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

After one of the most remarkable turning points in Earth’s history, namely the latest Palaeozoic mass extinction, marine and terrestrial life remained strongly disturbed prior to full recovery, with an unusually long delay of more than five million years (e.g., Erwin, 1993, Erwin, 2006, Retallack, 1995, Eshet et al., 1995, Bowring et al., 1998, Kozur, 1998a, Kozur, 1998b, Rampino and Adler, 1998, Jin et al., 2000, Benton, 2003, Benton et al., 2004, Peng and Shi, 2009). The assumed cause(s) for this biotic crisis are still under discussion and include: (1) large-scale volcanic activity of Siberian flood basalt (e.g., Renne and Basu, 1991, Campbell et al., 1992, Conaghan et al., 1994, Renne et al., 1995, Kozur, 1998a, Kozur, 1998b, Reichow et al., 2002, Reichow et al., 2009, Kamo et al., 2003, Visscher et al., 2004, Courtillot and Olson, 2007, Isozaki, 2007, Ganino and Arndt, 2009, Svensen et al., 2009) and contemporaneous volcanism in South China (Yin et al., 1992, Kozur, 1998a, Kozur, 1998b); (2) ocean anoxia reaching unusually shallow depths (Wignall and Hallam, 1992, Isozaki, 1994, Isozaki, 1997, Kajiwara et al., 1994, Wignall and Twitchett, 1996, Wignall et al., 1998, Kato et al., 2002, Kidder and Worsley, 2004, Nielsen and Shen, 2004, Grice et al., 2005, Kump et al., 2005, Riccardi et al., 2006, Riccardi et al., 2007, Hays et al., 2007, Xie et al., 2007a, Algeo et al., 2008, Grasby and Beauchamp, 2009), with degassing of CO₂ (Knoll et al., 1996, Woods et al., 1999), H₂S (Kump et al., 2005, Kaiho et al., 2006, Riccardi et al., 2007, Meyer and Kump, 2008, Meyer et al., 2008), methane (Heydari and Hassanzadeh, 2003, Retallack et al., 2003), or a combination of these (Ryskin, 2003); (3) an oceanic acidification crisis due to increase in atmospheric CO₂ concentrations (Heydari et al., 2003, Fraiser and Bottjer, 2007, Payne et al., 2007); (4) low atmospheric oxygen levels (Berner, 2002, Berner, 2005, Retallack et al., 2003, Huey and Ward, 2005, Berner et al., 2007); (5) worldwide depletion of stratospheric ozone (e.g., Kozur, 1998a, Kozur, 1998b, Visscher et al., 2004, Kump et al., 2005, Sephton et al., 2005, Beerling et al., 2007); and (6) climate change caused by strong volcanism (volcanic winter), impact of a celestial body (but see Koeberl et al., 2004), or both (Stanley, 1988, Campbell et al., 1992, Erwin, 1993, Kozur, 1998a, Kozur, 1998b, Retallack et al., 1998, Jin et al., 2000, Mory et al., 2000, Kaiho et al., 2001, Becker et al., 2001, Becker et al., 2004). Reviews dealing with the latest Permian extinction were recently published by Erwin et al., 2002, Racki and Wignall, 2005, Isozaki, 2007, Isozaki, 2009a, Isozaki, 2009b, Knoll et al., 2007, Twitchett, 2007a, Twitchett, 2007b, Wignall, 2007, Bottjer et al., 2008, Krassilov and Karasev, 2009, Posenato, 2009.

The mass extinction and environmental changes close to the Permian–Triassic (P–T) boundary were accompanied by major perturbations in the global carbon cycle, as marked by a pronounced negative carbon-isotope excursion. Unusually high ¹³C-enrichments in Permian carbonates were reported in the 1960s and 1970s (e.g., Compston, 1960, Osaki, 1973), marking the highest values of the Phanerozoic (Veizer et al., 1980). Wilgus (1981) proposed a global negative carbon-isotope trend across the P–T transition from her own data (western USA) and from compilation of literature (Compston, 1960, Osaki, 1973, Galimov et al., 1975, Kirkland and Evans, 1976, Veizer and Hoefs, 1976, Magaritz and Schulze, 1980, Veizer et al., 1980), and this was later confirmed by Clemmensen et al. (1985), but only from a few data points in the E-Greenland successions.

Chen et al. (1984) presented the first detailed δ¹³C trend for the P–T boundary section at Meishan (South China), the Global Stratotype Section and Point (GSSP) of the Permian–Triassic boundary, showing a decline in excess of 6‰ in the latest Permian, reaching the lowest values at −3‰. Subsequent higher resolution carbon-isotope datasets, published by Xu et al. (1986) for Meishan, Shangsi (both South China) and Sovetashen (Armenia) and by Holser and Magaritz (1987) for the Kuh-e-Ali Bashi section near Jolfa (NW Iran), proposed that the negative peak is situated close to the P–T boundary. Further results from a shallow-marine section that has higher sedimentation rates than Meishan, the Tesero in Southern Alps (Italy), subsequently indicated that the negative carbon-isotope shift is gradual (Magaritz et al., 1988). Still later, Baud et al. (1989) showed that this δ¹³C excursion is present in several classic P–T boundary successions of the Tethys Realm as well as in the marginal seas, such as the sections at Jolfa (Holser and Magaritz, 1987), Vedi (Armenia), Sovetashen (Armenia), Emarat (N-Iran), Idrijca (W-Slovenia), Çürük Dağ (SW-Turkey), Kemer Gorge (Antalya), Nammal Gorge (Salt Range, Pakistan), Thongde (Zanskar, Himalaya, India), Shangsi (S-China), Meishan (S-China) and Kaki Vigla (Greece). Its amplitude was generally in the 4–7‰ range (e.g., +4 to 0‰ for Çürük Dağ; +4 to −3‰ for Shangsi). This feature argues strongly that the P–T boundary negative shift is global in nature. As the same time, Oberhänsli et al. (1989), while reporting a similar gradual δ¹³C decline for the San Antonio section of Southern Alps, also published the data for the Schuchert Dal section of Jameson Land at East Greenland that shows Late Permian values of about −3‰ followed by sudden two-peak negative shifts in excess of 20‰ near the P–T boundary, a feature contrasting with all investigated marine low- and mid-latitude boundary sections.

The pilot study of Holser and Magaritz (1985) for the Reppwand section of Carnic Alps (Austria) that did not show any discernible changes in carbonate δ¹³C at the P–T boundary was followed by the Holser et al. (1989) study of the adjacent Gartnerkofel core (see also Magaritz and Holser, 1991) that presents one of the most impressive carbon-isotope curves for the P–T boundary interval. The Gartnerkofel succession was deposited in a shallow-marine environment of the western Tethys more than 150 km away from the shoreline (Buggisch, 1974, Buggisch, 1978, Holser et al., 1989). Its moderately high sedimentation rates enabled generation of a high-resolution carbon-isotope record that shows a gradual negative trend, similar to that in Tesero (Magaritz et al., 1988), with a minimum very close to the biostratigraphic P–T boundary. A second two-peak minimum is reported for the I. isarcica Zone. The Holser et al. (1989) paper was followed by a plethora of publications that report both the carbonate and the organic δ¹³C trends for the above and additional successions. These were reviewed for example in Scholle, 1995, Corsetti et al., 2005, and they are listed in Fig. 1 and Table 1. The reported amplitudes, shapes, durations and particularly the assigned exact stratigraphic positions of the negative δ¹³C peaks appear to have been somewhat variable (e.g., Twitchett, 2007a, Metcalfe et al., 2008), but nevertheless recorded in a wide range of marine and continental deposits. If global, due to short residence time of carbon in the ocean, the peaks and troughs of δ¹³C excursions should represent coeval time markers and enable trans-continental stratigraphic correlations.

Here we review details of several P–T boundary carbon-isotope trends in the context to their biostratigraphic backgrounds, with updated biostratigraphy that may differ somewhat from that presented in the original publications. This review may then enable construction of a general P–T boundary δ¹³C_carb curve, and recognition of geological factors that may impacted the ocean/atmosphere system during the latest Permian.