodont fauna of Iranis replaced three times by a cool-water fauna (Kozur, 2005, 2007)and the warm-water fauna returned in any case, indicating strongtemperature fluctuations.

The Lebedeva Horizon (Fig. 4) comprises mainly tuffs and hasthe same conchostracan fauna as the lower Khungtukun Horizon.This fauna is well correlated with the Megasitum vanum–Tripem-phigus sibiricus–Falsisca turaica Zone at Dalongkou and containsthe same species. The tuffs of the Siberian Traps begin somewhatearlier, in the Tutonchana Horizon that is present only locally,and in the uppermost part of the underlying Gagary-ostrov Hori-zon which lies below the appearance of the Upper Permian Falsiscaspecies. This succession is also present at Dalongkou, where theUpper Permian Falsisca begins 65 m above the base of the Guodik-eng Fm (Fig. 4).

Based on Milankovich cyclicity, the Megasitum vanum–Tripem-phigus sibiricus–Falsisca turaica Zone at Dalongkou is somewhatshorter than 100,000 years and this duration can also be assumedfor the similar fauna of the Lebedeva and the lower KhungtukunHorizons. Accepting the duration for the lower C. changxingensis–C. deflecta and the C. nodosa Zones of the Iranian conodont scheme(Kozur, 2005, 2007), the base of the Lebedeva Horizon should beclose to the base of the C. nodosa Zone. As the tuffs below the Sibe-rian Trap plateau basalts begin not much deeper, in the upper Ga-gary-ostrov Horizon, the beginning of this explosive volcanismwas most probably within the C. bachmanni Zone. Thus, the volca-noclastics below the plateau basalts of the Putorana Horizon corre-spond in the Tethyan conodont zonation to the C. nodosa to C.hauschkei zones (= approx. 330,000 years; Kozur, 2005, 2007), andthe upper C. bachmanni Zone. Evidently the entire duration of vol-canoclastic activity is somewhat less than 400,000 years. As thecontinuous gradual drop in d13Ccarb begins within the C. bachmanniZone, it is contemporaneous with this volcanoclastic sequence(Fig. 7). The commencement of the wide-spread thick plateau bas-alts at the base of the Putorana Horizon coincides with the eventboundary (= event horizon) of the main extinction event in marinebeds and this basalt volcanism straddle the P–T boundary.

A somewhat younger age was reported for the Daldykan intru-sions in the Noril’sk area by Ivanov et al. (2005), but this was pos-sibly a later event and no evidence exists for an uninterruptedvolcanic activity throughout the Early Triassic. Lo et al. (2002) as-sumed that the Emeishan flood magmatism of China was contem-poraneous with the Siberian Trap volcanism, but most authorsconsider it to have occurred around the Guadalupian–Lopinginanboundary (e.g. Chung et al., 1998; Wignall, 2001; Zhou et al.,2002; Isozaki, 2007, 2009), suggesting that the impact of volcanicactivity at the P–T boundary might have been extensive, the moreso that the major pulse of explosive felsic volcanism occurred alsoin the latest Permian C. meishanensis–H. praeparvus Zone in SouthChina. This volcanism was also related to the P–T boundary mainextinction event (Yin et al., 1992; Kozur, 1998a,b, 2007).

5.3. The impact of volcanism on ocean and atmosphere

The stratigraphic review shows that the Siberian Trap volca-nism was larger in size and of longer duration than assumed earlier

Fig. 7. Estimated duration and intensity of mafic Siberian Trap and S-China explosive felsic volcanism across the P–T boundary in comparison to the carbon-isotope trendfrom Abadeh (Korte et al., 2004a). (1) C. nodosa 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. Carbon isotope curve from Abadeh section I. Numeric ages: 252.6 Ma for the event horizon is a calculation by Mundil et al. (2004). The 252.6 Ma for the P–Tboundary was calculated by Kozur (2003) using the 252.5 for the very base of the I. isarcica Zone by Mundil et al. (2001) and the �100,000 a duration for the H. parvus Zone(= one short-eccentricity cycle). The same age was calculated by Mundil et al. (2004) who stated that the age of the biostratigraphic PTB ‘‘is slightly but unresolvably younger”than 252.6 Ma. The 252.4 ± 0.3 Ma is from Bed 25 (lower Boundary Clay) at Meishan (Mundil et al., 2004).


(e.g. Renne et al., 1995; see also Berner, 2002), indicating poten-tially an enormous impact on the global carbon cycle and life onEarth. This appears to be mirrored in the carbon isotope curve(Fig. 7), the more intense the volcanic activity, the lower the carbonisotope values (see also Racki and Wignall, 2005), with the climaxclose to the P–T boundary fitting perfectly with the d13C-minimum.

Several factors likely influenced the carbon isotope values of theocean–atmosphere system. Because outgassed volcanic carbondioxide with a d13C of ��5‰ (but see Hansen, 2006 for a dissent-ing opinion) was probably only 2–3‰ lower (Figs. 2, 3 and 5) thanthe coeval atmospheric-d13C, it is unlikely that this alone couldhave produced a >4‰ negative shift (e.g. Wignall, 2001; Berner,2002) despite enormous quantities of exhaled CO2.

The Siberian Trap eruptions were postulated to have destabi-lized significant quantities of methane hydrate (low d13C values)that was contained in permafrost soils (Dorritie, 2002; Racki andWignall, 2005; Retallack and Jahren, 2008; see also Krull et al.,2000). This may have provided an additional source of 13C-de-pleted carbon. Methane hydrates are reported from present-dayup to 1000 m thick permafrost at high-latitudes (e.g. Dorritie,2007). Because of abundant plant fossils, including ferns and otherhigher plants in the Late Permian Siberian Trap area, it is debatablewhether much soil was actually frozen at that time. Thawing per-mafrost is therefore an unlikely major source for the isotopicallydepleted carbon.

Carbon release by combustion due to Trap volcanism may havesupplied significantly more 13C-depleted CO2 to the atmosphere/hydrosphere system than any of the above mechanisms, with lavaflows, sills, dykes and large sub-volcanic bodies burning organic-rich sediments and coals. Thermogenic release of methane fromcoal beds was previously proposed as a causative factor for the

Toarcian Oceanic Anoxic Event (McElwain et al., 2005; Svensenet al., 2007) and for the Triassic–Jurassic boundary (Korte et al.,2009a; Svensen et al., 2009). Svensen et al. (2004, 2009), Payneand Kump (2007) and Retallack and Jahren (2008) assumed sucha scenario also for the P–T boundary. The quantity of 13C-depletedCO2 can only be speculated at. Svensen et al. (2004) proposed a2.5‰ drop from volcanic baked organic-rich sediments due to injec-tion of Palaeocene–Eocene-age dyke swarms. Considering that theextent of the P–T volcanic events was much larger, the impactshould have been commensurate. Dolerite sills, with thicknessesof up to 1000 m, have been documented to intrude the organic-richand petroleum bearing pre-Trap sedimentary strata here and inadjacent areas (e.g. Meyerhoff, 1980; Kontorovich et al., 1997;Nakashima, 2004; Erwin, 2006; Knoll et al., 2007; Svensen et al.,2009). Such extra source of isotopically depleted CO2 would consid-erably enhance the estimates of Wignall (2001) and Berner (2002)and could potentially be the major cause of the 4–7‰ d13C decline.

Anoxic and dysoxic conditions have been reported for the LatePermian and earliest Triassic Palaeotethys ocean (e.g. SouthernAlps, parts of the Dinarides, NE Iran; Wignall and Twitchett,1996), for marginal seas (W-Australia; Grice et al., 2005), low-lat-itude Panthalassa (Isozaki, 1994, 1997), the western United States(Wignall and Hallam, 1992; Woods and Bottjer, 2000) and Borealseas (Wignall et al., 1998). Isozaki (1997) and Kato et al. (2002)suggested that these conditions existed in deep-sea pelagic oceansfor several million years, from the late Changhsingian to the earlyOlenekian, resulting in superanoxia (see also Wignall and Twitch-ett, 2002); such conditions might have contributed to the latest-Permian mass extinction (Wignall and Twitchett, 1996). However,ocean stratification is unlikely to persist laterally over several1000 km and last for several million years (Zhang et al., 2001;


see also Broecker and Denton, 1989; Hotinski et al., 2001) becausegeothermal heating would cause ocean overturn (Mullarney et al.,2006; Emile-Geay and Madec, 2009). To complicate matters, Kaku-wa (2008) demonstrated that the duration of anoxia for low-lati-tude Panthalassa had been overestimated. This argument isweakened, however, by the fact that Kakuwa (2008) did not mea-sure Pr as it is necessary to know the concentration of this elementin order that anoxia be reliably identified using Ce-anomalies (Katoand Isozaki, 2009). Moreover, at the same time that anoxia pre-vailed in the Boreal seas, low-latitude Panthalassa, Palaeotethysand adjacent seas, oxic conditions existed in the southern high-lat-itude Panthalassa (Arrow Rock, Waipapa Terrane, Northern Island,New Zealand; see Takemura et al., 2003; Yamakita et al., 2003) andin the Neotethys (Kozur, 1998a,b, 2007; Heydari et al., 2003; Ri-choz, 2006). At Jolfa (= Kuh-e-Ali Bashi, NW Iran) anoxic conditionswere proposed for sediments 2 m below the event horizon, withinthe red Paratirolites Limestone, because of a negative cerium anom-aly by Kakuwa and Matsumoto (2006). The Paratirolites beds, how-ever, represent a Hallstatt Limestone facies (ammonitico rosso)that is characterized by a diverse benthos consisting exclusivelyof faunal elements (such as bairdiid-dominated ostracod fauna)that could only live in oxygen-rich bottom water (see Belousova,1965, for the adjacent Dorasham succession) and these fossilassemblages are present also in the Iranian and Transcaucasianrealms. Filter-feeder ostracods (Cavellinidae and Hollinacea) indic-ative of low-oxygen bottom water conditions (common in dysoxicPalaeotethyan sediments) are absent. Consequently, this ceriumanomaly is not an indicator of anoxic water conditions (see alsoKato and Isozaki, 2009).

Wide-spread anoxia across the P–T boundary, despite extendinginto unusually shallow marine waters, was not a worldwide phe-nomenon and it is therefore debatable if global anoxia could havebeen the main cause of mass extinction, at the more so that theextinction happened also at coeval horizons (e.g. ‘event-hori-zon’ = EH = base of the C. meishanensis–H. praeparvus Zone) evenat localities where no anoxia was present (e.g. Jolfa, NW Iran: Kozur,1998a,b, 2005). In such localities the mass extinction took placewithin red sediments containing a rich and diverse benthos. Localeuxinic waters that have reached shallow-marine environmentswith lethal consequences for all shallow-marine organisms (Kajiw-ara et al., 1994; Wignall and Twitchett, 1996; Grice et al., 2005;Meyer and Kump, 2008; Meyer et al., 2008) were likely only a con-tributory factor that may have played a role also in the amplitude ofthe negative carbon isotope excursion (Malkowski et al., 1989; Kor-te et al., 2004a; Wignall et al., 2005; Riccardi et al., 2007; Algeoet al., 2007a, 2008) because anoxic deep water is considerably en-riched in 12C (Deuser, 1970). This scenario of anoxic deep waterupwelling (perhaps by oceanic overturn) or chemocline upwardexcursions (e.g. Kump et al., 2005; Riccardi et al., 2006, 2007; Gra-sby and Beauchamp, 2009) may be supported also by sulphur iso-tope values of seawater sulphate. In anoxic seawater, bacterialsulphate reduction (BSR) results in preferential incorporation ofisotopically light 32S in the resultant H2S, leaving remaining seawa-ter sulphate enriched in 34S (Chambers and Trudinger, 1979; Kaplanand Rittenberg, 1964). At the P–T boundary, extreme d34SCAS (car-bonate associated sulphur) variations of short duration are reportedfrom several sections (Kaiho et al., 2001, 2006a, 2006b; Newtonet al., 2004; Riccardi et al., 2006; Gorjan et al., 2007), althoughwhether they represent a record of seawater is presently debated(Shields et al., 2004; Riccardi et al., 2006; Marenco et al.,2008a,b). It is perhaps pertinent to point out that the long-term in-crease in seawater d34S across the P–T boundary is well establishedfrom the latest Permian to Early Triassic, based on sulphate miner-als (Holser and Kaplan, 1966; Claypool et al., 1980; Cortecci et al.,1981; Wilgus, 1981; Holser and Magaritz, 1987; Chen and Chu,1988), and the indications of a similar trend are also in the d34SCAS

record (Kampschulte and Strauss, 2004; Newton et al., 2004). Therise that commenced already in the Late Permian (Newton et al.,2004; Korte et al., 2009b), and the enhanced BSR, may have beendue either to deep-water anoxia prior to the P–T boundary or to ele-vated bioproductivity caused by enhanced nutrient supply (Korteet al., 2009b; see also Kampschulte et al., 2001). In order to affectseawater d13C, anoxia must have expanded to shallow-water envi-ronments, but such an event was likely of short duration thus not ofmajor impact on the gradual and prolonged carbon isotope negativetrend. In addition, the onset of wide-spread shallow marine anoxiaoccurred around the EH at the base of the C. meishanensis–H. prae-parvus Zone (e.g. Wignall and Twitchett, 2002), lagged the onset ofthe gradual d13C decline by about 400,000 years. This indicates thatits effect could only relate to the latest portion of the negative shift.Only some short-term negative carbon-isotope excursions of <1‰

in the latest Permian (above the EH) and earliest Triassic at NhiTao, coincident with the sulphur isotope trends, could be a resultof anoxia (see Algeo et al., 2008).

5.4. Paleotemperature and weathering

A gradual reduction in biological productivity, possibly due toclimate changes, at first by repeated short-term cooling and acidrain due to SO2 aerosols, and subsequently by long-term globalwarming with very dry climate and strong temperature differencesbetween summer and winter in large parts of Pangea accentuatedby prolonged and largely explosive Siberian Trap volcanism, com-menced about 0.5 Ma prior to the P–T boundary. For example, coaldeposition rapidly declines and finally ends prior to the P–Tboundary. This gradual reduction in bioproductivity might havecontributed to the negative d13C-trend. On the other hand, super-imposed shorter 13C-enriched peaks in the C. meishanensis–H. prae-parvus Zone and M. ultima–S. ? mostleri Zone (Fig. 3) may reflectingnutrient availability (Payne and Kump, 2007; but see Saltzman,2005) or ocean anoxia.

How did seawater temperatures evolve from the Later Permianto the Early Triassic? Consideration of all reliable oxygen isotopedata from brachiopods and conodonts (Fig. 6) suggests an approx-imate 2 �C warming from the late Wuchiapingian (late Dzhulfian)to the late Changhsingian (late Dorashamian), with superimposedintra-Changhsingian and latest Changhsingian short-term cool-ing-periods. For time-intervals with no d18O data available, shortintervals of immigration of cool-water conodonts that replace en-tirely the low-latitude, warm-water ones (Kozur, 2005, 2007) areclearly discernable in the low-latitutde Iranian sections. One inter-val, comprising the upper C. changxingensis–C. deflecta Zone andmost of the C. zhangi Zone, can be precisely correlated with theshort reversed magnetostratigraphic interval of the Germanic Ba-sin (Szurlies, 2007; ‘0r’ according to Bachmann and Kozur, 2004;Kozur and Bachmann, 2005). This time span correlates to the shortmagnetostratigraphic reversed interval in the Nedubrovo Forma-tion of the Russian Platform (Lozovsky et al., 2001; Kozur, 2007)with mafic tuff fallout several 1000 km away from the SiberianTrap eruption centres (Tunguska Basin). Thus, the cool-waterimmigration spikes occurred during a strong explosive eruptionphase of the Siberian Traps (Kozur, 2005, 2007).

No oxygen isotope data from conodonts or brachiopods areavailable for the latest Permian and the earliest Triassic (Korteet al., 2004a), but the long-term trend (Fig. 6) shows a �1‰ de-crease from the late Changhsingian to the Early Triassic, indicatinga warming in equatorial regions of about 4 �C. Note that the Aba-deh and the Pufels conodonts were deposited at likely waterdepths of �10–90 m (Kozur, 2005) and of a few metres to�20 m, respectively. In addition, the Abadeh successions show ashallowing-upward trend, culminating in the Gandarian withwater depths similar to those of the Southern Alps. Because the


water at Pufels and at Abadeh shallowing-upward, this alone couldcause a decline in the d18O values, but a general warming of about4 �C in the equatorial regions is consistent with the propositionthat high-latitude seawater would be more influenced by globalclimatic trends. For example, during the Early Triassic (Brahma-nian = Induan), a low sea-surface pole-to-equator temperature gra-dient is indicated by endemism and latitudinal distribution ofammonoids (Brayard et al., 2006; Galfetti et al., 2007) and by dee-ply weathered Antarctic paleosols at palaeolatitudes of up to 77�S(Retallack and Krull, 1999). This observation requires a consider-able warming in high-latitudes, consistent with the appearanceof low-latitude warm-water biota in Spitsbergen during the earlyGandarian (Wignall et al., 1998). The warming of the Early Triassichas been attributed to increased atmospheric CO2 due to enhancedvolcanism (Wignall and Hallam, 1993) or to release of methanefrom gas hydrates (e.g. Erwin, 1993; Racki and Wignall, 2005).

The Late Permian to Early Triassic 87Sr/86Sr-trend (Fig. 3) ischaracterized by a moderate rise from 0.7070 at the base of theDzhulfian to 0.7082 in the late Olenekian (Korte et al., 2003). Therise commences in the Capitanian and has been attributed to thecessation of basaltic volcanism in the Palaeotethys (Korte et al.,2006) or to an initial rifting of Pangea leading to enhanced conti-nental flux (Kani et al., 2008; Isozaki, 2009) associated with a cool-ing event (Isozaki et al., 2007). The increase becomes morepronounced in the Early Triassic due to enhanced weathering re-flected in continental silicate rocks that resulted in wide-spreadclastic sedimentation, such as the Werfen Group of the WesternCarpathians and the Alps, the Germanic and Alpine Buntsandstein,the Moenkopi Formation of western North America and the Vetlu-ga Group of the Russian Platform (Korte et al., 2003). This likely re-sulted in large-scale sequestration of atmospheric carbon dioxide(e.g. Berner, 2004), analoguous to weathering episodes seen inthe 87Sr/86Sr-trend at the P–T boundary (Fig. 3). The 87Sr/86Sr levelsoff (Fig. 3) during the interval covered by the negative carbon iso-tope excursion (see also Twitchett, 2007). McArthur and Howarth(2004) originally attributed this to inaccuracies of geochronologi-cal constraints across the P–T boundary, but the reality of this lev-eling-off was later substantiated by better data, as presumablyreflecting the weathering of non-radiogenic basaltic rocks fromthe Siberian Traps (Holser and Magaritz, 1987).

Note nevertheless that the levelling-off trend in Sr isotope ra-tios precedes slightly the onset of volcanism and other as yet unac-counted factors must have therefore influenced the strontium-isotope ratios of seawater. For example, an enhanced weatheringof carbonates or evaporites on continents could have supplied lessradiogenic Sr to the oceans (Twitchett, 2007), although such a sce-nario is not supported by sedimentological and geochemical data(Ward et al., 2000; Sephton et al., 2005; Martin and Macdougall,1995).

6. Conclusion

Because of the good temporal correlation and the large magni-tude of the Siberian Traps, we propose that contemporaneous vol-canism was the main cause of the negative carbon isotopeexcursion at the P–T boundary. Volcanically outgassed CO2 andCO2 released by thermal metamorphism of pre-Trap sedimentson the Siberian platform seem likely to have been of particularimportance. Deep anoxic oceanic waters, reaching unusually shal-low depths, may also have been a contributing factor during the la-ter stages. The long duration of the carbon-isotope trend suggeststhat all short-lived events, such as a sudden breakdown in oceanicbiological activity due to mass extinction, sudden release of oce-anic methane hydrates or impact related causes can be ruled outas major causes of the negative carbon isotope excursion.

Acknowledgements

We acknowledge the analytical support of J. Hunziker (Lau-sanne) and C. Spötl (Innsbruck), technical assistance by M. Wim-mer (Innsbruck) and D. Lutz (Erlangen), and discussion with G.Shields (UC London) and C. Heubeck (Berlin). T. Algeo (Universityof Cincinnati) and P. Wignall (University of Leeds) are acknowl-edged for the reviews and pertinent comments. We thank theDeutsche Akademie der Naturforscher Leopoldina (BMBF-LPD9901/838 and 9901/8-116), the CSIR (India) and the Freie Univer-sität Berlin for contributions to financing this project and for pro-viding the necessary facilities. The sampling in Iran, carried outby H.W. Kozur, was supported by the Deutsche Forschungsgeme-inschaft and by the Geological Survey of Iran in Teheran and Täbris.

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