Evidence for recurrent changes in Lower Triassic oceanic circulation of the Tethys: The δ13C record from marine sections in Iran

Palaeogeography, Palaeoclimatology, Palaeoecology 252 (2007) 355 – 369 www.elsevier.com/locate/palaeo

Evidence for recurrent changes in Lower Triassic oceanic circulation of the Tethys: The δ 13 C record from marine sections in Iran Micha Horacek a,⁎, Sylvain Richoz b,c , Rainer Brandner d , Leopold Krystyn c , Christoph Spötl d a

d

Institute of Mineralogy and Petrology, University of Graz, Universitaetsplatz 2, 8010 Graz, Austria b Geological Museum, UNIL-BFSH2, 1015 Lausanne, Switzerland c Department of Palaeontology, Vienna University, 1090 Vienna, Austria Institute of Geology and Paleontology, University of Innsbruck, Innrain 52, 6020 Innsbruck, Austria Accepted 30 November 2006

Abstract Stable carbon isotope curves derived from Lower Triassic carbonate rocks from three Iranian sections are established to investigate changes in the carbon cycle during the Early Triassic in this area. The sections are located in the south-center (Abadeh), north (Amol), and northwest (Zal) of Iran. All three curves show a similar pattern starting out with high δ13C values in the uppermost Permian decreasing across the Permian–Triassic boundary, an increase toward more positive values during the Griesbachian that slowly increase further up during the Dienerian, followed by a positive excursion to values as high as + 8‰ near the Dienerian/Smithian boundary. During the Smithian values return to below 0‰, whereas second positive excursion to values higher than + 3‰ is recorded at the Smithian/Spathian boundary, again followed by a drop in δ13C into the Spathian and a final excursion to positive values at the Spathian/Anisian boundary. The results from these Iranian sections are consistent with previous studies from Italy and China, thus strongly suggesting that the recorded δ13C variability represents at least Tethys-wide geochemical signals. Moreover, the new curves reveal evidence of high-amplitude, frequent oscillations pointing toward rapid and profound changes in the global carbon cycle during the Lower Triassic. Stratification of the ocean interrupted by episodic overturning transporting deep water to the ocean surface is a viable mechanism to account for the recorded isotope variations. Provided that the δ13C curve is representative of the global Lower Triassic ocean, it has high potential for accurately dating sedimentary successions via chemostratigraphy. © 2007 Elsevier B.V. All rights reserved. Keywords: P/T boundary; Triassic; Isotopes; Ocean circulation; Anoxia

1. Introduction

⁎ Corresponding author. Current address: Austrian Research Centers Seibersdorf Research, Seibersdorf, Austria. E-mail addresses: [email protected] (M. Horacek), [email protected] (S. Richoz), [email protected] (R. Brandner), [email protected] (L. Krystyn). 0031-0182/$ - see front matter © 2007 Elsevier B.V. All rights reserved. doi:10.1016/j.palaeo.2006.11.052

The Permian–Triassic boundary (PTB) has been intensively studied, as it represents the most devastating extinction event in Earth's history (Erwin, 1993). The extinction mechanisms and processes, however, still remain poorly understood. These mechanisms include a bolide impact, ocean poisoning, large-scale methane

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release, and a sudden temperature rise as well as a fall (e.g., see Hallam and Wignall, 1997; Erwin et al., 2002 for reviews of the possible causes). The period following the extinction event has been much less intensely studied. Available data show that the Lower Triassic was a period transitional between the greatest mass extinction ever and the faunal recovery of the subsequent Mesozoic ecosystems (Flügel, 1994). Paleontological investigations draw a picture of a delayed recovery of Lower Triassic taxa (e.g., Schubert and Bottjer, 1995) facing partially harsh living conditions, e.g., oceanic anoxia (Wignall and Twitchett, 2002) and increased siliciclastic sedimentation on the shelves (Twitchett and Wignall, 1996). The recovery only started at the end of the Lower Triassic or the beginning of the Middle Triassic (Flügel, 1994; Kozur, 2003). Marine calcite incorporates inorganic carbon of the ambient seawater with little fractionation. Unless exposed to extensive diagenesis (e.g., during burial, emergence and soil formation, large amounts of diagenetic fluids) marine carbonate rocks most likely retain a carbon isotope signature close to their original isotope composition (Scholle and Arthur, 1980; Marshall, 1992). Consequently, the δ13C value of marine carbonates most often is a robust proxy of past ocean δ13C variations. Many mass extinction events are accompanied by shifts in the marine δ13C composition, indicating that these events are accompanied by environmental changes which involve the carbon cycle or the extinction itself might be responsible for a change in the global carbon cycle (for a review see e.g., Walliser, 1996; Veizer et al., 1999). One of the best-known examples is the PTB, which has been extensively studied over the past few decades (e.g., Baud et al., 1989; Holser et al., 1989). The Early Triassic following the PTB has remained a far more unfamiliar period with an unclear paleoecological situation. Horacek et al., (2007-this volume) studied the δ13C variations in the Lower Triassic of northern Italy and showed large oscillations similarly to those recently found in China (Payne et al., 2004), Turkey and Oman (Richoz, 2006). To verify these trends and to improve our understanding about the ecological and environmental changes during the Early Triassic, we examined shallow-water carbonate sections in three different regions of the Iranian microcontinent representing different oceans (i.e., Palaeotethys vs. Neotethys) in order to establish a continuous highresolution marine δ13C record.

dentist's drill, avoiding weathered surfaces, veins and stylolites. From most specimens several spots were drilled to investigate a possible diagenetic overprint of different parts of the specimen. The samples were analyzed at the isotope laboratory at the University of Innsbruck using a Gasbench II connected to a Finnigan Delta plus XL isotope ratio mass spectrometer (for analytical details see Spötl and Vennemann, 2003). Values are quoted relative to VPDB and reproducibility was better than 0.1‰ (1σ) for δ13C and δ18O.

2. Methods

Abadeh is located in central southern Iran, about 150 km southeast of Isfahan (N32°53′43″, E53°12′17″, altitude 2000 m). For a general description see Taraz et al. (1981). The PTB interval has recently been studied

Sampling was performed on cut carbonate hand specimens under a binocular microscope using a

3. Locations and stratigraphy of study sections During the Latest Permian and the Early Triassic the Iranian microcontinent was located in the central Tethys near the equator (Besse et al., 1998) and belonged to a chain of terranes referred to as Cimmeria (Sengör, 1984). Together with certain parts of Turkey, Afghanistan, Tibet and Indochina this archipelago formed a barrier between the Palaeo- and Neotethys (Fig. 1) during the Lower and Middle Triassic. Of the three sections (Fig. 2), the first is situated at Kuh-e-Hambast, close to Abadeh in central Iran. The second site is exposed near the village of Zal, northeast of Tabriz (Azerbaidjan province) in northwestern Iran. Both sections are located within an Upper Permian to Lower Triassic, NW–SE-striking half-graben-system containing triangular wedges of synrift sediments. Tilting with hanging-wall subsidence and footwall uplift resulted in up to 800-m-thick sediments in the southwest (Abadeh, Zal) and only some tens of metres of terrestrial deposits in the northeast (north of Isfahan; see Yazdi and Shirani, 2002). Within the Lower Triassic, two shallowing upward megasequences are recognized. They may have been produced by pulsed extension (tectonostratigraphic cycles) or by eustatic sea-level changes. The third location lies in the Elburz mountain range along the road from Amol (near the Caspian Sea) to Tehran and is also known as Mangol (Hirsch and Suessli, 1973). Comparison of the isotopic signals of these three sections is of special interest and importance with respect to the palaeogeographic setting, as Abadeh and Zal belonged to the northern Neotethys margin whereas Amol was part of the southern Palaeotethys shelf. 3.1. Abadeh

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Fig. 1. Early Triassic paleogeography after Stampfli and Borel (2001). 1: Dolomites, northern Italy; 2: Zal, northwestern Iran, 3: Abadeh, central Iran; 4: Amol, northern Iran.

by Gallet et al. (2000), Heydari et al. (2000, 2003), Korte et al. (2004) and Richoz (2006). 3.1.1. Lithostratigraphy The sampled section starts at about −40 m below the PTB with dark gray to black bioclastic limestones rich in microfossils with few intercalations of pale gray limestone with small caverns and poor in microfossils (Fig. 3). It is overlain by fine grained pale red to gray limestone (−20 m) containing less abundant microfossils. Then follows the boundary clay (0–0.3 m), which is present in many sections worldwide and shows a high carbonate content in the Abadeh section. It is overlain by a 1.8 mthick gray limestone bank with several layers of dm-thick lenses with dome-shaped structures. Taraz et al. (1981) described them as thrombolites (microbial build-ups). Other authors (e.g., Korte et al., 2004; Richoz, 2006; Baud et al., 2007) followed this interpretation. On the other hand, Heydari et al. (2003) reported here a 1 m-thick layer of inorganically precipitated synsedimentary carbonate cement. In our thin sections we recognized fan-shaped calcite crystals of cm- to dm-length with square-ended terminations suggesting a probable aragonite precursor

and inorganic formation. The mudstone in between these crystal fans contains a mesh of filaments (?fungi). Microbial mats locally overlay these carbonate seafloor crusts. This extraordinary facies is followed by tens of metres of thin bedded, partly bioturbated micritic mudstones (vermicular limestones), peloidal packstone layers and intercalated marls. The thin layering is interpreted as distal storm layers, deposited at/below the storm wave base. The cm-thick tempestites, which episodically interrupt the background sedimentation of the vermicular limestone facies, are characterized by restricted downward bioturbation. In the first 15 m of the sequence they are rich in bivalves (Claraia). In contrast to the interpretation offered by Heydari et al. (2003) we fail to see an indication for a major drop in sea level. Strongly bioturbated, thin-bedded dark-gray vermicular limestone and marly limestones with thin distal storm layers containing peloidal packstones with hummocky cross lamination follow for another 300 m. The shallowing upward sequence is topped by a 80-m-thick oolite/ oncolite rich limestone with microfossils and bivalve shells forming a cliff. Ooids of the cross-laminated oolitic grainstones are well preserved with an intra-cortical

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Fig. 2. Map of Iran with sample locations. After Stepanov et al., 1969, and Mette and Mohtat 2004, modified.

M. Horacek et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 252 (2007) 355–369 Fig. 3. Correlation of Lower Triassic δ13C isotope curves and carbonate lithology in the three Iranian systems tract. TST = transgressive systems tract. Ages: W = Wuchiapingian, Ch = Changhsingian, G = Griesbachian, P = Permian. 359

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moldic porosity. The pores are infilled by crystalline equant calcite cement which is partly replaced by ferroan dolospar. This points to a primary aragonitic composition of the ooids. The Early Triassic was characterized by paleoceanographic conditions favoring aragonitic precipitates in seawater with an elevated Mg/Ca ratio and elevated sulphate levels (Sandberg, 1983; Hardie, 1996). Upsection the second shallowing-upward megasquence follows with marls and thin platy limestone with bivalve tempestites, oolites with microgastropods and flat-pebble conglomerate layers which are covered by platy darkgray, homogenous limestone with ooids/oncoids and gastropod shell detritus layers gradually changing its colour to pale gray and yellow and back to dark gray at 720 m. After a small tectonic overthrust the sequence ends with well-bedded dolomites of the highstand sequence tract. The storm dominated, mixed carbonate/silciclastic lithofacies of the second cycle is well comparable to the Werfen facies of the Dolomites of Italy (Mostler, 1982). 3.1.2. Biostratigraphy The PTB was identified 0.3 m above the boundary shale by the first appearance of Hindeodus parvus. Griesbachian strata are defined by conodonts (Hindeodus, Isarcicella), bivalves (Claraia griesbachi) and rare ammonoids (Ophiceras, Mesokantoa). The Dienerian base is marked at 15 m by the first appearance of Neospathodus dieneri (Gallet et al., 2000), and the same species also provides Dienerian ages until 200 m above the boundary shale. No age-diagnostic conodonts were found in the interval from 200 to 700 m but the Dienerian/Smithian boundary should be expected around 500–550 m based on a lithostratigraphic comparison (onset of storm-induced sediments such as oolites and tempestites) with Zal. The Smithian/Spathian boundary is drawn tentatively above the last occurrence of Hadrodontina found around 750 m above the PTB since the genus is usually common in pre-Spathian rocks (Perri, 1991). 3.1.3. Isotope stratigraphy The sampled interval (Fig. 3) starts in the Wuchiapingian with high C-isotope values comparable to those found in many other sections (Baud et al., 1989; Holser et al., 1989; Shao et al., 2000). At the Dhzulfian/ Dorashamian (Wuchiapingian/Changhsingian) boundary, there is a small decrease to values higher than +3‰, again followed by values around +4‰. Close to the PTB values steeply decrease to about −0.5‰ just below the PTB. The curve shows increasing values in the Griesbachian peaking at +2‰ followed by a gentle decrease to +1‰ (see Richoz, 2006 for more details). In the Dienerian the

δ13C values rise again to +4‰. A profound rise to a maximum of +8‰ is found at around 500 m at the Dienerian/Smithian boundary (Fig. 3, peak A) followed by a steep drop to low values less than −2‰. The minimum is succeeded by a large step to positive values around +5‰ at ca. 750 m (Fig. 3, peak B) with a declining trend in the uppermost samples. 3.2. Zal The Zal section is located about 2 km north of the village of Zal, which lies ca. 5 km west of the road from Tabriz to Julfa, about 30 km south of Julfa (N38°43′47″, E45°36′13″, 1600 m) (Fig. 2). The section starts in the Permian, where a steep small gorge has been eroded into Wuchiapingian sediments and the PTB is found on the northern slope of the gorge. Two volcanic dikes cut the Griesbachian sediments parallel to bedding. The entire Lower Triassic succession reaches ca. 700 m in thickness (Fig. 4). 3.2.1. Lithostratigraphy Similar to the Abadeh section, we also find two shallowing-upward megasequences capped by shallowwater carbonates. The sampled section starts 20 m below the PTB in the Upper Permian consisting mainly of gray nodular limestone and intercalated black shales, indicating a deep-water setting. In the uppermost Permian the prominent Paratirolites bed is well marked. It is conformably overlain by a gray, yellow to red boundary clay (defined as 0 m), marking the eventstratigraphic boundary to the Lower Triassic. The Griesbachian starts with the Elikah Formation featuring yellow-gray, laminated, platy limestones, bearing Claraia shells, oncoids and some stromatolitic structures at the base above the boundary clay. Several cm- to dmthick layers of carbonate seafloor cement were recognized 3 m above the boundary clay in a sequence of 6 m thickness. Crusts of equant calcite and minor celestine are overlain by microbial films and mats with oncolitic structures. These limestones are followed by thickbedded massive limestone layers (20 m), gray to darkgray and weakly bioturbated to laminated. At 43 m above the boundary clays, the sequence is interrupted by a 30-m-thick volcanic sill and continues above with dark-gray to black nodular limestone rich in Claraia at 95 m. Then follows gray, laminated limestone (130 m) that hosts another sill at ca. 178 m. Upsection at ca. 190 m, there is again dark-gray nodular limestone that is partly laminated or weakly bioturbated, containing abundant pyrite in some layers. The overlying compact but thin-bedded limestone (230 m) is partly fine-

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Fig. 4. Comparison of the δ13C isotope curve from the Zal section (Iran) and from L`Om Picol/Uomo (Italy; Horacek et al., 2007-this volume). P = Permian, Ch = Changhsingian, Bel = Bellerophon Formation, T = Tesero Oolite, A = Andraz Member. TST = Transgressive systems tract. For legend see Fig. 3.

laminated with rare layers containing oncoids. An intercalation of massive gray limestone and dominating gray to red oolitic/oncolitic layers starts at 270 m. It is followed by intensely burrowed pale gray to yellow marls with still intact fine lamination and upsection again oncolites (410 m) of yellow, red or gray colour. A cliff-forming limestone bank of well-preserved oolitic grainstones with large-scale trough cross stratification starts at 429 m. As in Abadeh, oomoldic porosity is present, but no evidence of dolomitisation. The limestones are overlain by dominantly marls (474 m), intercalated by several minor tempestite cycles with megaripples and few layers rich in microfossils. Upwards the carbonate component in the marls

decreases untiluntilu 670 m, followed by marls with an increasing carbonate content and limestones, brown to gray which become increasingly dolomitized upsection. The last samples (665 m) belong to the base of a succession of massive, thick-bedded dolomite rocks comparable to the Middle Triassic Shotori Dolomite of central Iran (Seyed-Emami, 2003). 3.2.2. Biostratigraphy The sampled section starts in the Dorashamian (Changhsingian) about 7 m above the Dzhulfian/ Dorashamian boundary. The PTB is determined by the first appearance datum (FAD) of Hindeodus parvus, 0.7 m above the base of the boundary shale (Korte et al.,

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2004). Isarcicella isarcica follows 11 m higher up and the bivalve Claraia dominates the remaining Griesbachian part of the section, which is much thicker than in Abadeh. Analogous to Abadeh and Amol, the base of the Dienerian is drawn closely below the level where the rich Claraia presence ends (136 m). A late(?) Dienerian age is still indicated by Neospathodus cf. dieneri at 480 m and a late Smithian age is suggested by Pachycladina at 570 m. 3.2.3. Isotope stratigraphy The δ13C record (Fig. 3) starts in the Dorashamian (Changhsingian) and shows values above + 3‰. Towards the PTB, there is a steep decrease with a minimum around 0‰. After a small positive excursion to + 2‰ the values vary mostly between 0.5 and + 1‰ in the Griesbachian with two declines to almost 0‰ before they climb again to +2‰. In the Dienerian, the δ13C values are well above + 2‰ and increase to a minor maximum of + 4‰, followed by lower values around +3‰. The Dienerian/Smithian boundary is marked by a pronounced positive excursion with values rising to a maximum of almost + 6‰ (Fig. 3, peak A), followed by a steep and undisturbed decrease to very low values below − 2‰. The minimum with − 2.7‰ is at ca. 600 m, just beneath a jump to higher values reaching + 3‰ (Fig. 3, peak B). This peak is again followed by decreasing values to a minimum below − 1‰, immediately beneath another jump to values above +2‰ at 750 m (Fig. 3, peak C) and decreasing values following afterwards. 3.3. Amol The Amol section is located directly on the road from the Caspian Sea to Tehran via Amol, approximately 10 km south of Amol (Fig. 2). The uppermost Permian sediments are almost completely hidden by talus, as is the PTB. Earlier reports describing the lithology and fossil content from this section referred to it as Mangol (Hirsch and Suessli, 1973).

Then follows a mixed carbonate/siliciclastic lithofacies with green and red marls, intercalated with minor cycles of distal to proximal shelly storm layers, flat-pebble conglomerates, oolites with microgastropods and mudstone beds of typical southern alpine Werfen facies (e.g., Horacek et al., 2007-this volume). At ca. 150 m, this succession is overlain by well-bedded, partly laminated dolomite. 3.3.2. Biostratigraphy Above brachiopod-dated Upper Permian limestones, about 50 m of Griesbachian Claraia- and conodontbearing rocks (Hindeodus, Isarcicella) were reported (Hirsch and Suessli, 1973). Hadrodontina aequabilis found at 50 m marks the Griesbachian/Dienerian boundary interval (Perri, 1991), and Hadrodontina found around 150 m indicate a Smithian age (Perri, 1991) for the uppermost limestones. As in Zal and other regions of Iran, the overlying dolomites are probably of Spathian age. 3.3.3. Isotope stratigraphy The isotope record (Fig. 3) starts in uppermost Permian carbonates with a single value at + 2‰. The PTB itself is not exposed, but as the outcrop above starts with thrombolitic structures known to be restricted to the first few metres above the PTB (Baud et al., 1997) we assume that only a very small part of the basal Triassic rocks is not exposed. The earliest Triassic samples show a decrease to low isotope values from 0 to − 1.5‰ and a slow increase again to 0‰. After a sample gap the values increase to a first maximum just below + 2.5‰, followed by a decrease to ca. + 1‰. Upsection follows a steep rise to a peak above +5‰, followed by a drop to values beneath − 2‰ and a jump again to positive values almost reaching + 3‰, with the uppermost samples showing heterogeneous values with an overall decreasing trend. 4. Discussion 4.1. Sequence stratigraphy and correlations

3.3.1. Lithostratigraphy At the base of the sampled section dark gray to black limestone of Dorashamian age is present. The PTB is covered and could therefore not be investigated. The basal Triassic features gray limestone that contains stromatolitic structures, suggesting that only a very small part of the lowermost Triassic is not exposed. Further upwards follows gray thick-bedded, bioturbated limestone (15 m), composed of carbonate mudstone, by and then with some oncoids or layers of microfossils.

Though of different thickness, the three sections can be reliably correlated using conodonts for the PTB and the Griesbachian segment, which is only 15 m thick in Abadeh and 30 m thick in Amol, but more than 100 m thick in Zal. Following the well-constrained boundary in Abadeh, the base of the Dienerian is similiarly drawn in Zal and Amol, i.e., closely below the distinct Claraia acme (Fig. 3). The Dienerian/Smithian boundary is well identified although fossils are rare, as it is probably

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located closely to the pronounced positive excursion that can be found in all sections. Neospathodus dieneri found at 480 m just beneath the maximum at Zal section indicates an age close to the Dienerian/Smithian boundary, whereas Pachycladina at 570 m gives a Smithian age. Comparison with Payne et al. (2004), who put the boundary directly at the δ13C peak, gives a Smithian age with the occurrence of Neospathodus waageni shortly after the peak. The Smithian/Spathian boundary is indirectly constrained and drawn above at ca. 730 in Abadeh, 600 m in Zal, and 150 m in Amol (always above the last Hadrodontina conodonts). As no conodont samples were taken further upsection, stratigraphic constraints are provided by correlation with existing δ13C isotope curves only. The lithology of the three investigated sections is very similar, with boundary clay at the PTB (where exposed), carbonate seafloor cement crusts and microbial structures at the base of the Lower Triassic followed by dominantly laminated or vermicular limestone, overlain by oolitic and oncolitic limestones intercalated with marls. We recognized two shallowing-upward sequences, or 3rd-order transgressive-regressive megacycles, which are capped by oolitic sand shoals at the end of the Dienerian, respectively by peritidal dolomites in the lower Spathian. The first megacycle starts in the Upper Permian and has a maximum flooding surface in the condensed section of the red limestones with firm-or hardgrounds in the Paratirolites beds of the Dorashamian. At the PTB, there are no sedimentological indications for a sea-level regression. Contrary to Heydari et al. (2003) and Kozur (2003), we notice that the boundary clay, carbonate seafloor cement crusts, and stromatolitic layers (or “thrombolites”) are not indicative of a shallow-water environment. Quite on the contrary, the deficit of the carbonate sediment production caused by the breakdown of the bioproduction mimics a transgressive systems tract (TST). A further transgressive event is seen in the Dienarian of the Abadeh section at ca. 150 m and in the Zal section at ca. 300 m, correlatable with the isotope curve (see Fig. 3). The second megacycle starts with a well-developed transgressive surface on top of the highstand oolitic shoals. An isochronous correlation of the transgression is not possible in this case: the transgressive surface (TS) in the Abadeh section lies above, and in the Zal section below, the readily correlatable peaks of the carbonisotope curve. This may be an indication for synsedimentary tectonics in the Lower Triassic along the passive continental margin. Also the strongly variable thickness of the sections supports this interpretation. The uniform development of the sections points to

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similar depositional environments during the Lower Triassic. 4.2. Carbon-isotope stratigraphy 4.2.1. Trends The Iranian Upper Permian samples show high δ13C values typical of other sections elsewhere (e.g., Baud et al., 1989) followed by a profound shift to lower values across the PTB. The values increase again during the Griesbachian, with a short-timed and well-indicated negative excursion at the Griesbachian/Dienerian boundary in all sections (Fig. 3). The isotope values continue to rise in the Dienerian culminating in a large positive excursion exceeding 5‰ in Amol, 6‰ in Zal, and 8‰ in Abadeh (peak A). Near this peak is the Dienerian/Smithian boundary. This peak is followed by a steep drop to values below − 2‰. Almost immediately after these minimum values follows a jump to values above + 2.5‰ in Amol, around +3‰ in Zal, and exceeding + 5‰ in Abadeh (peak B), which is correlated with the well-known positive excursion in the early Spathian. Isotope data from sediments further upsection were only obtained from the Zal section, where a steady decrease to values less than − 3‰ is succeeded by a final positive peak (ca. + 3‰; peak C) that is correlated with the Spathian/Anisian boundary. The observed δ13C isotope trend across the PTB mirrors the global curve known from several sections worldwide (e.g., Baud et al., 1989; Holser et al., 1989). High Permian values are followed by a steep decrease to minimum values around 0‰ or even slightly negative values (e.g., Musashi et al., 2001). In the Griesbachian, a gradual rise to higher values is observed (e.g., Holser et al., 1989; Krystyn et al., 2003), although the Iranian δ 13 C isotope patterns show some additional lowamplitude oscillations, probably reflecting regional variations. The positive excursion near the Dienerian/ Smithian boundary (peak A) can be found in other sections, e.g., in Italy (Horacek et al., 2007-this volume), China (Payne et al., 2004; these authors place the Dienerian/Smithian boundary exactly at the maximum of this peak), Turkey and Oman (Richoz 2006). This isotopic peak is followed by a steep decline to negative values which ends with another positive excursion (peak B) also found in Italy and China right at the Smithian/Spathian boundary (Payne et al., 2004; Horacek et al., 2007-this volume). Positive excursions found in sections from Pakistan (Baud et al., 1996) and Oman (Hauser et al., 2001 and Richoz, 2006) can also be correlated with this event. In the Spathian, low δ13C values following the positive peak before a rise to

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another positive values exceeding + 2‰ (peak C) have also been identified in sections from Romania, Albania, the Himalaya (Atudorei, 1999) as well as in China (Payne et al., 2004) with the latest positive excursion most probably representing the Spathian/Anisian boundary. 4.2.2. Diagenetic influences The significance of stable isotope values of marine (bulk-rock) carbonates as paleoenvironmental proxies depends heavily on the degree of diagenetic overprint of these rocks. We addressed this issue by carefully taking micro-samples of homogeneous micrites using a micro drill. Lower values of samples close to dykes show some diagenetic influence of magmatic fluids; where this effect is visible it is limited to few metres adjacent to the dikes. There is no correlation of low oxygen and low carbon isotope values (Fig. 5), which might infer diagenetic alteration of the isotope values (Curtis and Coleman, 1986). The high degree of correlation of temporal changes in the carbon isotopic composition between the Iranian sites and other marine sections elsewhere (Payne et al., 2004; Horacek et al., 2007-this volume) argues strongly in favour of minimal diagenetic effects on the δ13 C values of our samples. 4.2.3. Environmental factors The high isotope values in the Upper Permian have been attributed to coal formation during the PermoCarboniferous (Retallack, 1999) and high marine bioproductivity (Kakuwa, 1996) and the onset of anoxia in the Panthalassa deep sea (Knoll et al., 1996; Kakuwa, 1996). Several hypotheses have been put forward to explain the steep fall of the carbon isotope curve at the PTB: oxidation of organic matter (Retallack, 1995; Retallack et al., 1996; Looy et al., 1999), methane release (Krull and Retallack, 2000), ocean anoxia (Wignall and Twitchett, 1996; Isozaki, 1997), collapse of bioproduction (Erwin, 1994), ocean poisoning by hypercania (Knoll et al., 1996), and extensive forest fires and volcanism (Campbell et al., 1992; Renne et al., 1995). Most of these hypotheses were rejected by Berner (2002), who identified the release of methane as the key process for lowering the δ13C values. Note that Berner's calculations were based on the δ13C curve from Meishan by Xu and Yan (1993). Meanwhile, revision of the C-isotope stratigraphy of this section has yielded a much less pronounced isotope minimum (Jin et al., 2000). More recent studies suggest that the ocean circulation was the dominant factor forcing environmental changes at the PTB and during the Early Triassic

(Corsetti et al., 2005). Nevertheless the oxic or anoxic condition of the sea at the PTB interval is still a highly controversial topic, e.g., Wignall and Twitchett (2002) favoured anoxic shallow-water conditions, whereas others (e.g., Kakuwa, 1996; Heydari et al., 2003) presented evidence of oxic conditions. We speculate that there was a stagnant ocean with anoxic bottom water during the late Permian that was suddenly mixed during the PTB interval, bringing suboxic to anoxic deep waters to shallow-water levels (e.g., Algeo et al., 2007-this volume). The suboxic conditions probably remained for some times due to escaping methane from the seafloor sediments and/or H2S that were partly oxidized in the water column and, in this way, were removing the available oxygen from the water. It is likely that there were suboxic conditions even in very shallow water that caused the formation of framboid pyrite and hampered, but not necessarily prevented bioturbation in the sediments. A possible increase in seawater temperature across the PTB might additionally have lead to less available oxygen due to reduced gas solubility. The increase in δ13C subsequent to the boundary minimum has been attributed to “improving environmental conditions” (Holser et al., 1989). This is only true, however, if the hypothesis of a total collapse in bioproductivity causing the negative isotope excursion is valid. The rising isotope values might as well be explained by a decrease in methane release or an alteration of the ocean circulation. The profound peak at the Dienerian/Smithian boundary (peak A, with a maximum of +8‰ in Abadeh) followed by the negative excursion (with minimum values below −2.5‰) is difficult to explain solely by changes in bioproductivity. Even optimum conditions for the biota only result in moderate positive excursions (e.g., Lini et al., 1992). A plausible explanation is an increase in ocean anoxia/stratification (Woods et al., 1999), in combination with an increased input of nutrients from the continents, evidenced by increased siliciclastic input, e.g., in the Southern Alps (Twitchett and Wignall, 1996) and by a rise of the Sr- isotope values within the Early Triassic, indicating increased continental weathering (Korte et al., 2003). Higher nutrient load leads to an increase in bioproductivity and efficient removal of 12Cenriched organic matter from the surface-ocean water, giving rise to an increase in δ13C. The negative excursion following the peak can be explained by mixing of the ocean waters, leading to an unstratified and oxygenated ocean rich in 12C derived from the oxidation of organic matter, as stratified oceans are getting unstable due to warming of ocean bottom waters (Zhang et al., 2001).

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Fig. 5. δ13C versus δ18O plot of the isotope values of the investigated sections. No correlation of the carbon and oxygen isotopes can be detected.

Episodic repetitions of these changes in ocean circulation might also have caused the high frequency oscillations of δ13C like the positive peak at the Dienerian/Smithian boundary followed by a negative excursion and similar patterns at the Smithian/Spathian and Spathian/Anisian boundaries. Although the mechanism leading to this chain of processes remains to be fully understood, it appears that the inferred global reorganizations of ocean circulation and likely responses of the atmosphere were not unique in Earth history. Similar carbon-isotope oscillations occurred

episodically in the Paleozoic and also in the Precambrian (see Munnecke et al., 2003) and may have had comparable global mechanisms. Munnecke et al. (2003) amended a model developed by Jeppsson (1990) and Bickert et al. (1997) to account for these ecological changes. Humid intervals (termed H-periods) are characterized by increased siliciclastic input, “fouling of the reefs”, deposition of black shales on the shelves, a well-mixed ocean and an anti-estuarine circulation (Fig. 6). Arid intervals (termed A-periods) are identified by growing reefs, increased evaporation of the shallow ocean, reduced

366 M. Horacek et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 252 (2007) 355–369 Fig. 6. Models summarized by Munnecke et al. (2003) to explain ecological changes in the Silurian. Some features are not present at Lower Triassic times, information is lacking for some others (see text).

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siliciclastic input, ocean stratification and an estuarine circulation. This model, developed for the Silurian shows several features that also have been identified in the Lower Triassic: increased siliciclastic input in the Smithian during an interval of low δ13C values evidenced at least in one region (Twitchett and Wignall, 1996), increasing carbonate production starting at the Smithian/Spathian boundary with coeval higher δ13C (Table 1). But besides many parallels between the Munnecke et al. (2003) model and the Lower Triassic, some parameters cannot be addressed due to a lack of data (e.g., oxic or anoxic conditions on the deep shelf) or seem to differ (e.g., absence of reefs in the Lower Triassic). Possible mechanisms forcing oscillating alterations in ocean circulation as shown by the carbon isotope data are temperature falls or rises as well as strengthening or weakening of thermohaline circulation, although modelling by Zhang et al. (2001) suggested that it is difficult to sustain a stratified ocean for more than a few kiloyears. An explanation might be the restriction of a

Table 1 Typical features of the model presented by Munnecke et al., 2003, many of them have been reported from Lower Triassic events (For more explanations see text) Dienerian– Smithian– Spathian– Smithian Spathian Anisian Positive δ13C excursion Positive δ18O excursion Carbonate factories and reefs on shallow shelves Abundant stromatolites/ oncolites Oxygenated sediments in deep shelf environments Anoxic sediments below and above in deep shelf environments Hiatus near the base in shallow shelf environments Intercalated siltstones and sandstones on proximal shelves No glacial sediments in high latitudes Extinction events in the marine environment Low diverse plankton community Short duration High atmospheric CO2 Low permanent/long-term ecological impact

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layered ocean to the Tethys and variable intensity of water exchange with the Panthalassa, due to tectonics (Suzuki et al., 1998) or sea-level changes. The sparse data from the Panthalassa Ocean otherwise do not show significant deviations from the Tethys trend (e.g., Isozaki, 1997; Musashi et al., 2001). Alternatively, Kump et al. (2005) invoked the idea of massive hydrogen sulphide release from the deep ocean to the surface water and into the atmosphere to produce the extinction at the Permian/Triassic boundary as well as the retarded biotic recovery in the Lower Triassic. Unfortunately, existing S isotope data (e.g., Newton et al., 2004) do not provide unequivocal evidence either. 5. Conclusions The PTB in Iran represents a TST and there is no evidence for a shallowing event. The carbon isotope curves of Lower Triassic marine sections from northwestern, northern and south-central Iran show a very similar pattern defining a specific trend for the Early Triassic seawater in the Palaeo- and Neotethys, verifying previous results from Italy and China. We therefore suggest that the pattern represents a trend of at least Tethys-wide but probably worldwide significance. The isotope shifts can best be explained by rapid oceanographic changes, i.e., switching from ocean stratification to vigorous circulation. Possible trigger mechanisms are volcano-tectonic activities (Suzuki et al., 1998), or changes in temperature causing stratification or circulation. Comparison with other events in the Proterozoic and Paleozoic (Munnecke et al., 2003) reveal many similarities, indicating that similar events occurred episodically during the geological past. Changes in ocean circulation are regarded as the most likely causes for the delayed biotic recovery from the PTB mass extinction event.

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Funding for this study was partially provided by a travel grant from the Austrian Ministry of Education and Science, a Gandolf Dölter Scholarship and a PhD scholarship of the University of Graz to M.H., and by a grant of the Österreichische Akademie der Wissenschaften to R.B. and L.K. within IGCP project 467, and by a grant of the Swiss National Science Foundation project No. 20–67941.02 and a travel grant from the Swiss Academy of Science to S.R. Extensive support in the field was provided by B. Hamdi and the Geological Survey of Iran. Constructive and thoughtful reviews by

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P.J. Marenco, E. Heydari, and T. Bickert, and of an earlier version by W. Piller, are acknowledged. References Algeo, T.J., Ellwood, B.B., Nguyen, T.K.T., Rowe, H., Maynard, J.B., 2007. The Permo-Triassic boundary at Nhi Tao, Vietnam: evidence for recurrent influx of sulfidic watermasses to a shallow-marine carbonate platform. Palaeogeogr. Palaeoclimatol. Palaeoecol. 252, 304–327 (this volume). doi:10.1016/j.palaeo.2006.11.055. Atudorei, N.V., 1999. Constraints on the Upper Permian to Upper Triassic marine carbon isotope curve. Case studies from the Tethys. Unpubl. PhD Thesis, Lausanne, 160 pp. + appendix. Baud, A., Margaritz, M., Holser, W.T., 1989. Permian-Triassic of the Tethys: Carbon isotope studies. Geologische Rundschau 78, 647–677. Baud, A., Atudorei, N.V., Sharp, Z., 1996. Permian and Early Triassic evolution of the Northern Indian margin: carbon isotope and sequence stratigraphy. Geodinamica Acta 9 (2), 57–77. Baud, A., Cirilli, S., Marcoux, J., 1997. Biotic response to mass extinction: the lowermost Triassic microbialites. Facies 36, 238–242. Baud, A., Richoz, S., Pruss, S., 2007. The Lower Triassic anachronistic carbonate facies in space and time. Global and Planetary Change 55, 81–89. Berner, R.A., 2002. Examination of hypotheses for the Permo-Triassic boundary extinction by carbon cycle modelling. Proceedings of the National Academy of Sciences 99 (7), 4172–4177. Besse, J., Torq, F., Gallet, Y., Ricou, L.E., Krystyn, L., Saidi, A., 1998. Late Permian to Late Triassic paleomagnetic data from iran: constraints on the migration of the Iranian block through the Tethyan Ocean and initial destruction of Pangaea. Geophysical Journal International 135, 77–92. Bickert, T., Pätzold, J., Samtleben, C., Munnecke, A., 1997. Paleoenvironmental changes in the Silurian indicated by stable isotopes in brachiopod shells from Gotland, Sweden. Geochimica et Cosmochimica Acta 61, 2717–2730. Campbell, I.H., Czamanske, G.K., Fedorenko, V.A., Hill, R.I., Stepanov, V., 1992. Synchronism of the Siberian Traps and the Permian–Triassic boundary. Science 258, 1760–1763. Corsetti, F.A., Baud, A., Marenco, P.J, Richoz, S., 2005. Summary of Early Triassic carbon isotope records. C. R. Palevol. 4, 473–486. Curtis, C.D., Coleman, M.L., 1986. Controls on the precipitation of early diagenetic calcite, dolomite and siderite concretions in complex depositional sequences. In: Gautier, D.L. (Ed.), Roles of Organic Matter in Sediment Diagenesis. SEPM Special Publication, vol. 38, pp. 23–34. Tulsa, Oklahoma. Erwin, D.H., 1993. The Great Paleozoic Crisis: Life and Death in the Permian. Columbia University Press, New York. 327 pp. Erwin, D.H., 1994. The end-Permian mass extinction. Nature 367, 231–236. Erwin, D.H., Bowring, S.A., Jin, Y.G., 2002. The end-Permian mass extinctions. In: Koeberl, C., MacLeod, K.G. (Eds.), Catastrophic Events and Mass Extinctions: Impacts and Beyond. Geological Society of America Special Paper, vol. 356, pp. 363–383. Flügel, E., 1994. Pangean shelf carbonates: controls and paleoclimatic significance of Permian and Triassic reefs. Geological Society of America, Special Paper 288, 247–266. Gallet, Y., Krystyn, L., Besse, J., Saidi, A., Ricou, L.-E., 2000. New constraints on the Upper Permian and Lower Triassic geomagnetic polarity timescale from the Abadeh section (central Iran). Journal of Geophysical Research 105, 2805–2815.

Hallam, A., Wignall, P.B., 1997. Mass Extinction and their Aftermath. Oxford University Press. 307 pp. Hardie, L.A., 1996. Secular variation in seawater chemistry: an explanation for the coupled secular variation in the mineralogies of marine limestones and potash evaporites over the past 600 m.y. Geology 24, 279–283. Hauser, M., Martini, R., Burns, S., Dumitrica, P., Krystyn, L., Matter, A., Peters, T., Zaninetti, L., 2001. Triassic stratigraphic evolution of the Arabian–Greater India embayment of the southern Tethys margin. Eclogae Geologicae Helvetiae 94 (1), 29–62. Heydari, E., Hassandzadeh, J., Wade, W.J., 2000. Geochemistry of central Tethyan Upper Permian and Lower Triassic strata, Abadeh region, Iran. Sedimentary Geology 137, 85–99. Heydari, E., Hassanzadeh, J., Wade, W.J., Ghazi, A.M., 2003. Permian–Triassic boundary interval in the Abadeh section of Iran with implications for mass extinction: Part 1 — Sedimentology. Palaeogeography, Palaeoclimatology, Palaeoecology 193, 405–423. Hirsch, F., Suessli, P., 1973. Lower Triassic conodonts from the Lower Elikah formation, Central Alborz Mountains (Northern Iran). Eclogae Geologicae Helvetiae 66, 525–531. Holser, W.T., Schönlaub, H.-P., Attrep Jr., M., Boeckelmann, K., Klein, P., Margaritz, M., Orth, C.J., Fenninger, A., Jenny, C., Kralik, M., Mauritsch, H., Pak, E., Schramm, J.M., Stattegger, K., Schmöller, R., 1989. A unique geochemical record at the Permian/ Triassic boundary. Nature 337, 39–44. Horacek, M., Brandner, R., Abart, R., 2007. Carbon isotope record of the P/T boundary and the Lower Triassic in the Southern Alps: evidence for rapid changes in storage of organic carbon; correlation with magnetostratigraphy and biostratigraphy. Palaeogeogr. Palaeoclimatol. Palaeoecol. 252, 347–354 (this volume). doi:10.1016/j.palaeo.2006.11.049. Isozaki, Y., 1997. Permo-Triassic boundary superanoxia and stratified superocean: records from lost deep sea. Science 276, 235–238. Jeppsson, L., 1990. An oceanic model for lithological and faunal changes tested on the Silurian record. Journal of the Geological Society (London) 147, 663–674. Jin, Y.G., Wang, Y., Wang, W., Shang, Q.H., Cao, C.Q., Erwin, D.H., 2000. Pattern of marine mass extinction near the Permian–Triassic boundary in South China. Science 289, 432–436. Kakuwa, Y., 1996. Permian–Triassic mass extinction event recorded in bedded chert sequence in southwest Japan. Palaeogeography, Palaeoclimatology, Palaeoecology 121, 35–51. Knoll, A.H., Bambach, R.K., Canfield, D.E., Grotzinger, J.P., 1996. Comparative earth history and Late Permian mass extinction. Science 273, 452–457. Korte, C., Kozur, H.W., Bruckschen, H.W., Veizer, P., 2003. Strontium isotope evolution of Late Permian and Triassic seawater. Geochimica et Cosmochimica Acta 67 (1), 47–62. Korte, C., Kozur, H.W., Joachimski, M.J., Strauss, H., Veizer, J., Schwark, L., 2004. Carbon, sulfur, oxygen and strontium isotope records, organic geochemistry and biostratigraphy across the Permian/Triassic boundary in Abadeh, Iran. International Journal of Earth Sciences 93, 565–581. Kozur, H.W., 2003. Integrated ammonoid, conodont and radiolarian zonation of the Triassic and some remarks to stage/substage divisions and the numeric ages of the Triassic ages. Albertiana 28, 57–74. Krull, E.S., Retallack, G.J., 2000. δ13C depth profile from paleosols across the Permian–Triassic boundary: Evidence for methane release. Geological Society of America Bulletin 112, 1459–1472. Krystyn, L., Richoz, S., Baud, A., Twitchett, R.J., 2003. A unique Permian–Triassic boundary section from the Neotethyan

M. Horacek et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 252 (2007) 355–369 Hawasina Basin, central Oman mountains. Palaeogeography, Palaeoclimatology, Palaeoecology 191, 329–344. Kump, L.R., Pavlov, A., Arthur, M.A., 2005. Massive release of hydrogen sulfide to the surface ocean and atmosphere during intervals of oceanic anoxia. Geology 33 (5), 397–400. Lini, A., Weissert, H., Erba, E., 1992. The Valanginian carbon isotope event: a first episode of greenhouse climate conditions during the Cretaceous. Terra Nova 4, 374–384. Looy, C.V., Brugman, W.A., Dilcher, D.L., Visscher, H., 1999. The delayed resurgence of equatorial forests after the Permian–Triassic ecologic crisis. Proceedings of the National Academy of Sciences of the United States of America 96, 13857–13862. Marshall, J.D., 1992. Climatic and oceanographic isotopic signals from carbonate rock record and their preservation. Geological Magazine 129, 143–160. Mette, W., Mohtat-Aghai, P., 2004. Late Permian and Early Triassic microfossil assemblages of Iran. Berichte des Instituts für Erdwissenschaften Karl-Franzens Universität Graz 9, 263–265. Mostler, H., 1982. Bozener Quarzporphyr und Werfener Schichten. In: Mostler, H. (Ed.), Jahrestagung der Österreichischen Geologischen Gesellschaft: Excursion Guidebook, Innsbruck, vol. 4, pp. 43–80. Munnecke, A., Samtleben, C., Bickert, T., 2003. The Ireviken Event in the lower Silurian of Gotland, Sweden — relation to similar Paleozoic and Proterozoic events. Palaeogeography, Palaeoclimatology, Palaeoecology 195, 99–124. Musashi, M., Isozaki, Y., Koike, T., Kreulen, R., 2001. Stable carbon isotope signature in mid-Panthalassa shallow-water carbonates across the Permo-Triassic boundary: evidence for 13C-depleted superocean. Earth and Planetary Science Letters 191, 9–20. Newton, R.J., Pevitt, E.L., Wignall, P.B., Bottrell, S.H., 2004. Large shifts in the isotopic composition of seawater sulphate across the Permo-Triassic boundary in northern Italy. Earth and Planetary Science Letters 218, 331–345. Payne, J.L., Lehrmann, D.J., Wei, J., Orchard, M.J., Schrag, D.P., Knoll, A.H., 2004. Large perturbations of the carbon cycle during recovery from the end-Permian extinction. Science 305, 506–509. Perri, M., 1991. Conodont biostratigraphy of the Werfen Formation (Lower Triassic), Southern Alps, Italy. Bollettino della Societa Paleontologica Italiana 30 (1), 23–46. Renne, P.R., Zichao, Z., Richards, M.A., Black, M.T., Basu, A.R., 1995. Synchrony and causal relations between Permian–Triassic boundary crises and Siberian flood volcanism. Science 269, 1413–1416. Retallack, G.J., 1995. Permian–Triassic life crisis on land. Science 267, 77–80. Retallack, G.L., 1999. Postapocalyptic greenhouse paleoclimate revealed by earliest Triassic paleosols in the Sydney Basin, Australia. Geological Society of America Bulletin 111, 52–70. Retallack, G.J., Veevers, J.J., Morante, R., 1996. Global coal gap between Permian–Triassic extinction and Middle Triassic recovery of peatforming plants. Geological Society of America Bulletin 108, 195–207. Richoz, S., 2006. Stratigraphie et variations isotopiques du carbone dans le Permien supérieur et le Trias inférieur de quelques localités de la Néotéthys (Turquie, Oman et Iran). Mémoires de Géologie de Lausanne 45, 275. Sandberg, P.A., 1983. An oscillating trend in Phanerozoic non-skeletal carbonate mineralogy. Nature 305, 19–22. Scholle, P.A., Arthur, M.A., 1980. Carbon isotope fluctuations in Cretaceous pelagic limestone: potential stratigraphic and petroleum exploration tool. American Association of Petroleum Geologists Bulletin 64, 67–87.

369

Schubert, J.K., Bottjer, D.J., 1995. Aftermath of the Permian–Triassic mass extinction event: paleoecology of Lower Triassic carbonates in the western USA. Palaeogeography, Palaeoclimatology, Palaeoecology 116, 1–39. Sengör, A.M.C., 1984. The Cimmeride Orogenic System and the tectonics of Eurasia. Geological Society of America, Special Paper 195, 1–82. Seyed-Emami, K., 2003. Triassic in Iran. Facies 48, 91–106. Shao, L., Zhang, P., Dou, J., Shen, S., 2000. Carbon isotope compositions of the Late Permian carbonate rocks in southern China: their variations between the Wujiaping and Changxing formations. Palaeogeography, Palaeoclimatology, Palaeoecology 161, 179–192. Spötl, C., Vennemann, T., 2003. Continuous-flow IRMS analysis of carbonate minerals. Rapid Communications in Mass Spectrometry 17, 1004–1006. Stampfli, G.M., Borel, G.D., 2001. A plate tectonic model for the Paleozoic and Mesozoic constrained by dynamic plate boundaries and restored synthetic oceanic isochrones. Earth and Planetary Science Letters 196, 17–33. Stepanov, D.L., Golshani, F., Stöcklin, J., 1969. Upper Permian and Permian–Triassic boundary in North Iran. Geological Survey of Iran, Report 12 72 pp. Suzuki, N., Ishida, K., Shinomiya, Y., Ishiga, H., 1998. High productivity in the earliest Triassic ocean: black shales, Southwest Japan. Palaeogeography, Palaeoclimatology, Palaeoecology 141, 53–65. Taraz, H., Golshani, F., Nakazawa, K., Shimizu, D., Bando, Y., Ishii, K., Murata, M., Okimura, Y., Sakagami, S., Nakamura, K., Tokuoka, T., 1981. The Permian and the Lower Triassic systems in Abadeh region, Central Iran. Mem. Fac. Sci. Kyoto Univ., Ser. Geol. Mineral, vol. 47(2), pp. 61–133. Twitchett, R.J., Wignall, P.B., 1996. Trace fossils and aftermath of the Permo-Triassic mass extinction: evidence from northern Italy. Palaeogeography, Palaeoclimatology, Palaeoecology 124, 137–151. Veizer, J., Ala, D., Azmy, K., Bruckschen, P., Buhl, D., Bruhn, F., Carden, G.A.F., Diener, A., Ebneth, S., Goddéris, Y., Jasper, T., Korte, C., Pawellek, F., Podlaha, O.G., Strauss, H., 1999. 87Sr/86Sr, δ13C and δ18O evolution of Phanerozoic seawater. Chemical Geology 161, 59–88. Walliser, O.H. (Ed.), 1996. Global events and event stratigraphy in the Phanerozoic. Springer, Berlin. 333 pp. Wignall, P.B., Twitchett, R.J., 1996. Oceanic anoxia and the end Permian mass extinction. Science 272, 1155–1158. Wignall, P.B., Twitchett, R.J, 2002. Extent, duration and nature of the Permian–Triassic superanoxic event. Geological Society of America, Special Paper 356, 395–413. Woods, A.D., Bottjer, D.J., Mutti, M., Morrison, J., 1999. Lower Triassic sea-floor carbonate cements: their origin and a mechanism for the prolonged biotic recovery from the end-Permian mass extinction. Geology 27, 645–648. Xu, D.U., Yan, Z., 1993. Carbon isotope and iridium event markers near the Permian/Triassic boundary in the Meishan section, Zhejiang Province, China. Palaeogeography, Palaeoclimatology, Palaeoecology 104, 171–176. Yazdi, M., Shirani, M., 2002. First research on marine and nonmarin sedimentary sequences and micropaleontologic significance across Permian/Triassic boundary in Iran (Isfahan and Abadeh). Journal of China University of Geosciences 13 (2), 172–176. Zhang, R., Follows, M.J., Grotzinger, J.P., Marshall, J., 2001. Could the Late Permian deep ocean have been anoxic? Paleoceanography 16 (3), 317–329.