Permian Palaeotemperature and Ice Volumes History:

Evidence from Carbon and Oxygen isotopes

Die Temperatur- und Vereisungsgeschichte des Perms:

Rekonstruktion anhand von Kohlenstoff- und

Sauerstoffisotopenverhältnissen

Der Naturwissenschaftlichen Fakultät der Friedrich-Alexander-Universität Erlangen-Nürnberg zur Erlangung des Doktorgrades Dr.rer.nat.

vorgelegt von Chen Bo, M.Sc. aus Hunan, China

Als Dissertation genehmigt von der Naturwissen- schaftlichen Fakultät der Friedrich-Alexander-Universität Erlangen-Nürnberg

Tag der mündlichen Prüfung: 19. 07. 2012

Vorsitzender der Promotionskommission: Prof. Dr. Rainer Fink

Erstberichterstatter: Prof. Dr. Michael M. Joachimski

Zweitberichterstatter: Prof. Dr. Shen Shuzhong

Table of Contents

Zusammenfassung ...... 1

Abstract ...... 1

Chapter 1: Introduction ...... 9

1.1 The Permian ...... 9

1.2 Introduction to stable isotope geochemistry ...... 18

1.2.1 Principles ...... 18

1.2.2 Phosphate palaeothermometry ...... 21 1.2.3 Conodont elements as proxy for phosphate oxygen isotope analysis

...... 23 1.2.4 Factors influencing the oxygen isotope composition of biogenic

apatite ...... 25

1.3 Analytical methods ...... 27

1.4 Focus of this study ...... 31

Chapter 2: Carbon and conodont apatite oxygen isotope records of

Guadalupian–Lopingian boundary sections: climatic or sea-level signal? ...... 33

Abstract ...... 33

2.1 Introduction ...... 34

2.2 Geological setting ...... 36

2.3 Results ...... 38

2.4 Discussion ...... 40

2.4.1 Carbon isotopes ...... 40

2.4.2 Oxygen isotopes and palaeotemperature evolution ...... 42

2.5 Conclusions ...... 48

Chapter 3: Climatic warming in the latest Permian and the Permian–Triassic mass extinction ...... 51

Abstract ...... 51

3.1 Introduction ...... 51

3.2 Geological setting ...... 52

3.3 Results ...... 54

3.4 Discussion ...... 55

Chapter 4: Permian ice volume and palaeoclimate history: oxygen isotope proxies revisited ...... 61

Abstract ...... 61

4.1 Introduction ...... 62

4.2 Geological setting ...... 65

4.3 Results ...... 68

4.4 Discussion ...... 70

4.4.1 Permian oxygen isotope record of brachiopod calcite ...... 71

4.4.2 Permian oxygen isotope record of conodont apatite ...... 75

4.5 Conclusions ...... 83

Chapter 5: Conclusions ...... 87

References ...... 89

Tables ...... 117

Acknowledgements ...... 139

Curriculum vitae ...... 141

Zusammenfassung

Zusammenfassung

Um die Paläoklima- und Eisvolumengeschichte des Perms rekonstruieren zu können, wurde eine hochaufgelöste Sauerstoffisotopenkurve basierend auf der Analyse von Conodontenapatit aus Südchina, dem Iran und den USA (n= 356) erarbeitet. Die Rekonstruktion fokussierte im speziellen auf die Zeitabschnitte der Guadalupian–Lopingian und Perm–Trias Übergange, da diese Phasen durch dramatische Einschnitte in der globalen Faunendiversität charakterisiert sind. Sauerstoffisotopenanalysen an einzelnen Conodontengattungen belegen, dass die Gattungen Streptognathodus und Hindeodus v.a. im Oberflächenwasser lebten und damit die Temperaturen an der Wasseroberfläche wiedergeben. Das Habitat von Gondollela war abhängig vom Meeresspiegel, wobei sowohl eine Lebensweise in offenmarinen Oberflächengewässern als auch in tieferen Partien der Wassersäule möglich war. Diese Ergebnisse der Untersuchungen zeigen, dass im Falle des Perms Sauerstoffisotopenanalysen nicht an gemischten Conodontenassoziationen, sondern nur einzelnen Gattungen durchgeführt werden sollten. Unterpermische Conodonten aus Südchina zeigen relativ hohe Sauerstoffisotopenwerte zwischen 22–23‰ VSMOW, welche unter Annahme, dass das Eisvolumen im frühen Perm (Karbon-Perm Vereisung) vergleichbar zu jenem der Eisschilde des Letzten Glazialen Maximum (LGM) war, auf relativ warme tropische Temperaturen von 26 bis 30 °C hinweisen. Im Unterschied zu sedimentologischen Untersuchungen, die ein Abschmelzen der großen Gondwana Eisschilde im Sakmarian andeuten, weisen die Sauerstoffisotopenwerte (Abnahme im δ18O um 2‰) auf eine spätere Erwärmung und ein späteres Abschmelzen der Gletscher im Kungurian hin. Zwei Guadalupian-Lopingian Grenzprofile (Penglaitan und Tieqiao Profile/Südchina) wurden detailliert auf Änderungen in der Sauerstoffisotopie

1 Zusammenfassung untersucht. Kohlenstoffisotopenanalysen an Karbonaten wurden durchgeführt, um Änderungen im Kohlenstoffkreislauf dokumentieren zu können. Im Unterschied zu früheren Studien konnte keine bedeutende negative Kohlenstoffisotopenexkursion im mittleren Capitanian nachgewiesen werden. In den Sauerstoffisotopenverhältnissen ist im späten Captanian (J. xuanhanensis–C. postbitteri hongshuensis Conodonten Subzone) ein Anstieg um 1.5‰, in der C. postbitteri postbitteri Subzone eine Abnahme um 1‰ und in der C. dukouensis bis C. asymmetrica Zone eine weitere Abnahme um 2‰ zu verzeichnen. Dieser Verlauf in den Sauerstoffisotopenwerten lässt eine Erwärmung der Wassertemperatur um etwa 4 °C im späten Capitanian (J. postserrata bis J. granti Zone), eine Abkühlungvon 6 bis 8 °C im Übergangsbereich Guadalupian–Lopingian sowie frühen Wuchiapingian und eine erneute Erwärmung im Wuchiapingian (C. dukouensis bis C. liangshanensis Zone) erkennen. Der Temperaturanstieg im späten Capitanian kann mit der Hauptphase des Emeishan Plateaubasaltvulkanismus korreliert werden. Diese Gleichzeitigkeit lässt darauf schließen, dass die vulkanische

Emission von CO2 zu einer erhöhten Treibhausgaskonzentration in der Atmosphäre, einer Erwärmung, einem intensivierten hydrologischen Zyklus, einem erhöhten Nährstoffeintrag in die Ozeane und letztlich zu einer erhöhten Bioproduktivität geführt haben könnte. Eine erhöhte Primärproduktion ist durch die positive Kohlenstoffisotopenexkursion im späten Capitanian dokumentiert. Da die Meeresspiegelkurve parallel mit der rekonstruierten Paläotemperaturkurve verläuft, ist allerdings zu vermuten, dass die Änderungen in den Sauerstoffisotopenverhältnissen nicht ausschließlich auf klimatische Temperaturänderungen, sondern auch auf Änderungen in der Wassertiefe (und damit der Wassertemperatur), in der die Conodonten gelebt haben, zurückzuführen sind. Die Sauerstoffisotopenwerte von Conodonten (Südchina) aus dem

2 Zusammenfassung

Changhsingian zeigen einen markanten Anstieg zu Werten um 22‰ VSMOW. Dieser Anstieg weist ebenfalls auf eine Abwanderung von Clarkina in größere Wassertiefen in Folge des Meeresspiegelanstiegs im Changshingian und auf eine klimatische Abkühlung hin. Sauerstoffisotopenanalysen an Conodonten des Perm–Trias Übergangsbereichs (Meishan und Shangsi Profile/Südchina) zeigen eine Verringerung der Werte um 2‰ VSMOW in späten Perm und damit eine Erwärmung der tropischen Oberflächenwässer um 8 °C an. Der Trend zu leichteren Sauerstoffisotopenwerten geht einher mit dem weltweit nachgewiesenen Trend zu leichteren Kohlenstoffisotopenwerten (gemessen an

Karbonat- und organischem Kohlenstoff), welche auf eine erhöhte CO2 Emission infolge des Sibirischen Plateaubasaltvulkanismus und daran geknüpfter Prozesse zurückgeführt wird. Letztere könnten zu einer erhöhten Treibhausgaskonzentration in der Atmosphäre und der dokumentierten Erwärmung geführt haben. Die Koinzidenz der klimatischen Erwärmung mit dem Aussterbeereignis im Bereich der Perm–Trias Grenze deutet an, dass die klimatische Erwärmung einer der Gründe für das bedeutendste Aussterbeereignis in der Erdgeschichte war. Der Vergleich der in dieser Studie an Conodontenapatit erarbeiteten Sauerstoffisotopenwerten mit an Brachiopodenkalzit gemessenen Werten (publizierte Studien) zeigt, dass Brachiopoden als auch Conodonten von unterschiedlichen Paläokontinenten z.T. deutlich unterschiedliche Sauerstoffisotopenwerte aufweisen. Diese Unterschiede werden auf regional unterschiedliche Klimabedingungen (z.B. Verhältnis Evaporation zu Niederschlag) zurückgeführt. Eine Rekonstruktion von globalen Temperaturänderungen oder Änderungen im Eisvolumen kann daher nur anhand eines Vergleichs von mehreren Isotopenkurven aus unterschiedlichen Regionen erfolgen. Eine Interpretation einer Isotopenkurve, in welcher Daten

3 Zusammenfassung aus verschiedenen Regionen zusammengeführt wurden, ist nicht zielführend.

4 Abstract

Abstract

High-resolution oxygen isotope record based on 356 measurements of conodont apatite from several low latitudinal sections in South China, USA and Iran was composed in order to unravel Permian palaeotemperature and ice volume history. The reconstruction focussed especially on the Guadalupian–Lopingian (G–L) and Permian–Triassic boundary intervals since catastrophic decreases in biodiversity occurred during this time periods. Oxygen isotope analyses of different conodont taxa suggest that Streptognathodus and Hindeodus lived in near-surface seawater and recorded surface waters temperature, whereas the habitat of gondolellid genera was variable depending on sea level, with both near-surface and deeper waters as potential life habitat. The results of this study document that oxygen isotope studies on Permian conodonts should be performed on mono-generic conodont samples and that oxygen isotopes not only provide valuable palaeoclimatic information but also may help to constrain the life habitat of conodonts. The oxygen isotope record measured on Early Permian conodonts from South China exhibits relative high values between 22 and 23‰ VSMOW during the glaciated Early Permian, translating into warm seawater temperature between 26 and 30 °C, assuming that Early Permian ice volumes were comparable to the Pleistocene glacial maxima. In contrast to the earlier view that the Late Palaeozoic Ice Age (LPIA) terminated in the late Sakmarian, the South China conodont apatite oxygen isotope record suggests waning of the ice sheets in the Kungurian. Ice melting is indicated by a pronounced decrease in δ18O of 2‰ VSMOW, which is interpreted as reflecting the combined effect of climatic warming and glacial ice melting. Oxygen isotopes of conodont apatite as well as carbon isotopes of whole rock carbonates were studied in the G–L boundary sections Penglaitan and Tieqiao (Southwest China). Carbon isotopes of whole rock samples were

5 Abstract studied in order to document changes in the carbon cycle. In contrast to earlier studies, no major negative excursion in δ13C is observed in the middle Capitanian. Instead, a positive δ13C excursion is observed in the latest Capitanian with a 1.5‰ increase registered in the J. xuanhanensis Zone to C. postbitteri hongshuiensis Subzone, followed by a decrease of 1‰ within the C. postbitteri postbitteri Subzone and a 2‰ decrease in the C. dukouensis to C. asymmetrica Zone. The oxygen isotope record reconstructed from gondollelid conodonts suggests warming of water temperatures of about 4 °C in the late Capitanian (J. postserrata to J. granti Zone), cooling of about 6 to 8 °C across the G–L boundary and in the earliest Wuchiapingian, and again significant warming in the Wuchiapingian (C. dukouensis to C. liangshanensis Zone). The temperature increase in the late Capitanian can be correlated with the main phase of Emeishan volcanism suggesting that climatic warming may have resulted in an intensified hydrological cycle, fertilization of the oceans and enhanced primary productivity, the latter documented in the positive late Capitanian carbon isotope excursion. However, changes in sea-level seem to parallel reconstructed water temperatures suggesting that changes in water depth in combination with superimposed climatic changes may have been responsible for the observed temperature changes. Oxygen isotope record measured on Late Permian conodonts from South China show a significant increase to 22‰ VSMOW in the Changhsingian. This increase is interpreted as a consequence of climate cooling and Clarkina moving to deeper waters as consequence of the Changhsingian sea level rise. High-resolution oxygen isotope records from the Permian–Triassic boundary interval in the Meishan and Shangsi sections (South China) show a decrease by 2‰ in the latest Permian translating into warming of low latitudinal surface waters of 8 °C. The oxygen isotope shift coincides with the negative shift in carbon isotope ratios of carbonates suggesting that the addition

6 Abstract of isotopically light carbon to the ocean-atmosphere system by Siberian trap volcanism and related processes resulted in higher greenhouse gas levels and global warming. The major temperature rise started immediately before the main extinction phase, with maximum and harmful temperatures documented in the latest Permian (Meishan: bed 27). The coincidence of climate warming and the main pulse of extinction suggest that global warming was one of the causes of the collapse of the marine and terrestrial ecosystems. In addition, very warm climate conditions in the Early Triassic may have played a major role in the delayed recovery in the aftermath of the Permian–Triassic crisis. The Permian conodont apatite δ18O record is compared with published brachiopod calcite δ18O records. Brachiopods and conodonts from different palaeocontinents show significantly different δ18O values, suggesting differences in local climatic conditions (e.g. evaporation/precipitation ratio). As a consequence, secular changes in palaeotemperature and oxygen isotope composition of Permian sea water cannot be reconstructed from records combined from different areas, but have to be based on records from a specific area.

7

8 Chapter 1 Introduction

Chapter 1: Introduction

1.1 The Permian

The Permian (298.9 to 252.2 Ma) represents the final period of the Paleozoic Era and was first introduced in 1841 by Murchison for sedimentologically diverse deposits in the Ural Mountains and named after the ancient kingdom of Permia. The subdivision scheme of the Permian has been discussed until Jin et al. (1997) proposed an integrated Permian chronostratigraphic scheme which subdivided the Permian into the three series/epochs Cisuralian, Guadalupian and Lopingian with the Ural Mountains in Russia, the Guadalupe Mountains in the USA and South China as the representative areas, respectively. This subdivision scheme has been accepted by the International Commission on Stratigraphy with 9 stages defined to further subdivide the proposed epoch/series (Fig. 1.1). To date, for 6 out of 9 stages Global Boundary Stratotype Section and Points (GSSP) defining the stages boundaries were established (Davydov et al., 1998; Glenister et al., 1999; Jin et al., 2006a, b), preferentially by using conodont evolutionary successions. The Permian was a unique period during which global tectonics as well as environmental and faunal communities underwent dramatic changes. During the Permian, the supercontinent Pangaea was completed by the amalgamation of Siberia during the Uralian orogeny; all Pangaea straddled almost from the North to the South Pole and was surrounded by a single ocean named Panthalassa (Fig. 1.2). In its eastern part, Pangaea was incised by the Palaeotethys, an ocean basin between Laurasia and Gondwana. The Palaeothethys narrowed during the Permian as consequence of the opening of the Neothethys, an oceanic basin between Gondwana and the northward drifting Cimmerian terranes (Fig. 1.2).

9 Chapter 1 Introduction

Fig. 1.1 Chronostratigraphic scheme and major events in Permian. Chronostratigraphic scheme and absolute ages after Gradstein et al. (2004). Global diversity and extinction rates from Payne et al. (2012), glacial intervals in Gondwana after Fielding et al. (2008c), pCO2 estimates after Royer (2006) and Montañez et al. (2007), dashed line represents uncertainty between the two data sets. “Kamura event” based on Isozaki et al. (2007a, b), Eruption ages of Emeishan and Siberian Trap volcanism after Mundil et al. (2004) and Sun et al. (2010) respectively. Anoxia after Isozaki (1997), sea level curve after Haq and Schutter (2008), PD–present day. Pangaea was moving northwards with an average latitudinal displacement of 15–20° from the latest Carboniferous to latest Permian (Scotese, 2001) and the configuration of the continents resulted in a pronounced asymmetry with Gondwana in the southern hemisphere being larger than its northern counterpart Laurusia. The northward movement as well as the asymmetric

10 Chapter 1 Introduction distribution of the Permian landmasses may have had a profound impact on the Permian climate regime, ocean circulation, and the marine and terrestrial ecosystems (Link, 2009; Shi and Waterhouse, 2010).

Fig. 1.2 Palaeogeographic reconstruction for the Early (A) and Late Permian (B) (A: http://www2.nau.edu/rcb7/280moll.jpg; B: http://www2.nau.edu/rcb7/260moll.jpg) The Carboniferous–Permian has been long recognized as one the most significant “icehouse” episodes in Earth history (Fig. 1.3). The glacial sediments of the Late Palaeozoic Ice Age (LPIA) have been largely deposited in middle to high latitude Gondwana, whereas coeval strata in low latitudinal areas are characterized by cyclic sedimentary sequences as consequence of glacio-eustatic sea level changes driven by the waning and waxing of ice sheets (Heckel, 1986, 1994, 2008; Smith and Read, 2000; Wright and Vanstone, 2001). Initially, the glaciers of LPIA were considered as a single, large ice sheet that has waned and waxed persistently from the latest Mississippian to the early Permian (Veevers and Powell, 1987; Frakes and Francis 1988; Crowley and Baum, 1991, 1992; Ziegler et al. 1997; Hyde et al., 1999). However, recent studies have challenged this view since glacial deposits found in various regions of Gondwana occur in several phases separated by phases with non-glacial deposits. In addition, the waning and waxing of the different ice centers were not synchronous in the different areas (Isbell et al., 2003; Fielding et al., 2008a, b, c). The new findings imply that a more dynamic climate regime

11 Chapter 1 Introduction consisting of multiple glacial and non-glacial phases may have prevailed during the LPIA. The initial glaciations of LPIA characterized by small size and short-lived glaciers occurred in northern South America and northern Africa during the Late Devonian and early Mississippian (Caputo and Crowell, 1985; López-Gamundí, 1997; Crowell, 1999; Isbell et al., 2003; Caputo et al., 2008), followed by slightly larger extent ice sheets appearing in the Precordilleran region of Argentina and the Parnaíba and Amazon Basins of Brazil in the Viséan (López-Gamundí, 1997; Isbell et al., 2003; Limarino et al., 2006; Caputo et al., 2008; Gulbranson et al., 2010; Taboada, 2010). Stable isotope proxies suggest that substantial glaciers or ice sheets have been widespread in the Tournaisian by exhibiting a dramatic increase in carbon and oxygen isotope ratios (Mii et al., 1999; Saltzman et al., 2000; Saltzman, 2002, 2003a, b; Buggisch et al., 2008; Grossman et al., 2008), which is interpreted as indicative of enhanced organic carbon burial resulting in lowering of atmospheric CO2, climatic cooling, and consequently ice masses build-up. The glaciers further extended in South America (Henry et al., 2008; Holz et al., 2008; Rocha-Campos et al., 2008) and first appeared in eastern Australia during the Serpukhhovian (Fielding et al., 2008a, b, c), with considerable ice masse may being contained in this time glaciers since it corresponds to a significant increase in the oxygen isotope ratios of conodont apatite and brachiopods shells (Mii et al., 1999, 2001; Buggisch et al., 2008). Starting from the Bashkirian, the glaciers progressively expanded and became widespread across South America, South Africa and Australia (Isbell et al., 2003; Fielding et al., 2008a, b, c; Stollhofen et al., 2008), and finally occurred at Oman and Arabia at the beginning of the Moscovian (Martin et al., 2008). The extent of galcation may have shrunk during the late Pennsylvanian since the palynomorphs (Pfefferkorn et al., 2008) and stable isotopes (Frank et al., 2008)

12 Chapter 1 Introduction indicate climatic warming culminating in the Kasimovian to Gzhelian. In the latest Pennsylvanian, the glaciers expanded again and reached its peak extension during the Asselian and early Sakmarian, with ice sheets or glaciers occurring in most areas of Gondwana (López Gamundi, 1987; Veevers and Tewari, 1995; Isbell et al., 2003; Fielding et al., 2008a, b, c; Isbell et al., 2008a, c; Martin et al., 2008; Mory et al., 2008; Rocha-Campos et al., 2008; Stollhofen et al., 2008). The major ice caps are suggested to have melted in the Sakmarian (Veevers and Powell, 1987; Isbell et al., 2003; Korte et al., 2005, 2008; Zeng et al., 2012), however, the exact timing of the deglaciation is controversial since glacial deposits are found to appear intermittently in the late Early and Middle Permian of Australia and Siberia (Chumakov, 1994; Fielding et al., 2008c). The deglaciation represent the transition from an icehouse to a greenhouse state, a process interpreted to have had a substantial influence on the physical and chemical processes and thus on the ecosystems (Falcon-Lang, 2004; Joachimski et al., 2006; Clapham and James, 2008; Grossman et al., 2008; Heckel, 2008; Isbell et al., 2008a) and the early Permian can therefore serve as an analogue for understanding the current environmental changes resulting from the authropogenic rise in atmospheric CO2 (Gastaldo et al., 1996; Montañez et al., 2007; Isbell et al., 2008a). The “icehouse” mode of the LPIA is supported by model predictions suggesting that atmospheric carbon dioxide partial pressure (pCO2) was remarkably low (GEOCARB III model Berner and Kothavala, 2001; Royer et al., 2004), whereas oxygen concentrations were highest in the Earth's past 600 million year history (Berner, 1999, Fig. 1.3). The detailed pCO2 record reconstructed from the carbon isotope composition of palaeosols (latest

Pennsylvanian to Permian) shows a strong covariance between pCO2 and the 4 glacial events reported from Australia (Royer, 2006; Montañez et al., 2007;

13 Chapter 1 Introduction

Fielding et al., 2008c; Fig. 1.3), suggesting that CO2 played an important role in driving climate changes.

Fig. 1.3 Phanerozoic diversity (Sepkoski, 1981) and main “icehouse” periods (blue bars at the top; LPIA–Late Paleozoic Ice Age, Frakes et al., 1992), atmospheric pCO2 (GEOCARBIII Berner and Kothavala, 2001; Royer et al., 2004) and oxygen concentration (Berner, 1999). The onset of the LPIA as well as the consequent glacial/interglacial oscillations are suggested to have led to a pronounced reduction in extinction and origination rates, a decrease in the latitudinal diversity gradient of marine invertebrates (Stanley and Powell, 2003; Angiolini et al., 2005; Powell, 2007), and may have had an important influence on the distribution and composition of palaeotropical coal-forming forests on Pangea as well as the Northern and Southern China block situated in the eastern Palaeotethys (Cleal and Thomas, 2005). The Permian is characterized by the end-Guadalupian (259.0 Ma) and end-Permian (252.28 Ma) mass extinction events (Shen et al., 2010, 2011), the latter representing the most severe extinction event in Earth history. The end-Guadalupian extinction had been considered as part of the

14 Chapter 1 Introduction prolonged end-Permian crisis (Raup and Sepkoski, 1982). But subsequent studies recognized it as a separate event, independent of the end Permian extinction event (Jin et al., 1994; Stanley and Yang, 1994). The end-Guadalupian crisis was estimated to have eliminated approximately 58% of all marine invertebrate genera, comparable in magnitude to the end-Triassic or end-Cretaceous extinctions (Stanley and Yang, 1994; Knoll et al., 1996). The triggering mechanisms for this event were intensely debated, however, several significant environmental fluctuations have been found to have occurred approximately coincident with the mass extinction event, and thus were proposed as potential triggers. The pronounced regression occurring at the Roadian–Wuchiapingian boundary (Fig. 1.1; Haq and Schutter, 2008), resulted in erosional unconformities developed across many shelfal or platform successions around the World (Jin et al., 2006a) and the significant reduction of shallow water habitats may have contributed to the biotic crisis (Jin et al., 1994, 1995; Mei et al., 1998; Hallam and Wignall, 1999; Wang and Sugiyama, 2000; Yang et al., 2004). Emeishan volcanism was dated to have erupted in south China during the end-Guadalupian (Fig. 1.1; Zhou et al., 2002; Wignall et al., 2009a; Sun et al., 2010), thus time-equivalent to the mass extinction. This temporal link has led some researchers to suggest a cause-effect relationship between volcanism and the mass extinction (Courtillot et al., 1999; Wang and Sugiyama, 2000; Zhou et al., 2002; Wignall et al., 2009a, b), with climate warming and ocean acidification resulting from the injection of volcanic CO2 (Ganino and Arndt, 2009) as the potential trigger. On the other hand, Capitanian cooling named as the “Kamura event” (Fig. 1.1) has been proposed by Isozaki et al. (2007a, b) based on high δ13C values measured on shallow marine Capitanian carbonates from mid-Panthalassan atolls in Japan and western Tethyan successions in Slovenia (Isozaki et al., 2007a, b, 2011). The high δ13C values were interpreted as indicative of high primary productivity,

15 Chapter 1 Introduction enhanced organic carbon burial and thus low atmospheric CO2 levels culminating in comparatively cool climatic conditions. According to Isozaki et al. (2007a, b, 2011), climate cooling may have eliminated Guadalupian marine organisms that were adapted to warm tropical condition. In addition, catastrophic methane outburst and oceanic anoxia were discussed as other potential causes of the end-Guadalupian mass extinction (Isozaki, 1997; Retallack et al., 2006). The most severe decrease in biodiversity during the Phanerozoic occurred at the latest Permian, with 90% of all marine species becoming extinct, and a dramatic decline in terrestrial diversity (Retallack, 1995; Erwin et al., 2002). The mass extinction occurred in a relatively short time period of less than 200.000 years with ecosystems collapsing synchronously in the marine and terrestrial realms (Shen et al., 2011). The potential causes of the P–T mass extinction are one of the hottest topics in geosciences with many studies focusing on the triggering mechanisms; yet there is still no unambiguous answer. Currently, the mainstream views concerning the causes are (i) long-lived anoxia or euxina in the oceans with

CO2- and H2S-rich deep waters intruding into shallow depths and poisoning life (Fig. 1.1; e.g. Wignall and Hallam, 1992; Erwin, 1993; Isozaki, 1997; Knoll et al., 1996; Kump et al., 2005), and, (ii) the eruption of the Siberian flood basalts and associated progresses (Fig. 1.1; e.g. Renne et al., 1995; Kamo et al., 2003; Reichow et al., 2009; Svensen et al., 2009; Korte et al., 2010; Shen et al., 2011). Anoxic or euxinic conditions in both deep and shallow waters have been detected in many P–T boundary sections around the world using a variety of palaeoredox proxies including organic biomarkers, pyrite framboid sizes, sulfur isotopes, uranium isotopes and Ce anomalies (Isozaki, 1997; Nielsen and Shen, 2004; Grice et al., 2005; Shen et al., 2007; Cao et al., 2009; Bond and Wignall,

16 Chapter 1 Introduction

2010; Brennecka et al., 2011). However, the intensity and timing of redox state changes in the oceans remain uncertain, especially in shallow-marine environments with considerable spatial and temporal variability (Algeo et al., 2007a, b, 2008; Kakuwa, 2008; Bond and Wignall, 2010). Anoxia were previously suggested to have triggered the catastrophe by the reduction of oxygenated shallow-marine habitats (Wignall and Hallam, 1993), or by hypercapnia resulting from an overturn of the stratified oceans (Knoll et al., 1996). Recently, the release of hydrogen sulfide gas from euxinic seas was suggested to have caused ocean poisoning and damage of the ozone layer (Kump et al., 2005), but this idea is not supported by the modeling suggesting that, in the end-Permian, massive H2S escape from anoxic oceans seems unlikely to have seriously damaged Earth’s ozone shield or induced mass poisoning of terrestrial biota (Harfoot et al., 2008). The Siberian flood basalts represent one of the largest continental flood basalt (CFB) provinces on Earth with an estimated volume of 4 ×106 km3 (Courtillot et al., 1999). Radiometric dating of both the basalt eruptions and the mass extinction tend to favor the view that both events were essentially synchronous (Kamo et al., 2003; Mundil et al., 2004; Reichow et al., 2009; Shen et al., 2011). In early studies, sulfate aerosols resulting from the volcanic eruptions were interpreted to have triggered global cooling, potential ice sheet buildup and a glacio-eustatic lowering of sea level (Campbell et al., 1992; Renne et al., 1995). However, there are no known glacial deposits reported for this time interval. In contrast, the massive injection of greenhouse gases derived from the volcanic eruptions (Sobolev et al., 2011) or from contact metamorphism of organic carbon-rich shales and petroleum-bearing evaporites by intrusions of basaltic magmas (Svensen et al., 2009) and/or the destabilization of methane hydrates preserved in permafrost soils (e.g. Retallack and Jahren, 2008) have

17 Chapter 1 Introduction been proposed as a more plausible triggering mechanism. The massive inputs

CO2 are interpreted to have led to global warming, ocean acidification, ocean anoxia and potentially the destruction of atmospheric ozone, which, alone or in combination, may have resulted in a significant stress for the global ecosystems and may have caused mass mortality. This hypothesis best explains the timing, pace as well as the selectivity of the end-Permian mass extinction (Knoll et al., 2007; Shen et al., 2011) and is in accord with the major environmental perturbation observed at the P–Tr boundary interval such as the negative carbon isotope excursion in both organic matter and carbonate (for a review see Korte and Kozur, 2010), widespread wildfires and accelerated deforestation on land (Shen et al., 2011; Shen et al., 2011) as well as sudden prevalence of marine hypoxia and anoxia (e.g. Brennecka et al., 2011). High resolution oxygen isotope records are crucial in understanding the palaeoclimatic and biological events in the Permian by providing a powerful tool to quantitatively estimate ice volume changes and the termination of the LPIA, and to constrain the timing and magnitude of temperature perturbations during the end-Guadalupian and P–T mass extinction events.

1.2 Introduction to stable isotope geochemistry

1.2.1 Principles

While all isotopes of a given element share the same number of protons within the atom's nucleus, each isotope differs in its number of neutrons. A given element may in principle have any number of neutrons, and the number of nucleons (protons and neutrons) in the nucleus termed as the mass number, is different for two isotopes of any element. The absolute abundance of rare isotopes (such as 18O) is generally difficult to be determined accurately. Therefore, the relative difference in the isotope ratios of a sample and an isotopically well-characterized standard is generally measured and expressed in

18 Chapter 1 Introduction the δ notation as

δ= (Rsample–Rstd/Rstd) ×1000 (Ⅰ) 18 16 Rsample and Rstd are the ratios of the heavy/light isotope (e.g. O/ O) in the sample and standard, respectively. The δ values are reported in parts per thousand (per mill, ‰), with a positive value indicating that the ratio of the heavy vs. light isotope in the sample is higher than in the standard and conversely. All isotopes for a certain element have the same number and distribution of electrons, but have different masses, which cause subtle differences in their physicochemical properties. For example, molecules bearing heavy isotopes are generally more stable and have higher dissociation energies than those comprised of light isotopes. As a consequence, chemical bonds formed by light isotopes are easier to break than those formed by heavy isotopes. The mass difference causing differences in the reaction rates and during physical processes (e.g. diffusion, evaporation) is more pronounced amongst the light elements (e.g. H, O, C, S, and N). The partitioning of isotopes between different substances or two phases of the same substance with different isotopic compositions is termed as ‘fractionation’. The fractionation factor (α), which quantifies the isotopic fractionation between two substances or two phases of the same substance labeled as A and B, is defined as α = RA/RB. Using the δ notation, α can be expressed as

(Ⅱ)

The isotopic fractionation between two phases is also reported as ΔA-B =

δA – δB. The relationship between Δ and α is

19 Chapter 1 Introduction

3 δA –δB＝ΔA-B 10 lnαA-B (Ⅲ)

Isotope fractionation effects may be classified as: (1) equilibrium and (2) kinetic fractionation effects.

(1) Equilibrium isotope fractionation Equilibrium isotope fractionation is the partial separation of isotopes between two or more substances in chemical equilibrium, and is a type of mass-dependent isotope fractionation. It can be considered as the effect of atomic mass on bond energy with the heavier isotopes forming stronger chemical bonds than lighter isotopes. Consequently, a subtle preference is observed for the heavier isotopes are partitioned into one phase relative to another. Following Schauble (2004), (i) equilibrium isotope fractionation between two phases generally is dependent on temperature; (ii) elements characterized by higher mass difference of the light and heavy isotopes normally show a larger fractionation, and (iii), the heavy isotope is preferentially partitioned into the site with strong and short chemical bonds.

(2) Kinetic isotope fractionation. Kinetic isotope effects are generally associated with rapid, incomplete or unidirectional processes such as evaporation, dissociation, biologically mediated reaction as well as diffusion. Kinetic effects cause deviations from the simple equilibrium processes due to different reaction rates of different isotope species resulting from the vibration differences or indirectly from differences in the bonding energy.

Since equilibrium fractionation is a function of temperature, theoretically, the precipitation temperature of any two phases can be determined by calculating their isotopic fractionation provided the phases were precipitated in

20 Chapter 1 Introduction isotopic equilibrium and the temperature dependence of the equilibrium fractionation is known. Amongst the light stable isotopes, oxygen isotopes have been documented that to represent an ideal palaeothermometer. Oxygen occurs in nature as the three stable isotopes 16O, 17O, and 18O, with respective natural abundances of 99.76, 0.04 and 0.20% (Rosman and Taylor, 1998). Owing to the higher abundances of 16O and 18O and the larger mass difference, the 18O/16O ratio is of major interest in oxygen isotope palaeothermometry. In his classic paper concerning the thermodynamic properties of isotopic substances, H. Urey proposed in 1947 that changes in water temperature from which calcium carbonate was precipitated could lead to measurable variations in the 18O/16O ratios of calcium carbonate. Reconstructing temperatures of ancient oceans should thus be possible by measuring the oxygen isotope ratios of fossil biogenic calcite. This approach became feasible in 1950 with McCrea establishing the classic phosphoric acid method, by which the oxygen in carbonate is converted to CO2 by reaction of calcium carbonate with 100％ phosphoric acid at a constant temperature:

2+ 3- 3CaCO3+2H3PO4 3CO2+3H2O+3Ca +2 (PO4) (Ⅳ)

Since only two-thirds of the carbonate oxygen is liberated as CO2, an isotopic fractionation occurs between evolved CO2 gas and water. In order to be able to correct for this fractionation, the reaction temperature has to be constant.

1.2.2 Phosphate palaeothermometry

The idea to use oxygen isotopes in biogenic phosphate as a palaeothermometer was proposed by Urey et al. (1951). Parallel oxygen isotope analyses of carbonate and phosphate co-existing in marine shells were initially

21 Chapter 1 Introduction considered to have potential of providing accurate palaeotemperatures independent of the isotopic composition of the water from which the shells were precipitated (Urey et al., 1951; Tudge, 1960; Longinelli, 1965). However, subsequent studies showed that this approach was not useful due to the almost constant offset in the 18O/16O ratios of co-existing carbonate and phosphate precipitated at sedimentary temperatures (Longinelli, 1966; Longinelli and Nuti, 1973a). On the other hand, in case that the oxygen of water from which apatite was precipitated can be approximated, the oxygen isotope composition ratios of fish teeth and bones, conodont elements, mammal bones and marine phosphorites (Longinelli, 1965, 1966; Longinelli and Nuti, 1968a, b, 1973a, b; Kolodny et al., 1983; Shemesh, 1983; Longinelli, 1984; Luz et al., 1984) can serve as a powerful palaeotemperature proxy due to the following reasons: (i) Various studies have shown that the exchange of oxygen isotopes

between water and PO4 is extremely rapid and governed by equilibrium fractionation under multiple enzyme catalyzed metabolic

reactions involving PO4 (Kolodny et al., 1983; Blake et al., 1997). As a consequence, biogenic apatite is precipitated in oxygen isotopic equilibrium with ambient water and can be used to reconstruct palaeotemperatures. (ii) Apatite phosphate is relatively resistant to secondary alteration of the oxygen isotope ratios. Shemesh et al. (1983) compared the oxygen isotope ratios of cherts, limestone and phosphorites (Mishhash Formation of Israel/Campanian), and suggested that oxygen isotope record is considerably better preserved in phosphates than in either carbonates or cherts (Shemesh et al., 1983). This observation is consistent with subsequent experimental studies which exhibits that the phosphate component of bioapatite is very resistant to oxygen

22 Chapter 1 Introduction

isotopic exchange with water at low temperatures (Blake et al., 1997; Zazzo et al., 2004). The phosphate water fractionation equation as a function of temperature was initially established by Longinelli and Nuti (1973a) based on an empirical calibration of the δ18O values of phosphatic shells of living modern organism and average growth temperatures:

18 18 T (℃) =111. 4–4. 3(δ Ophosphate –δ Owater) (V)

18 18 T is the average growth temperature (°C), δ Ophosphate and δ Owater refers to the oxygen isotope ratios of the phosphate and environmental water, respectively. The temperature equation was refined by Kolodny et al. (1983):

18 18 T (℃) =113. 3–4. 38(δ Ophosphate–δ Owater) (VI)

More recently, Pucéat et al. (2010) published a new phosphate-water oxygen isotope fractionation equation based on oxygen isotopes determined on fish raised in aquariums at controlled temperature and with monitored water oxygen isotope composition. The new equation reveals a similar slope but an offset of about 2.2‰ compared with the old equations. A correction term for 18 the δ O value of NBS120c was introduced since different laboratories accept different values for standard NBS120c.

18 18 18 T (℃) =118.7–4.22(δ Ophosphate+(22.6–δ ONBS120c)–δ Owater) (VII)

1.2.3 Conodont elements as proxy for phosphate oxygen isotope analysis

Conodonts are a group of relatively small, eel-shaped marine animals which were abundant in Ordovician to Triassic oceans and which have been interpreted as one of the earliest jawless vertebrates (Donoghue et al., 2000; Sweet and Donoghue, 2001). The tooth-like conodont elements found in sediments are considered as the fossilized skeletal remains presumably

23 Chapter 1 Introduction representing a complex feeding or filtering apparatus (Donoghue et al., 2000; Sweet and Donoghue, 2001) composed of carbonate-fluorapatite with the following chemical composition:

Ca5Na0.14(PO4)3.01(CO3)0.16F0.73(H2O)0.85 (Pietzner et al., 1968)

Whereas the external morphology of conodont elements is diverse, the internal structure was demonstrated to be more or less identical. Conodont elements are composed of an upper crown and a lower basal body (Sweet et al 1988). Lamellar tissue of the crown typically consist of more densely packed lamellae with comparatively coarse prismatic apatite crystallites (typically 1–30 μm in length), whereas the basal body are more coarsely lamellar in structure and composed of finer cuboidal crystallites with a higher organic matter content, making the basal body more prone to mechanical abrasion and post-depositional alteration. The crown consists mainly of hyaline and albid tissue. Although hyaline and albid tissue are porous, permeability is considered as a more critical factor in controlling interaction and chemical exchange of phosphate-bound oxygen with oxygen from pore fluids (Trotter et al., 2007), with low permeability usually representing higher potential for retaining primary geochemical signatures. Interestingly, hyaline and albid tissue are found to have very low permeability because of their poor pore connectivity or special crystal shape and size (Trotter et al., 2007). Due to their carbonate-fluorapatite mineralogy and the dense crystalline structure, conodont elements are generally interpreted to have a high potential for the preservation of the primary oxygen isotope ratios (Pucéat et al., 2004). In addition, conodont elements are abundant in Paleozoic to Triassic marine sediments and additionally represent important index fossils for the biostratigraphic subdivision of the Paleozoic. These qualities make conodont apatite particularly attractive for oxygen isotope analyses (Joachimski and

24 Chapter 1 Introduction

Buggisch, 2002; Joachimski et al., 2006; Buggisch et al., 2008; Trotter et al.,

2008; Joachimski et al., 2009; Buggisch et al., 2010).

1.2.4 Factors influencing the oxygen isotope composition of biogenic apatite

The oxygen isotope composition of marine biogenic apatite is dependant on temperature and the oxygen isotopic composition of the sea water from which it was precipitated. The oxygen isotope ratio of sea water is controlled by evaporation, freshwater admixture by riverine runoff or the ice volume effect (Fig. 1.4). Evaporation will result in an enrichment of 18O in sea water, whereas admixture of freshwaters will shift the oxygen isotope composition of sea water to lighter values. Precipitation can lower δ18O of sea water through direct addition of 18O depleted rainfall, or via discharge of 18O depleted freshwater by rivers (Fig. 1.4). The δ18O of precipitation becomes increasingly lighter in higher latitudes since the air mass transport from the source to high latitudes results in cooling of the air masses, condensation and a preferential “rainout of 18O” which causes atmospheric vapour to become more and more depleted in 18O (Fig. 1.4). The δ18O composition of rivers is also related to latitude, with high latitude river systems having lower oxygen isotope ratios than those from low latitudes. In arid areas, enhanced evaporation will result in higher salinities and higher δ18O values of sea surface waters, whereas in regions close to the outlet of rivers, freshwater input will lower salinity and δ18O of surface waters. The volume of glacial ice sheets formed in high latitudes and polar regions represents an additional factor that affects the oxygen isotope composition in sea water on a global scale (Fig. 1.4). Since glaciers are generally sourced by 18O depleted precipitation in high latitudes or high altitudes, the formation of large ice sheets will result in 18O in sea water to become enriched. In contrast, ice sheet melting will result in a decrease of the 18O/16O ratio of sea water. For

25 Chapter 1 Introduction example, the ice sheets growth during the Last Glacial Maximum is suggested to have led to 1‰ increase in δ18O of sea water (Schrag et al., 1996) corresponding to 120–130m sea level lowering (e.g. Fairbanks, 1989; Yokoyama et al., 2000).

Fig. 1.4 Reservori size and oxygen isotope composition of water reservoris as well as fluxes between different reservoris. Modified after Sharp (2007). The question whether there was a secular change in the oxygen isotope composition of sea water is one of the controversial issues in isotope geochemistry. Basically, two conflicting views have been proposed: (i) The δ18O values of ocean water have been buffered at values around 0‰ VSMOW) over geological time by interaction with the oceanic crust (Gregory, 1991; Lécuyer and Allemand, 1999; Muehlenbachs et al., 2003). (ii) The δ18O value of ocean water was progressively depleted in 18O with increasing age because marine carbonates and silica becomes progressively lower with increasing age (Veizer et al., 1999;

26 Chapter 1 Introduction

Wallmann, 2001; Jaffrés et al., 2007).

To date, there is no agreement on this dispute, but new methods such as the clumped isotope palaeothermometer, in which temperature calculation is independent of the oxygen isotope ratios of the ocean (e.g. Came et al., 2007) and oxygen isotope records measured on biogenic phosphate being more resistant to diagenetic alteration may shed new light on this issue. Most important, for shorter time periods (e.g. < 1 Ma), a potential secular change in the oxygen isotope composition of the oceans can be assumed to be negligible. Consequently, any changes in the oxygen isotope ratios are interpreted to reflect change in temperatures, salinity and ice volume. In addition, knowledge concerning the habitat of analyzed conodont taxa is crucial for palaeotemperature reconstructions, since oxygen isotope ratios of surface dwelling taxa may reflect the combined effect of temperature, salinity and ice volume, whereas δ18O of taxa dwelling at greater water depths will be determined preferentially by changes in ice volume, with temperature and salinity variations becoming minor with greater water depths.

1.3 Analytical methods

The early attempts to use conodonts apatite for phosphate oxygen isotope analysis (Luz et al., 1984) were hampered mainly by the small size of the conodonts elements. However, with advances in the analytical techniques in recent years, oxygen isotope analysis became feasible on minor sample quantities. The early method for determining the oxygen isotope composition in biogenic phosphate was established by Tudge (1960). The PO4 group contained in phosphate was isolated as BiPO4 and the O2 liberated by conventional fluorination using BrF5. The generated O2 was converted to CO2 for mass spectrometric analyses (Tudge, 1960, Longinelli, 1966). This method was used

27 Chapter 1 Introduction for more than 30 year. However, the BiPO4 method had several drawbacks: (i) the reaction procedure is tedious and time consuming, (ii) relatively large samples are required, and most important (iii), BiPO4 is hygroscopic. In the

1990’s, a more rapid and easier analytical technique isolating the PO4 group of biogenic apatite as Ag3PO4 was presented by Crowson et al. (1991) and O'Neil et al. (1994). Ag3PO4 is routinely converted either to CO2 or CO. Either,

Ag3PO4 is heated with graphite in sealed silica tubes (O'Neil et al., 1994), releasing CO2 for mass spectrometric analysis, or, Ag3PO4 is reacted with graphite at high-temperatures using a TC-EA (high-temperature conversion-elemental analyzer) to produce CO (Kornexl et al., 1999; Vennemann et al., 2002). Vennemann et al. (2002) made a detailed comparison for the three different methods (including fluorination) and showed that the conventional fluorination method provides the most precise and accurate measurements as 100% oxygen is liberated from Ag3PO4. However, fluorination is time consuming, needs dangerous fluorine chemicals and most important requires large amounts of Ag3PO4 (4–5 mg). High-temperature reaction with graphite in isolated silica tubes is simple, rapid, and economic and with analytical reproducibilities comparable to the fluorination method, but it also requires relatively large amounts of sample. Additionally, the accuracy of this method varied in different laboratories, potentially due to combined effects of incomplete extraction of oxygen, oxygen exchange with the silica tube and/or differences in reaction temperature (O'Neil et al., 1994; Vennemann et al., 2002). In addition, differences in the size or type of the silica tube as well as the kind of graphite used may also contribute to the inaccuracy (Stephan, 2000; Vennemann et al., 2002). Therefore, this method must be calibrated separately in each laboratory by well-defined Ag3PO4 standards calibrated by conventional fluorination. The high-temperature carbon reduction by TC-EA represents a relatively rapid and simple method and has the

28 Chapter 1 Introduction advantage that relatively small samples are needed for analysis (1 mg for three to four replicate analyses). In addition, in situ oxygen isotope analysis by laser ablation (Cerling and Sharp, 1996), laser fluorination (Jones et al., 1999) or SHRIMP II ion microprobe (Trotter et al., 2008) have been developed recently for determining the oxygen isotope composition in apatite. These methods share an important advantage of high spatial resolution up to 100 µm, but also suffer from a lower analytical reproducibility and nonsystematic isotope offsets in comparison to the conventional methods, due to the variable contribution of oxygen-bearing carbonate and hydroxyl groups in biogenic apatite (Wenzel et al., 2000).

In this study, the PO4 group in conodont apatite was isolated as Ag3PO4, and the oxygen isotope analyses performed using a TC-EA coupled online to a ThermoFinnigan Delta Plus mass spectrometer. The chemical isolation of the phosphate group was as follows Joachimski et al. (2009). Conodonts were extracted by dissolving carbonate rocks in 10% acetic acid. Heavy liquid separation was performed using sodium–polytungstate. Conodont elements were picked from the heavy fraction under a binocular microscope. Monogeneric conodont assemblages were selected for separate isotope analyses in case of abundant conodont elements. Mixed (multi-generic) assemblages were used in case the abundances were not sufficient to compose mono-generic samples. Approximately 0.5 to 1 mg of conodont apatite was weighed into small polyethylene beakers and dissolved by adding 5 ml 2M HNO3. 2M KOH and 2M HF were added to the 2+ solutions for neutralization and precipitation of Ca as CaF2. The solutions were centrifuged and the supernatant transferred to a clean beaker using a pipette.

Silver amine solution (0.34 g AgNO3, 0.28 g NH4NO3, 1ml NH4OH in 30 ml deionized water) was added in order to precipitate the dissolved phosphate

29 Chapter 1 Introduction as trisilverphosphate. The beakers were put in an oven for 8 h at 60 °C. As ammonium was slowly evaporating, the pH of the solution decreased and

Ag3PO4 crystals started to precipitate (O'Neil et al., 1994).

The Ag3PO4 crystals were repeatedly rinsed with deionized water. After drying at 50 °C in oven for 1–2 hours, the Ag3PO4 crystals were homogenized using a small agate mortar. The homogenized Ag3PO4 powder was weighed into silver foil (0.2 to 0.3 mg per sample) and transferred to the sample carousel of the TC-EA. Samples and internal standards were generally measured in triplicate (He flow rate was 80 ml/s, reactor temperature was set to 1450 °C, column temperature was 90 °C). At 1450 °C, the silver phosphate is reduced and CO forms as the analyte gas (Vennemann et al., 2002). CO was transferred in helium stream through a gas chromatograph via a Conflo III interface to the mass spectrometer. Conodont apatite oxygen isotope values are reported in ‰ relative to VSMOW. Reproducibility of triplicate sample measurements was generally <±0.2‰ (1 std. dev.). The average oxygen isotope composition of the standard NBS 120c was 22.48 ± 0.13‰ VSMOW (n=65). Palaeotemperatures were calculated using the temperature equation for biogenic apatite published by Pucéat et al. (2010). In addition, few carbonate samples were collected systematically from the Penglaitan section for carbon isotope analysis. Carbonate powders for whole rock carbon isotope analysis were drilled from freshly cut rock surfaces. When possible, micrites were selected preferentially. Carbonate powders were reacted with 100% phosphoric acid at 70 °C using a GasBench II connected to a ThermoFisher Delta V Plus isotope ratio mass spectrometer (IRMS). All values are reported in per mil relative to V-PDB by assigning a δ13C value of +1.95‰ and a δ18O value of −2.20‰ to NBS19. Reproducibility was monitored by replicate analyses of laboratory standards and was better than ±0.05‰ (1 std.

30 Chapter 1 Introduction dev.) for carbon isotopes.

1.4 Focus of this study

In this study, a high-resolution oxygen isotope record for the Permian based on 356 measurements of conodont apatite from several low latitudinal sections in South China, USA and Iran is presented in order to clarify the following scientific issues: (i) Providing a high-resolution conodont apatite oxygen isotope record for the Permian. (ii) Constraining the magnitude and termination of the Late Paleozoic ice age. (iii) Unraveling potential palaeoclimatic changes during the catastrophic declines in faunal diversity during the end-Guadalupian and the end Permian. In Chapter 2, the reconstruction of the palaeotemperature history for the Guadalupian to early Lopingian interval is presented. Inferred climatic changes are discussed in context of sea level fluctuations and Emeishan volcanism. A high-resolution oxygen isotope record across the Permian–Triassic boundary, providing precise constraints on the timing and magnitude of global warming during this critical interval is presented in chapter 3. Chapter 4 presents the oxygen isotope record for the Permian. The conodont oxygen isotope record is compared with published brachiopod calcite δ18O records with the discussion concentrating on ice volume history and the timing of the termination of the LPIA.

31

32 Chapter 2 Guadalupian–Lopingian boundary

Chapter 2: Carbon and conodont apatite oxygen isotope records of Guadalupian–Lopingian boundary sections: climatic or sea-level signal?

Main parts of this chapter are publisched in Chen, B., Joachimski, M.M., Sun, Y.D., Shen, S.Z, Lai, X.L., 2011. Carbon and conodont apatite oxygen isotope records of Guadalupian–Lopingian boundary sections: Climatic or sea-level signal? Palaeogeography, Palaeoclimatology, Palaeoecology 311, 145–153.

Abstract

The Guadalupian–Lopingian (G–L) boundary (260.4 Ma) is one of the major extinction events in Earth history that coincides with the Emeishan Large Igneous Province and palaeoclimatic changes. Carbon isotopes of whole rock samples were studied in order to document changes in the global carbon cycle. In contrast to earlier studies, we observe no major negative excursion in δ13C in the middle Capitanian. A positive δ13C excursion is observed in the latest Capitanian with a 1.5‰ increase registered in the J. xuanhanensis Zone to C. postbitteri hongshuiensis Subzone, followed by a decrease of 1‰ within the C. postbitteri postbitteri Subzone and a 2‰ decrease in the C. dukouensis to C. asymmetrica Zone. Oxygen isotopes of conodonts from two G–L boundary sections were measured in order to reconstruct conodont habitat and potential changes in water temperature. Oxygen isotope ratios of gondolellid conodonts are higher in comparison to oxygen isotope ratios measured on hindeodid conodonts suggesting that gondollelids lived in cooler and thus deeper waters compared with hindeodids. The oxygen isotope record reconstructed from gondollelid conodonts suggests warming of water temperatures of about 4 °C in the late Capitanian (J. postserrata to J. granti Zone), cooling of about 6 to 8 °C across the G–L boundary and in the earliest Wuchiapingian, and again

33 Chapter 2 Guadalupian–Lopingian boundary significant warming in the Wuchiapingian (C. dukouensis to C. liangshanensis Zone). The temperature increase can be correlated with the main phase of Emeishan volcanism suggesting that climatic warming may have resulted in an intensified hydrological cycle, fertilization of the oceans and enhanced primary productivity, the latter documented in the positive late Capitanian carbon isotope excursion. However, changes in sea-level seem to parallel reconstructed water temperatures suggesting that changes in water depth in combination with superimposed climatic changes may be responsible for the observed temperature changes. This study documents that oxygen isotope studies on Permian conodonts should be performed on mono-generic conodont samples and that oxygen isotopes not only provide valuable palaeoclimatic information but also may help to constrain the life habitat of conodonts.

2.1 Introduction

The end-Guadalupian or late Capitanian mass extinction (Shen and Shi, 2009; Wignall et al., 2009b) is considered one of the most severe extinction events in Earth history (Jin et al., 1994; Stanley and Yang, 1994). The cause of this mass extinction event is still debated with various triggering mechanisms being proposed to have initiated the major decrease in marine and terrestrial faunal diversity, especially, the significant loss of fusulinacean foraminifera (Stanley and Yang, 1994; Yang et al., 2004; Bond et al., 2010), corals (Wang and Sugiyama, 2000), and brachiopods (Shen and Shi, 2002, 2009). Radiometric and biostratigraphic dating indicated that the Emeishan volcanic eruption coincided with this mass extinction (Zhou et al., 2002; Wignall et al., 2009a; Sun et al., 2010). This temporal link has led some researchers to suggest that flood basalt volcanism was a potential trigger of the extinction event (Courtillot et al., 1999; Wang and Sugiyama, 2000; Zhou et al., 2002; Wignall et al., 2009b). In addition to volcanism, a catastrophic methane outburst and

34 Chapter 2 Guadalupian–Lopingian boundary oceanic anoxia were suggested as other potential causes of the mass extinction (Isozaki, 1997; Retallack et al., 2006). The most prominent sea-level lowstand during the Palaeozoic occurred during the Capitanian –Wuchiapingian interval (Haq and Schutter, 2008; Wignall et al., 2009b) and resulted in erosional unconformities developed in many shelfal or platform successions around the World (Jin et al., 2006a). The regression and reduction of shallow water habitats may have severely affected marine life and contributed to the biotic crisis. Capitanian cooling (Kamura event) was suggested to have influenced marine faunal assemblages during the late Guadalupian (Isozaki et al., 2007a, b). High δ13C values measured on shallow marine Capitanian carbonates from mid-Panthalassan sections were interpreted as indicative of high primary productivity, enhanced organic carbon burial and thus low atmospheric CO2 levels culminating in comparatively cool climatic conditions (Isozaki et al., 2007a, b). In addition, Fielding et al. (2008a) reported sedimentary evidence for Capitanian ice fields in eastern Australia. However, the precise causal links between these potential causes and the biotic crisis have not been resolved. For example, high-resolution palaeotemperature records which are crucial for the understanding of the relationship between volcanism, climatic changes, glaciations, sea-level changes and the end-Guadalupian mass extinction are not available for the Guadalupian–Lopingian transitional interval. This contribution focuses on the reconstruction of palaeoenvironmental changes during the late Guadalupian and early Lopingian. Carbon isotopes of whole rock carbonates were studied to confirm changes in the global carbon cycle during this major crisis. Oxygen isotopes of conodont apatite which have been shown to faithfully record palaeotemperatures of ancient oceans (Joachimski and Buggisch, 2002; Joachimski et al., 2004, 2006; Trotter et al., 2008) were used to reconstruct the palaeotemperature history in the Guadalupian and early Lopingian. Palaeotemperatures reconstructed from

35 Chapter 2 Guadalupian–Lopingian boundary oxygen isotopes of gondollelid and hindeodid conodonts from the Penglaitan and Tieqiao G–L boundary sections (Guangxi Province, Southwest China) are discussed in the light of conodont habitat, potential climatic changes as consequence of Emeishan flood basalt volcanism and changes in water depth as consequence of sea-level changes.

2.2 Geological setting

The study area is situated in Laibin County (Guangxi Province, Southwest China) where Permian sediments are extensively exposed along the emerged riverbanks of the Hongshui River (Fig. 2.1). During Permian times, this area was situated within the Yunnan–Guizhou–Guangxi Basin of the South China block (Sha et al., 1990; Shen et al., 2007), which was situated near the equator in the Eastern Tethys (Metcalfe, 2002). The Yunnan–Guizhou–Guangxi Basin represents the southern extension of the Jiangnan Basin which was continuously subsiding between the Yangtze and Cathaysian blocks during Permian times (Wang and Jin, 2000). The Tieqiao section is located on the northern bank of the Hongshui River, about 5 km southeast of Laibin. The Penglaitan section represents the Global Stratotype Section and Point (GSSP) of the Guadalupian–Lopingian series boundary (Jin et al., 2006a) and is exposed on the southern bank of the Hongshui River, some 20 km east of Laibin. Structurally, the Penglaitan and Tieqiao sections are located on the eastern and western flanks of the Laibin syncline, respectively (Fig. 2.1). Both sections are characterized by continuous deposition in the late Guadalupian and early Lopingian, are well-dated by means of conodont biostratigraphy (Mei et al., 1998; Jin et al., 2006a), and were not affected by significant structural deformation or metamorphism (Jin et al., 2006a). Whilst in most other sections in the world the G–L boundary is characterized by a depositional hiatus as a consequence of the end-Guadalupian

36 Chapter 2 Guadalupian–Lopingian boundary global regression, the Penglaitan section is considered as one of the most complete Guadalupian–Lopingian (G–L) boundary sections worldwide (Mei et al., 1998; Jin et al., 2006a).

Fig. 2.1 (A) Late Permian palaeogeographic reconstruction with location of the study area on South China plate (http://www2.nau.edu/rcb7/260moll.jpg), (B) palaeogeography of South China during the Wuchiapingian (Wang and Jin, 2000; Shen et al., 2007; palaeolatitude after Zhang, 1997; Fang et al., 1992) and modern outcrop area of Emeishan volcanic province (Ali et al., 2002), and, (C) location of Penglaitan and Tieqiao sections (Jin et al., 2006). The depositional environment of the Laibin sections during most of the Guadalupian to early Wuchiapingian was a basin to slope setting with water depth estimated around 200 m (Wang and Jin, 2000). The lithostratigraphic units are the Maokou and Wuchiaping/Heshan Formations which are dominated by mostly basinal carbonates and radiolarian cherts. The uppermost 10 m of the Maokou Formation is the Laibin limestone member, which is mainly composed of pale grey medium to thick bedded massive limestones in the lower part and thin to medium bedded crinoidal tuffaceous limestones in the

37 Chapter 2 Guadalupian–Lopingian boundary upper part (Shen et al., 2007). The Laibin limestone is interpreted to have been deposited during the sea-level lowstand in the late Capitanian.

2.3 Results

Carbon isotopes measured on whole rock carbonates show a considerable scatter which is attributed to enhanced diagenetic fluid flow during chertification. In the Penglaitan section, carbon isotopes ratios show a 1.5 to 2‰ increase from the J. xuanhanensis to C. postbitteri hongshuiensis Subzone, a decrease of 1‰ within the C. postbitteri postbitteri Subzone and a second decrease of almost 2‰ in the C. dukouensis to C. asymmetrica Zone (Fig. 2.2). The oxygen isotopic composition of conodont apatite ranges from 19.0 to 21.8‰ in the Penglaitan and 18.6 to 20.9‰ in the Tieqiao section (Fig. 2.2). In the Tieqiao section gondolellids from the J. aserrata and J. postserrata conodont zones show relatively homogenous δ18O values between 20.5 and 21.0‰. Mixed gondolellid/hindeoid sample reveal slightly lighter δ18O values in comparison to mono-generic gondollelid samples. Gondolellids from the J. xuanhanensis Zone show values around 20‰. A minimum value of 19.1‰ is measured for a mono-generic sample of Hindeodus. δ18O values start to increase from 19.9 to 20.9‰ in the J. granti to C. postbitteri postbitteri Subzone with the base of the C. postbitteri postbitteri Subzone representing the G–L boundary. Another decrease in δ18O from 20.9 to 19.0‰ is observed from the C. postbitteri postbitteri Subzone to the Clarkina leveni and C. liangshanensis Zone. In the Penglaitan section, a value of 20.9‰ is measured on gondollelids from the J. shannoni Zone (Fig. 2.2). δ18O values of gondolellids decrease in the J. prexuanhanensis and lower J. granti Zone. Lowest values of 19.0 and 19.1‰ are measured on hindeoids from the J. granti Zone. As in the Tieqiao section, δ18O of gondolellids start to increase in the J. granti Zone with

38 Chapter 2 Guadalupian–Lopingian boundary maximum values of 21.8‰ measured in the C. dukouensis Zone.

Fig. 2.2 Carbon isotopes of whole rock carbonate and oxygen isotopes of conodonts from the Penglaitan (A) and Tieqiao (B) sections. Lithology and conodont biostratigraphy after (Mei et al., 1998; Shen et al., 2007), bed numbers after (Sha et al., 1990), absolute ages after (Gradstein et al., 2004).

39 Chapter 2 Guadalupian–Lopingian boundary

2.4 Discussion

2.4.1 Carbon isotopes

The carbon isotope record measured in the Penglaitan GSSP section (Wang et al., 2004; this study) shows a 1.5‰ increase in the J. granti Zone and the C. postbitteri hongshuiensis Subzone and a decrease in δ13C of 1‰ within the C. postbitteri postbitteri Subzone and of 2‰ in the C. dukouensis to C. asymmetrica Zone (Fig. 2.3). Measured δ13C values are comparable to values reported by Isozaki et al. (2007b) from the Kamura sections although the exact correlation of the carbon isotope records is hampered by the circumstance that the Kamura sections are dated by (rare) fusulinids. For example, the G–L boundary in the Kamura sections has been placed tentatively at the transition of a zone barren of fusulinids and the Codonofusiella–Reichelina fusulinid Zone. However, the comparison of the Kamura carbon isotope record with the δ13C record of the Penglaitan section (Fig. 2.3) supports the earlier interpretation of Isozaki et al. (2007b) that the G–L boundary is characterized by a minor positive excursion. In both records, δ13C values increase from average values around +4‰ to maximum values of +5‰ at the G–L boundary followed by a decrease to values between +3.5 and +4‰ above the G–L boundary, relatively homogenous δ13C values between +3.5 and +4‰ in the C. dukouensis Zone and a further decrease at the transition of the C. dukouensis to C. asymmetrica Zone (Fig. 2.3). Wignall et al. (2009a) and Bond et al. (2010) reported a 5 to 7‰ negative excursion in δ13C in different sections within the Emeishan large igneous province spanning the time interval of the J. altudaensis and J. prexuanhanensis Zone (Fig. 2.3). Bond et al. (2010) used the negative δ13C shift in the Kamura sections that was originally assumed to have occurred above the G–L boundary by Isozaki et al. (2007b) to argue for a global

40 Chapter 2 Guadalupian–Lopingian boundary significance of the negative carbon excursion in the J. altudaensis and J. prexuanhanensis Zone. However, as shown above, the comparison of the carbon isotope records from the Penglaitan and Kamura sections suggests that the initial age assignment of the negative shift in δ13C in the Kamura section by Isozaki et al. (2007b) was correct. In addition, this large negative excursion is not confirmed by carbon isotope values from the Penglaitan section that are between +2.0 and +4.5‰ in the correlative time interval (Fig. 2.3). These high carbon isotope values cannot be explained by diagenetic alteration, since diagenesis generally results in a lowering of the δ13C values (Weissert et al., 2008). In summary, it is questionable whether the negative carbon isotope excursion reported by Bond et al. (2010) is of global significance.

Fig. 2.3 Comparison of carbon isotope records from Penglaitan (this study) with published records (Wang et al., 2004; Isozaki et al., 2007a, b; Wignall et al., 2009a; Bond et al., 2010). Correlation of isotope records from Penglaitan (Southwest China) and Japan (Isozaki et al., 2007a, b) based on carbon isotope ratios, absolute ages after (Gradstein et al., 2004).

41 Chapter 2 Guadalupian–Lopingian boundary

Short-term changes in the carbon isotope ratios of inorganic carbon are related to changes in ocean productivity or the varying input of isotopically light carbon whereas on a longer timescale (106–107 years) the carbon isotope ratios may reflect changes in the burial ratio of organic vs. inorganic carbon. The duration of the positive carbon isotope excursion in the end-Guadalupian is estimated to have been less than 106 years and may thus indicate an increase in primary productivity potentially triggered by enhanced nutrient inputs. Interestingly, Wignall et al. (2009b) observe a change to dysoxic conditions in the J. granti Zone that could be the expression of higher organic carbon inputs and enhanced O2 consumption in the deeper water column. In comparison, the Valanginian positive carbon isotope excursion (Weissert oceanic anoxic event) was explained by an increase in atmospheric pCO2 as consequence of Paranà-Etendeka volcanism, triggering a climate change, accelerated hydrological cycling and an indirect fertilization of the oceans (Erba et al., 2003). The temporal relationship between Emeishan volcanism and changes in carbon and oxygen isotope ratios in the Penglaitan, Tieqiao and Panthalassan sections is suggesting a comparable causal link between volcanism; climate and the fertility of end-Guadalupian oceans (see discussion below).

2.4.2 Oxygen isotopes and palaeotemperature evolution

The oxygen isotopic composition of biogenic apatite is dependent on temperature and δ18O of ambient water (Kolodny et al., 1983). In order to calculate sea water temperatures from δ18O of conodont apatite, the oxygen isotope ratio of Permian sea water has to be estimated. The oxygen isotopic composition of sea water may be influenced by the input of freshwater or higher evaporation, especially in near coastal settings. Both, the Penglaitan and Tieqiao section were deposited in slope to basinal settings representing fully marine conditions with major changes in salinity being unlikely. In addition,

42 Chapter 2 Guadalupian–Lopingian boundary the oxygen isotopic composition of oceanic waters may be influenced by the build-up of continental ice caps. Fielding et al. (2008a) described Capitanian glacial deposits from eastern Australia. Since ice is generally depleted in 18O in comparison to the oxygen isotope ratio of the ocean, Capitanian ocean waters may have been enriched in 18O. However, Capitanian ice masses are interpreted to have been of limited extent potentially representing only alpine-style glaciation (Fielding et al., 2008a; Birgenheier et al., 2010). Since alpine glaciers will have a very minor effect on the oxygen isotopic composition of the ocean, we tentatively assume that oceanic waters in the Yunan–Guizhou–Guangxi Basin were characterized by an oxygen isotope ratio of -1‰ VSMOW, a value generally assumed for normal marine waters in an ice-free world (Savin, 1977). Since conodonts were nektonic organisms, the life habitat of the studied conodont taxa is crucial for the interpretation of reconstructed palaeotemperatures. Conodont faunas analysed from the Penglaitan and Tieqiao sections are dominated by gondolellids (Jinogondolella and Clarkina) except for the short time interval of the Laibin Limestone member in which hindeodids occurred (Jin et al., 2006a). Jinogondolella is interpreted to have lived in the low-latitudinal warm water realm during the Guadalupian (Mei and Henderson, 2001). Clarkina is usually considered as a free-swimming taxon thriving in offshore (Orchard, 1996; Lai et al., 2001) and potentially deeper waters (Tian, 1993a, b). The habitat of Hindeodus is more difficult to constrain with some researchers considering it a shallow-water taxon living in warm waters (Orchard, 1996), whereas Lai et al. ( 2001) interpreted Hindeodus to have been a pelagic taxon that potentially could have lived in shallow as well as deeper waters. Interestingly, in both sections, monogeneric samples of Hindeodus from the Laibin limestone member show significantly lower δ18O values and thus

43 Chapter 2 Guadalupian–Lopingian boundary higher temperatures in comparison to monogeneric Jinogondolella or Clarkina samples from below and above (Fig. 2.2; Table 2.1). The 0.8‰ (Tieqiao) and 1.0‰ (Penglaitan) lower δ18O values measured on hindeodids translate into 3 to 4 °C warmer temperatures. These higher temperatures suggest a shallower surface water habitat for Hindeodus, whereas Jinogondolella and Clarkina may have lived in cooler, deeper waters. Applying a temperature-depth gradient typical for modern tropical oceans (Levitus and Boyer, 1994), the 4 °C lower temperatures reconstructed for gondollelids relative to hindeodids indicate water depths of 60 to 80 m (assuming a surface water habitat for Hindeodus), suggesting that Jinogondolella and Clarkina were thriving in subsurface waters. In the Tieqiao section, palaeotemperatures reconstructed from the δ18O of gondollelids were between 26 and 29 °C in the J. aserrata (Wordian) to J. postserrata Zone (Capitanian). Temperatures started to increase slightly in the J. xuanhanensis and J. granti zones to a maximum of 32 °C (Fig. 2.4). From the C. postbitteri hongshuiensis to C. postbitteri postbitteri Subzone, temperatures decreased to 26 °C. Starting with the C. dukouensis Zone temperatures increased again with relatively high temperatures between 33 and 36 °C recorded in the C. leveni and C. liangshanensis zones. A comparable pattern is observed in the Penglaitan section. Temperatures between 26 and 29 °C are calculated from gondolellids of the J. shannoni and J. xuanhanensis Zone. Gondollelids from the J. xuanhanensis and J. granti zones indicate temperatures between 29 and 30 °C. In the C. postbitteri postbitteri Subzone to C. asymmetrica Zone, gondollelid δ18O values suggest cooling with minimum palaeotemperatures of 23 °C (Fig. 2.4). Maximum temperatures around 33 to 36 °C reconstructed from gondollelids and hindeodids seem too warm if compared to modern tropical sea surface temperatures. Since calculated temperatures are dependent on the

44 Chapter 2 Guadalupian–Lopingian boundary oxygen isotopic composition of Permian sea water, the unrealistically warm temperatures suggest that the assumed oxygen isotopic composition of sea water in the Yunnan–Guizhou–Guangxi Basin may have been different than -1‰ VSMOW. Due to this uncertainty, we only focus on relative palaeotemperature changes assuming no major change in the oxygen isotope ratio of sea water during this relatively short time interval.

Fig. 2.4 Composite oxygen isotope record of conodonts from the Penglaitan and Tieqiao sections. Palaeotemperature for gondollelid (filled symbols) and hindeodid (open symbols) δ18O values calculated assuming an oxygen isotopic composition of Permian sea water of -1‰ VSMOW. Sea level changes after Mei et al. (1999a, b), duration of Emeishan volcanism after (Ali et al., 2002; Sun et al., 2010), absolute ages after (Gradstein et al., 2004). Explanation for stratigraphic units 1 to 5 see Fig. 2.2 Blue line – 3 point running mean. Oxygen isotopes of gondollelid conodonts from the Tieqiao and Penglaitan section indicate warming of water temperatures of about 4 °C in the

45 Chapter 2 Guadalupian–Lopingian boundary late Capitanian (J. postserrata to J. granti Zone), cooling of 6 to 8 °C across the G–L boundary and in the early Wuchiapingian (C. postbitteri postbitteri to C. asymmetrica Zone) and significant warming in the later Wuchiapingian (C. leveni to C. liangshanensis Zone). Warming in the late Capitanian could be related to the eruption of the Emeishan Large Igneous Province (LIP) which is extending over > 0.5 × 106 km2 in southwest China (Courtillot and Renne, 2003). Initial Emeishan volcanism was dated into the middle Capitanian J. altudaensis Zone (Sun et al., 2010). Large scale and widespread eruptions occurred in the J. xuanhanensis Zone (Sun et al., 2010) that could have contributed to the gradual temperature increase in the J. xuanhanensis to J. granti Zone (Fig. 2.4). Volcanic eruptions result in the emission of a range of volcanic gases with CO2 and SO2 being volumetrically the most important ones

(Sigurdsson, 1990). However, CO2 quantities released during basaltic flood eruptions are assumed to have been relatively small in comparison to atmospheric CO2 levels and may thus have had only a limited effect on global climate (Caldeira and Rampino, 1990; Self et al., 2006). Sediment derived CO2 contributions from contact metamorphism of carbonates, organic-carbon rich sediments and petroleum-bearing evaporates (Ganino and Arndt, 2009; Svensen et al., 2009) and/or destabilisation of methane hydrates contained in permafrost soils (e.g. Retallack and Jahren, 2008) were advocated to contribute additional quantities of greenhouse gases that could amplify climatic warming. However, organic carbon and methane have low to very low carbon isotope ratios and a significant addition of isotopically light carbon from contact metamorphism of organic matter or gas hydrate destabilization is not evident from the carbon isotope record of the Penglaitan and Panthalassan sections. The significant temperature decrease of 6–8 °C across the G–L boundary and in the early Wuchiapingian (uppermost C. postbitteri hongshuiensis to C. asymmetrica Zone; Fig. 2.4) postdates the main Emeishan eruptions as, for

46 Chapter 2 Guadalupian–Lopingian boundary example, Ali et al. (2002) and Sun et al. (2010) argued that the main phase of Emeishan volcanism ceased either before or around the G–L boundary. The post-eruption weathering of the continental effusive basalts may have contributed to climatic cooling by lowering atmospheric pCO2 (e.g. Berner et al., 1983). Silicate weathering consumes atmospheric CO2 by converting CO2 into bicarbonate, transferring bicarbonate by rivers to the ocean, and precipitating and depositing it as marine carbonate. Since basalts are more easily weathered than other crystalline silicate rocks (Dessert et al., 2003), weathering of large volumes of basalt under warm climatic conditions may have acted as a strong CO2 sink, potentially inducing climatic cooling. However, the observed temperature decrease is relatively large in comparison to cooling predicted by numerical modeling through chemical weathering of large igneous provinces. For example, weathering of the much larger Deccan Traps (2~3 ×106 km3 vs. 0.3 ×106 km3 for Emeishan basalts; Ali et al., 2005) was modeled to have resulted in a 20% reduction of atmospheric pCO2 causing global cooling of 0.6 °C (Dessert et al., 2001). Weathering of Siberian Trap basalts was estimated to have induced cooling of 1.1 °C (Grard et al., 2005).

Precipitation and temperature are the key parameters determining CO2 consumption during basalts weathering. Since Emeishan basalts erupted in low latitudes (high temperature and intense precipitation), a high weathering rate is expected (Dessert et al., 2003). However, it is questionable whether weathering of the comparatively small Emeishan basalt province was responsible for the observed cooling. In summary, the significant temperature changes in the latest Capitanian can hardly be explained the eruption and subsequent weathering of Emeishan basalts. An alternative explanation for the observed variations in δ18O and palaeotemperature are changes in water depth in conjunction with Capitanian and early Wuchiapingian eustatic sea-level changes. According to Shen et al.

47 Chapter 2 Guadalupian–Lopingian boundary

(2007) and Wignall et al. (2009b), shallowing is observed in the Capitanian with the base of the Laibin limestone member being interpreted as a sequence boundary. The carbonates of the Laibin limestone member are interpreted as lowstand deposits indicating a more proximal position of the Penglaitan (upper slope to basin) and a more distal position of the Tieqiao location (base of slope to basin). The palaeotemperature increase documented in lower δ18O values of gondollelids in the J. xuanhanensis to J. granti Zone coincides with shallowing and assumed reduction of water depth. A transgressive surface is placed at the top of bed 5b at Penglaitan and in the upper part of bed 5 at Tieqiao with substantial deeping occurring in the overlying beds in both sections (Wignall et al., 2009b). Interestingly, palaeotemperatures decrease significantly above Bed 6 in Tieqiao and bed 5 in Penglaitan suggesting that the cooler temperatures may reflect increasing water depth. Finally, increasing temperatures in the Wuchiapingian (C. postbitteri postbitteri Subzone to C. liangshanensis Zone) coincide with shallowing and the development of platform and reef carbonates (Shen et al., 2007). In summary, the coincidence of sea-level variations and temperature changes reconstructed from gondollelid δ18O is in support of our interpretation that Capitanian and Wuchiapingian gondollelids lived in subsurface waters. During sea-level rises, gondollelids probably migrated to deeper waters and recorded cooler temperatures whereas during shallowing episodes, gondolellids were forced to thrive in shallower waters with higher temperatures.

2.5 Conclusions

This study supports earlier views of Permian gondollelids having lived in subsurface and thus cooler waters (Orchard, 1996; Lai et al., 2001). Whether gondolellids were thriving within the water column or near to the sea bottom

48 Chapter 2 Guadalupian–Lopingian boundary cannot be decided. Reconstructed habitat depth of gondolellids from the Penglaitan and Tieqiao sections indicate water depths of 60 to 80 m. Water depths in the Yunan–Guizhou–Guangxi Basin were estimated to have been around 200 m during the Capitanian (Wang and Jin, 2000), however, no water depth estimates are available for the Laibin limestone member. Since gondollelids are interpreted to have lived in deeper waters, oxygen isotopes measured on Permian gondollelid conodonts may reflect a combination of changes in water depths and palaeoclimatic changes. Hindeodid conodonts, interpreted to have lived in surface and/or shallow waters, are considered a better option to reconstruct the surface water temperatures of Middle and Late Permian seas. However, hindeodids are not very abundant in many platform successions and thus will not allow the reconstruction of high-resolution oxygen isotope records. This study illustrates that, at least in the Permian, oxygen isotopes studies on conodont apatite should be performed on mono-generic conodont samples and that oxygen isotope analyses of monogeneric conodont samples have a great potential to shed additional light on the life habitat of conodont

49

50 Chapter 3 Permian–Triassic boundary

Chapter 3: Climatic warming in the latest Permian and the Permian–Triassic mass extinction

This chapter is lagerly based on Joachimski, M.M., Lai, X.L., Shen, S.Z., Jiang, H.S., Luo, G.M., Chen, B., Chen, J., Sun, Y.D., 2012. Climate warming in the latest Permian and the Permian–Triassic mass extinction. Geology 40, 195–198.

Abstract

High-resolution oxygen isotope records document timing and magnitude of global warming across the Permian–Triassic boundary (PTB). Oxygen isotope ratios measured on phosphate-bound oxygen in conodont apatite from the Meishan and Shangsi sections (South China) decrease by 2‰ in the latest Permian translating into warming of low latitudinal surface waters of 8 °C. The oxygen isotope shift coincides with the negative shift in carbon isotope ratios of carbonates suggesting that the addition of isotopically light carbon to the ocean-atmosphere system by Siberian trap volcanism and related processes resulted in higher greenhouse gas levels and global warming. The coincidence of climate warming and the main pulse of extinction suggest that global warming was one of the causes of the collapse of the marine and terrestrial ecosystems. In addition, very warm climate conditions in the Early Triassic may have played a major role in the delayed recovery in the aftermath of the Permian–Triassic crisis.

3.1 Introduction

The largest biotic extinction event in Earth History occurred at the Permian–Triassic boundary (PTB) affecting both marine and terrestrial life (e.g. Erwin, 1993). The causes of this mass extinction event have been discussed controversially and include among others the effects of Siberian Trap volcanism

(Renne et al., 1995), oceanic anoxia (Wignall and Hallam, 1992), oceanic H2S

51 Chapter 3 Permian–Triassic boundary degassing (Kump et al., 2005) as well as climatic changes caused by increasing atmospheric CO2 concentrations due to volcanism (Erwin, 1993; Korte et al., 2010) and thermal decomposition of organic matter (Svensen et al., 2009). Climatic changes, especially climatic warming in conjunction with elevated atmospheric CO2 levels as consequence of Siberian trap volcanism and contemporaneous volcanism in South China (Yin et al., 1992) were suspected to have played a major role during the Permian–Triassic transition. However, up to date no proxy data were available that document timing and magnitude of global warming during this critical time interval. We present the first high-resolution oxygen isotope paleotemperature records from two Chinese PTB sections. Oxygen isotopes were measured on conodonts, tooth-like microfossils composed of carbonate-fluorapatite, representing the only skeletal remains of the marine conodont animal assumed to have been one of the earliest jawless vertebrates (Donoghue et al., 2000). Oxygen isotopes analyzed on phosphate-bound oxygen of conodont apatite have been shown to be relatively inert to diagenetic alteration and to be a promising tool for reconstructing paleoclimatic changes (e.g. Joachimski et al., 2009). The presented oxygen isotope records document significant climatic warming in the latest Permian that is seen in context with major changes in the global carbon cycle.

3.2 Geological setting

Oxygen isotopes were studied on mono-generic conodonts from the Meishan (Zhejiang Province) and Shangsi (Sichuan Province) sections in South China. During the Late Permian, both sections were located on the Yangtze Platform at 20° N in subtropical latitudes (Liu et al., 1999) at a distance of 1.500 km from each other. The Meishan section represents the Global Stratotype Section and Point (GSSP) of the Permian–Triassic boundary (PTB) and has been

52 Chapter 3 Permian–Triassic boundary intensively studied by various research groups (for summary see Yin et al., 2001). The Changhsingian to Griesbachian strata were deposited in an intraplatform depression with water depths being shallower in comparison to other PTB sections in south China (e.g. Yin et al., 1995). The Changxing Formation (uppermost Wuchiapingian to Changhsingian) is represented by siliceous bioclastic micrite with thinly-interbedded cherts. The Yinkeng Formation (uppermost Changhsingian to Griesbachian) is composed of shales with minor intercalation of thin-bedded limestones. In the Shangsi section, the latest Permian (Wuchiapingian to Changhsingian) is represented by dark, organic-rich carbonate mudstones overlain by cherty micrites. Early Triassic (Griesbachian) strata are composed of thin-bedded alternations of micritic and argillaceous limestones with no cherty concretions. The hemipelagic facies of the Shangsi section indicates a deeper water to basinal depositional setting (Wignall et al., 1995). Conodonts play an important role for the correlation of PTB sections with the first appearance datum (FAD) of Hindeodus parvus marking the basal Triassic (Yin et al., 1988). In the Meishan section, the PTB is defined at the base of bed 27c (Fig. 3.1). The position of the PTB in the Shangsi section is more controversial. Lai et al. (1996) defined the boundary at the upper part of bed 28a based on the occurrence of Hindeodus turgidus at this level. Jiang et al. (2011) observe the first occurrence of Hindeodus parvus and thus the base of the Hindeodus parvus Zone at the base of bed 29c. However, based on the occurrence of N. taylorae (an auxiliary marker for the recognition of the PTB) in bed 28a, the authors suggest that the PTB should be defined within bed 28 and thus below the first occurrence of Hindeodus parvus. Oxygen isotopes were analyzed on platform “tooth” elements (P1) of conodont taxa Clarkina (or Neogondolella) and Hindeodus in case of abundant faunas (see table 3.1). Ramiform “tooth” elements of both taxa were used in case

53 Chapter 3 Permian–Triassic boundary of less abundant specimens. Clarkina is generally interpreted as an offshore, outer shelf, basinal or deep water taxon (e.g. Wardlaw and Collinson, 1984; Lai et al., 2001). The paleoecology of Hindeodus is more controversial. While some authors consider Hindeodus a shallow-water dweller (Wardlaw and Collinson, 1984; Orchard, 1996), others interpreted Hindeodus to have lived in varied environments (e.g. Kozur et al., 1996; Lai et al., 2001). In the Meishan and Shangsi sections, the abundance of Clarkina declines significantly toward the PTB with Hindeodus becoming more abundant toward the boundary and being the dominant taxon in the Early Triassic (Lai et al., 2001).

3.3 Results

Oxygen isotopes of Clarkina from the base of the studied interval (beds 22–23) in Meishan are between 21.3 and 22.3‰. Lower values between 20.7and 21.6‰ are measured from the base of bed 24 to bed 26 (Fig. 3.1). In this interval, oxygen isotope ratios of Hindeodus compare well to those of Clarkina with a maximum difference in δ18O of 0.4‰. Starting from bed 27, oxygen isotopes start to decrease, both for Clarkina and Hindeodus. Oxygen isotopes for Clarkina decrease to values of 20.4 and 20.0‰ while oxygen isotopes measured on Hindeodus decrease to values between 19.3 (beds 27a to 28) and 18.8‰ (bed 29). A comparable pattern is observed in the Shangsi section (Fig. 3.2). Oxygen isotope values measured on Clarkina from beds 26 and 27a are around 21.5‰. δ18O values start to decrease at bed 27c with values of 20.0‰ (Clarkina) and 19.5‰ (Hindeodus) measured in bed 27c and 28c, respectively. Hindeodids from beds 30–32 show δ18O values from 18.9 to 19.3‰, hindeoids from bed 33 have a value of 18.7‰. In general, P1 and ramiform conodont elements of both studied taxa, Clarkina and Hindeodus, show identical δ18O values given an analytical reproducibility of ± 0.2‰ (1σ) suggesting that both elements can be

54 Chapter 3 Permian–Triassic boundary used for paleotemperature reconstruction. The well comparable absolute δ18O values as well as δ18O patterns in the Meishan and Shangsi sections suggest that the recorded trends can neither be explained by local variations in salinity nor by diagenetic alteration.

Fig. 3.1 Oxygen isotopes of conodont apatite (this study) and carbon isotopes of whole rock carbonate (red dots - Jin et al., 2000; black dots - Xie et al., 2007) for Meishan C. Anoxic facies after Wignall and Hallam (1992). Paleotemperatures calculated assuming a δ18O value for Permian sea water of -1‰ VSMOW. TR–Triassic, PTB–Permian–Triassic boundary, SB– sequence boundary, TS–transgressive surface, 1–limestone, 2–marly limestone, 3–claystone, 4–siliceous marl, 5–black shale, 6–volcanic ash, 7–chert, 8–dysaerobic to anaerobic facies.

3.4 Discussion

Paleotemperatures were calculated using the equation given by Pucéat et al. (2010) assuming a δ18O value of -1‰ VSMOW for Late Permian seawater (a

55 Chapter 3 Permian–Triassic boundary value typically assumed for non-glacial time intervals; Savin, 1977) and no secular variations in seawater salinity. In the Meishan section, oxygen isotope ratios of Clarkina from beds 22–23 translate into average temperatures of 22 °C (Fig. 3.1). Slightly higher temperatures between 23–27 °C are registered in the overlying beds 24–26. In this interval, Clarkina and Hindeodus P1 elements indicate comparable temperatures. Starting with bed 27a, paleotemperatures increase with Clarkina showing a minor temperature increase in comparison to Hindeodus. While Hindeodus indicates warming of 8 °C to maximum paleotemperatures of 33–35 °C, paleotemperatures reconstructed from δ18O of Clarkina indicate warming of 5 °C with maximum paleotemperatures ranging between 29 and 30 °C. A comparable paleotemperature trend is reconstructed in the Shangsi section (Fig. 3.2). Paleotemperatures calculated from δ18O of Clarkina indicate temperatures between 23 and 24 °C in the latest Permian and increase to 32 °C above the event horizon. δ18O values of Griesbachian Hindeodus specimens translate into temperatures between 32 and 35 °C, comparable to paleotemperatures reconstructed from δ18O measured on hindeodids from Meishan. The well-comparable oxygen isotope ratios of Hindeodus and Clarkina in the Meishan section (bed 24a–24e) indicate that Hindeodus and Clarkina were thriving in comparable water depths with comparable temperatures. Sediments of the Meishan section were deposited in an intraplatform depression with estimated water depths of around 50 m for the Permian–Triassic transition (Zhang et al., 1996). In these water depths, temperatures are not expected to be very different from surface water temperature. Consequently, we argue that both taxa, Hindeodus and Clarkina, recorded surface water temperature. The Shangsi section represents a platform to slope setting and potentially a significantly deeper environment in comparison to Meishan. Nevertheless, δ18O values of Clarkina are well comparable to those of Meishan supporting our interpretation

56 Chapter 3 Permian–Triassic boundary that Clarkina and Hindeodus from both sections lived in near surface-waters and recorded surface water temperature. This interpretation is further supported by calculated warm water temperatures.

Fig. 3.2 Oxygen isotopes of conodont apatite and carbon isotopes of whole rock carbonate (Yan et al., 1989) for Shangsi. Paleotemperatures calculated assuming a δ18O value for Permian sea water of 1‰ VSMOW. Anoxic facies after Wignall et al. (1995). Position of Permian–Triassic boundary (PTB) based on first occurrence of Hindeodus parvus or occurrence of Neogondolella taylorae (PTB*), an auxiliary marker for the recognition of the PTB (see Jiang et al., 2011). Legend for lithologies see Fig. 3.1. The most striking feature of the conodont δ18O record is the significant temperature rise in the latest Permian. While Hindeodus suggests warming of 8 °C, Clarkina indicates a slightly minor temperature increase of 4–5 °C (Fig. 3.1). The increase in temperature cannot be attributed to shallowing and

57 Chapter 3 Permian–Triassic boundary warming of surface waters since a sea level rise and thus deepening of the environment was identified above bed 24d in the Meishan section (Zhang et al., 1996; Cao et al., 2010). The minor temperature increase recorded by Clarkina in comparison to Hindeodus suggests that this taxon migrated to slightly deeper and thus colder waters in conjunction with this sea level rise in the latest Changhsingian (e.g. Wignall and Hallam, 1992) and/or the development of a stratified and thermally less well-mixed water column. Reconstructed climatic warming is supported by latest Permian and earliest Triassic paleosols indicating significant warming across the PTB (Retallack, 1999). Warm Early Triassic climatic conditions and a low latitudinal temperature gradient are further indicated by the latitudinal distribution of ammonoids (Galfetti et al., 2007) and intensively weathered high latitudinal paleosols (Retallack and Krull, 1999). Global warming will result in minor

O2-contents of surface waters, increased respiration rates of organic carbon and potentially a more sluggish oceanic circulation culminating in widely documented shallow water anoxia (e.g. Wignall and Hallam, 1992; Knoll et al., 2007). This context is corroborated by the coincidence of climatic warming and the onset of oxygen-poor facies in Meishan (bed 25, Fig. 3.1) and Shangsi (bed 28a, Fig. 3.2). Since temperature is considered as a main controlling factor for the formation of microbial carbonates (Riding, 1992), warming of surface waters will have favored the widespread occurrence of cyanobacterial microbialites after the PT extinction (e.g. Xie et al., 2010). Reconstructed climatic warming in the latest Permian is seen in context with a major perturbation in the carbon cycle mirrored in the globally recorded negative carbon isotope excursion (see reviews in Korte and Kozur, 2010 and 18 13 13 Hermann et al., 2010). The apatite δ O and carbonate δ C records (δ Ccarb) from Meishan and Shangsi suggest parallel negative shifts at the transition of beds 23/24 (Meishan) and at the extinction horizon (Meishan and Shangsi). In

58 Chapter 3 Permian–Triassic boundary

13 the Meishan section, δ Ccarb values show a gradual decrease starting at the transition of bed 23/24 and a second more abrupt negative shift in the upper part of bed 24e below the event horizon (Fig. 3.1; Jin et al., 2000; Cao et al., 2002; 13 Xie et al., 2007). δ Ccarb values recover in beds 26–28 but reveal a second negative shift in beds 28–34 (early Griesbachian; Xie et al., 2007). In the Shangsi 13 section, the first negative shift in δ Ccarb starts in bed 22, the second more sudden shift occurs immediately below or at the event horizon (Fig. 3.2; Yan et al., 1989). Outgassing of CO2 during Siberian trap volcanism (Renne et al., 1995, Sobolev et al., 2011), thermal decomposition of isotopically light organic carbon around sill intrusions (Svensen et al., 2009), destabilization of 12C-enriched methane hydrates contained in permafrost soils (e.g. Retallack and Jahren, 2008) and a higher contribution of methane to the ocean-atmosphere system (Luo et al., 2010) were discussed as triggering mechanisms for the negative carbon isotope excursion. All of these processes would have resulted in higher atmospheric greenhouse gas levels and may have initiated the dramatic warming documented by the oxygen isotope records of conodont apatite. The main extinction of the Permian–Triassic crisis occurred during a relatively short time interval represented by beds 25 to 28 in Meishan (Shen et al., 2011); the highest extinction rate is observed at the base of bed 25 (Jin et al., 2000; Yin et al., 2007). Paleotemperatures start to increase below bed 25, with maximum and probably harmful temperatures reached at the base of bed 27. The coincidence of climate warming and the major phase of extinction indicate that rapid climate warming contributed to the collapse of both marine and terrestrial ecosystems. Additionally, as pointed out by Knoll et al. (2007), the very warm climate conditions in the Early Triassic may have played an important role in the timing and pattern of the slow recovery of marine and terrestrial ecosystems in the aftermath of the Permian–Triassic crisis.

59

60 Chapter 4 Permian composite oxygen isotope record

Chapter 4: Permian ice volume and palaeoclimate history: oxygen isotope proxies revisited

Main parts of this chapter were submitted for publication to Gondwana Research: Chen B., Joachimski Michael M., Shen S.Z., Lambert Lance, Lai X.L., Wang X.D., Chen J., Yuan D.X., Permian ice volume and palaeoclimate history: oxygen isotope proxies revisited. Abstract

A high-resolution oxygen isotope record based on 356 measurements of conodont apatite from several low latitudinal sections in South China, USA and Iran was composed in order to unravel Permian palaeotemperature and ice volume history. The conodont apatite δ18O record is compared to published brachiopod calcite δ18O records. Brachiopods and conodonts from different palaeocontinents show significantly different δ18O values, suggesting differences in local climatic conditions (e.g. evaporation/precipitation ratio). As a consequence, secular changes in palaeotemperature and oxygen isotope composition of Permian sea water cannot be reconstructed from records combined from different areas, but have to be based on records from a specific area. Oxygen isotope analyses of different conodont taxa suggest that Streptognathodus and Hindeodus lived in near-surface seawater and recorded surface water temperature, whereas the habitat of gondolellid genera was variable depending on sea level, with both near-surface and deeper waters as potential life habitat. The oxygen isotope record measured on conodonts from South China exhibits relative high values between 22 and 23‰ VSMOW during the glaciated Early Permian, translating into warm seawater temperature between 26 and 30 °C, assuming that the Late Palaeozoic ice volumes were comparable to the Pleistocene glacial maxima. In contrast to the earlier view that the Late Palaeozoic Ice Age (LPIA) terminated in the late Sakmarian, the

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South China conodont apatite oxygen isotope record suggests waning of the ice sheets in the Kungurian. Ice melting is indicated by a pronounced decrease in δ18O of 2‰ VSMOW, which is interpreted as reflecting the combined effect of climatic warming and glacial ice melting. Significant temperature fluctuations (4 °C warming succeeded by 6 to 8 °C cooling) are observed during the Guadalupian–Lopingian transition, interpreted as combined climate changes induced by Emeishan volcanism and changes in habitat depth of gondollelid conodonts. Oxygen isotope values increase to 22‰ VSMOW in the Changhsingian, which suggests climate cooling and Clarkina moving to deeper waters because of the Changhsingian sea level rise. Across the Permian–Triassic boundary, δ18O values decrease from 22 to 19‰ VSMOW, parallel to the significant negative carbon isotope excursion and the eruption of the Siberian Traps. The latter is pointing to a cause-effect relationship as a consequence of the massive release of volcanic greenhouse gases derived from the Siberian volcanism and related processes.

4.1 Introduction

The Permian (298.9 to 252.2 Ma) represents a transitional interval between the Paleozoic and Mesozoic eras in which global tectonic, environmental and faunal communities underwent dramatic changes that had a profound impact on the subsequent evolution of the Earth system. Following the final collision of Laurussia and Gondwana in the Late Carboniferous (Stampfli and Borel, 2002), the assembly of the supercontinent Pangea was accomplished with the amalgamation of Siberia during the Uralian orogeny (Scotese, 2001). The Palaeothethys was narrowing during the Permian as a consequence of the opening of the Neothethys, a new oceanic basin that resulted from the northward-drifting Cimmerian terranes, which had split off from Gondwana (Fig. 4.1). The formation of the supercontinent Pangea, as well

62 Chapter 4 Permian composite oxygen isotope record as its later breakup, has been suggested to have played an important role in the origin of the Late Palaeozoic Ice Age (LPIA) and global scale environmental changes (Link, 2009). In contrast to earlier interpretations arguing that the LPIA was a single, widespread glacial event spanning the late Mississippian to Early Permian (Veevers and Powell, 1987), recent studies have shown that the LPIA can be differentiated into discrete glacial intervals separated by warmer periods of time. The Permian acme of the LPIA was dated as Asselian to early Sakmarian (Isbell et al., 2003; Fielding et al., 2008a, b, c), with the ice caps probably waning in the Sakmarian (Veevers and Powell, 1987; Isbell et al., 2003; Korte et al., 2005; Zeng et al., 2012). However, the exact timing of this deglaciation is still controversial since glacial deposits are found to appear intermittently in late Early and Middle Permian strata in Australia and Siberia (Fielding et al., 2008c). The late Middle through latest Permian are characterized by the end-Guadalupian (259.0 Ma) and end-Permian (252.28 Ma) mass extinction events (Shen et al., 2010, 2011). Both events significantly affected global ecosystems, with 90% of all marine species becoming extinct, and a dramatic decline in terrestrial diversity registered during the end-Permian mass extinction (Retallack, 1995; Erwin et al., 2002). The triggering mechanisms for these mass extinctions are not clear. However, important environmental fluctuations, for example major sea level changes (Haq and Schutter, 2008; Wignall et al., 2009b; Chen et al., 1998), marine anoxia (Wignall and Twitchett, 1996; Isozaki, 1997; Wignall et al., 2009b), climate warming (Svensen et al., 2009; Shen et al., 2011; Joachimski et al., 2012) as well as climate cooling (Campbell et al., 1992; Isozaki et al., 2007a, b) or perturbations of the Earth’s carbon cycle (Chen et al., 1991; Jin et al., 2000; Cao et al., 2002, 2008; Payne et al., 2004; Wang et al., 2004; Xie et al., 2007; Korte et al., 2010; Chen et al.,

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2011; Luo et al., 2011; Shen et al., 2011) have been demonstrated to have been coincident with or to have occurred close to the mass extinction events. Interestingly, the Emeishan and Siberian Large Igneous Province are dated to have erupted during the end-Guadalupian (Zhou et al., 2002; Wignall et al., 2009a; Sun et al., 2010) and latest Permian, respectively (Campbell et al., 1992; Renne et al., 1995; Kamo et al., 2003; Mundil et al., 2004; Reichow et al., 2009; Svensen et al., 2009). The close temporal association between the volcanic events and the biotic crisis have led many researchers to assume a cause-effect relationship between volcanism and mass extinctions, with the volcanism resulting in the emission of greenhouse gases, destabilization of gas hydrates (Retallack and Jahren, 2008), and/or contact metamorphism of organic carbon-rich sediments (Ganino and Arndt, 2009; Svensen et al., 2009), all processes that potentially could have contributed to higher atmospheric greenhouse gas (CO2, CH4) levels. Climatic and environmental changes can be recorded in the stable isotope compositions of marine skeletons with variations in carbon and oxygen isotopes documenting changes in the global carbon cycle and palaeotemperature or ice volume, respectively. Brachiopod shells composed of low magnesium calcite have been preferentially used to reconstruct Permian oxygen isotope records. However, the reconstruction of global carbon and oxygen isotope records using brachiopod shells is complicated by a potential diagenetic alteration of shell calcite, local environmental effects and the lack of temporal- and spatial coverage (Grossman et al., 2008). Hence, an alternative proxy is required to reevaluate and refine the oxygen isotope signal from Permian oceans. Conodonts are tooth-like microfossils that represent the skeletal remains of the marine animal assumed to have been one of the earliest jawless vertebrates (Donoghue et al., 2000). They are widespread in Permian marine sediments and are used for the biostratigraphic subdivision of the

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Permian (e.g. Jin et al., 1997; Henderson and Mei, 2003; Shen and Mei, 2010). Due to their carbonate-fluorapatite mineralogy and the dense crystalline structure, they have a high potential to preserve primary oxygen isotope ratios (Pucéat et al., 2004). Consequently, conodont apatite has been considered as a robust proxy to reconstruct the palaeotemperature history of the Palaeozoic ocean (Joachimski and Buggisch, 2002; Joachimski et al., 2006; Buggisch et al., 2008; Trotter et al., 2008; Joachimski et al., 2009; Buggisch et al., 2010) We present a high-resolution oxygen isotope record based on conodont apatite from several low latitudinal Permian sections in South China, USA and Iran. The conodont apatite δ18O data presented here are compared with published coeval brachiopod calcite δ18O records. The Permian oxygen isotope records will be discussed in the context of ice volume and palaeoclimatic changes, as well as major volcanic and mass extinction events.

4.2 Geological setting

Samples were collected from sections in South China (Early Permian to Early Triassic), USA (Middle Permian) and Iran (Late Permian). During Permian time, the South China block was located in the eastern part of the Palaeotethys. The Guadalupe Mountains (Texas, USA) were situated in the western part of Euramerica, whereas the Abadeh section (Iran) was located on the Sanandaj-Sirjian microcontinent, which was a part of the northward drifting Cimmerian terranes (Kozur, 2007; Fig. 4.1). While the studied localities were located on different palaeocontinents, all were located in the low latitude tropical zone and characterized by continuously deposited marine sediments with abundant marine faunas. Early Permian conodont samples were collected from the Naqing section located about 44 km southwest of the city of in Luodian Guizhou Province, Southwest China (Fig. 4.1). The section is characterized by a continuous

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Carboniferous to Permian sedimentary succession with an abundant conodont fauna. The Permian succession of the Naqing section is dominated by dark–grey and grey limestones with frequent silicified or cherty horizons (Mei et al., 1999a; Henderson and Mei, 2003) deposited in a slope setting of the Dian-Qian-Gui basin, a deep-water basin within the South China Block (Fig. 4.1). Conodont biostratigraphy of the Naqing section was established by Mei et al. (2001) and refined by Henderson and Mei (2003).

Fig. 4.1 (A) Late Permian palaeogeographic reconstruction showing location of the Guadalupe Mountains (USA-1), South China Block (-2), and Abadeh section (Iran-3) (http://www2.nau.edu/rcb7/260moll.jpg). (B) Palaeogeography of South China during the Wuchiapingian (modified after Wang and Jin, 2000 and Shen et al., 2007; palaeolatitude after Zhang, 1997 and Fang et al., 1992), location of studied sections and the modern outcrop area of Emeishan volcanic province (Ali et al., 2002). Middle to Late Permian samples were mainly collected from South China, where Lopingian strata are represented by a complete marine depositional succession that contains highly diverse marine faunas (Jin et al., 1997; Shen and Mei, 2010). The studied Tieqiao and Penglaitan sections in laibin, Guangxi Province (Fig. 4.1) are characterized by Guadalupian to early Lopingian carbonates and radiolarian cherts that were deposited continuously in a slope to basin setting. Both sections are well-dated by means of conodont biostratigraphy (Mei et al., 1998; Jin et al., 2006a; Shen et al., 2007), with the Penglaitan section as the Global Boundary Stratotype Section and Point (GSSP)

66 Chapter 4 Permian composite oxygen isotope record for the base of Wuchiapingian and which is considered as one of the most complete Guadalupian–Lopingian boundary sections worldwide (Mei et al., 1998; Jin et al., 2006a). In most other shelf or platform successions around the World, the Capitanian–Wuchiapingian regression (Haq and Schutter, 2008) resulted in erosional unconformities (Jin et al., 2006a). The Dukou section in northern Sichuan Province (Fig. 4.1) exposes Lopingian slope to basin carbonates alternating with cherts. Conodont biostratigraphy indicates continuous deposition from the Wuchiapingian Clarkina dukouensis conodont Zone to lowest Triassic (Mei et al., 1994; Cao et al., 2010). Changhsingian to Griesbachian conodonts were collected from the Meishan section in Zhejiang Province (Fig. 4.1), which is composed of carbonates and shales deposited in a slope setting. The Meishan D section represents the GSSP for the base of the Changhsingian (Jin et al., 2006b; Wang et al., 2006) and the Permian–Triassic boundary (PTB; Yin et al., 2001). The Shangsi PTB section in Sichuan Province (Fig. 4.1) exposes siliceous limestones with interbedded shales deposited in a relatively deeper depositional setting in comparison to the Meishan section. Conodont biostratigraphy of the Shangsi section was intensively studied by Zhang et al. (1984), Li et al. (1989), Lai et al. (1996), Nicoll et al. (2002), Jiang et al. (2011), and Shen et al. (2011). Late Permian conodonts from platform and basin sediments were also studied from the Jiangya section in Hunnan Province (Fig. 4.1). For comparison, the Middle Permian (Guadalupian) was sampled in the Guadalupe Mountains of West Texas/USA (Fig. 4.1), where the GSSPs of the Guadalupian Series were established using the evolutionary lineages of the conodont genus Jinogondolella (Wardlaw et al., 2004). The Middle Permian strata exposed in the Guadalupe Mountains reflect a complete facies transect from the shoreline to the basin. Lithologies comprise shallow-water and reef carbonates on the shelf, intercalated siliciclastic and platform derived

67 Chapter 4 Permian composite oxygen isotope record carbonates on the slope, and mainly sand- and siltstones in the basin (Glenister et al., 1999; Wardlaw et al., 2000). The basinal siliciclastic deposits are considered to have been deposited continuously (Glenister et al., 1999), with the interbedded limestone units containing an abundant and well-preserved conodont fauna (Wardlaw et al., 2004). Both the siliciclastic and carbonate units are prolific petroleum reservoirs in the nearby subsurface, and the entire region is referred to as the Delaware Basin. Late Permian conodonts were further studied from the relatively deep water/open ocean section of Abadeh in central Iran. The Late Permian and Early Triassic succession of Abadeh has been intensively investigated by means of biostratigraphy (Taraz, 1971, 1973; Kozur et al., 1975, 1978; Taraz et al., 1981; Iranian-Japanese Research Group, 1981; Kozur, 2004, 2005, 2007; Shen and Mei, 2010), lithology (Heydari et al., 2003) and geochemistry (Korte et al., 2004, 2010). The conodont succession of the Abadeh section is considered to represent one of the most detailed zonations of the Permian–Triassic transition in the Tethys (Kozur, 2004, 2005, 2007) and considerable attention has been paid to conodont biostratigraphy due to controversial views as to whether this section represents a continuous open marine sedimentary sequence with complete conodont evolutionary series across the PTB (Sweet and Mei, 1999; Gallet et al., 2000; Heydari et al., 2001; Kozur, 2004, 2005, 2007). The stratigraphic range of the sampled intervals in sections from South China, USA and Iran are presented in Fig. 4.2.

4.3 Results

The conodont apatite composite oxygen isotope record is based on a total of 356 analyses (Fig. 4.3 and Table 4.1). 296 samples spanning the entire Permian were collected from various sections in South China. Additional Middle and Late Permian samples were collected from the Guadalupe

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Mountains of West Texas/USA (n=27) and the Abadeh section/Iran (n=33), respectively. Data for the Guadalupian–Lopingian (Tieqiao and Penglaitan, South China; n=64) and Permian–Triassic boundary sections (Meishan and Shangsi; n=53) have been previously published by Chen et al. (2011) and Joachimski et al. (2012), respectively.

Fig. 4.2 Stratigraphic range of sections studied in South China, USA and Iran. Absolute ages after Gradstein et al. (2004). Alternate ages (top) after Schmitz and Davydov (2012) and Shen et al. (2011). The oxygen isotope record based on conodonts from South China shows average values slightly above 22.0‰ for the Early Permian. Average δ18O values decrease from 22.0‰ in the Kungurian to 20.0‰ in the late Capitanian. A positive excursion with values increasing from 19.8 to 20.8‰ is observed in the Capitanian–Wuchiapingian transitional interval. In the early Wuchiapingian, δ18O average values decrease from 21.0‰ to minimum values of 19.5‰, followed by a prominent increase from 19.5 to 22.0‰ in the late Wuchiapingian to early Changsinghan. δ18O values for late Changhsingian and early Triassic conodonts from the well known PTB sections at Meishan and

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Shangsi show a significant decrease with values of 22.0‰ in the late Changhsingian, decreasing to 19‰ in the early Triassic.

Fig. 4.3 Composite oxygen isotope curve of conodont apatite (n=356) from sections in South China, USA, and Iran. Bold line represents Locfit regression calculated for samples from South China, dotted lines give 95% confidence interval. Absolute ages after Gradstein et al. (2004). Alternate ages (top) after Schmitz and Davydov (2012) and Shen et al. (2011). The Middle Permian (Guadalupian) conodonts from the Guadalupe Mountains (West Texas/USA) exhibit slightly heavier δ18O values in comparison to conodonts from south China, with values ranging from 20.8 to 22.0‰. Late Permian samples from the Abadeh section (Iran) show δ18O values of 19.8 to 21‰. The δ18O values from Iran agree well with those from South China in the late Wuchiapingian and latest Changhsingian, but are depleted in δ18O in the Changhsingian.

4.4 Discussion

The oxygen isotope composition of biogenic calcite or apatite is dependent

70 Chapter 4 Permian composite oxygen isotope record on temperature and the oxygen isotopic composition of the sea water in which an organism lived. Since the latter is dependent on variations in salinity and continental ice volume, oxygen isotopes can also be used to reconstruct changes in salinity and ice volume. Knowledge concerning the habitat of analyzed taxa is crucial for palaeotemperature reconstructions from δ18O, since oxygen isotope ratios of surface dwelling taxa may reflect the combined effect of temperature, salinity and ice volume, whereas δ18O of taxa dwelling at greater water depths will be determined preferentially by changes in ice volume, with temperature and salinity variations becoming minor with greater water depths.

4.4.1 Permian oxygen isotope record of brachiopod calcite

Previous oxygen isotope studies focused on the analysis of well-preserved Permian brachiopod shells (Compston, 1960; Popp et al., 1986; Morante, 1996; Korte et al., 2005, 2008; Mazzullo et al., 2007; Grossman et al., 2008; Angiolini et al., 2009; Mii et al., 2012), mainly in order to constrain the timing of LPIA deglaciation. For example, Korte et al. (2005, 2008) described a 2.5‰ decrease in δ18O of Sakmarian to Artinskian brachiopods as well as bivalves from both low (Urals) and high (Australia) latitudes. A pronounced decrease in δ18O in the late Artinskian is reported as well by Mii et al. (2012), who investigated well-preserved brachiopods shells from the Sydney Basin in eastern Australia which was situated between 50–65° southern latitude during the Permian (Stampfli and Borel, 2002). However, the oxygen isotope ratios, especially of the Artinskian samples, reveal a very large variation of up to 6‰. This large variation calls into question the state of preservation for those brachiopod shells. In contrast to the studies of Korte et al. (2005, 2008) and Mii et al. (2012), no time-equivalent decrease in δ18O was observed in coeval brachiopod shells

71 Chapter 4 Permian composite oxygen isotope record from the low-latitudinal U.S. Midcontinent (Grossman et al., 2008). The melting of the Early Permian Gondwanan ice sheets would be expected to result in a lower δ18O value for Permian sea water and associated shell calcite. However, the compilation of published Permian brachiopod δ18O data from low latitudes shows no general trend in the oxygen isotope ratios (Fig. 4.4). Most Permian δ18O values range from -0.5 to -4.5‰, with a relatively high variability of 3 to 4‰ for samples from specific time intervals (Fig. 4.4). High δ18O values ranging from -0.4 to -0.2‰ for middle to late Asselian brachiopods from the southern Urals were one of line of evidence used by Korte et al. (2005) to argue for a significant decrease in δ18O, and a deglaciation in the Sakmarian. However, coeval brachiopods from Northern Iran (Angiolini et al., 2009) and the U.S. Midcontinent (Mazzullo et al., 2007; Grossman et al., 2008) are characterized by lower δ18O values between -2.3 and -3.6‰, which translates into realistic palaeotemperatures for tropical seas during peak glaciation. Angiolini et al. (2009) suggested that the high δ18O values registered in the southern Ural samples are possibly the result of excess evaporation caused by regional aridity. High Sakmarian and Artinskian δ18O values measured on brachiopods from the Urals were also interpreted to represent regional aridification and excess evaporation (Grossman et al., 2008). Discrepancies in the oxygen isotope ratios measured on Middle Permian brachiopods from the Delaware Basin (Korte et al., 2005) and Spitsbergen (Mii et al., 1997) similarly suggest regional differences in aridity. The Spitsbergen brachiopod samples were collected from higher latitudes of 45°N (Stampfli and Borel, 2002), whereas the Delaware Basin was located in near-equatorial latitudes. Brachiopods from Spitsbergen are expected to have higher δ18O values representative of lower palaeotemperatures in higher latitudes. However, the δ18O values of brachiopods from Spitsbergen are on average -3.1±0.9‰, whereas the brachiopods from the Delaware Basin have an average value of

72 Chapter 4 Permian composite oxygen isotope record

-2.1±0.9‰ (Fig. 4.4). This discrepancy was interpreted by Korte et al. (2005) to reflect significantly higher evaporation rates in the Delaware Basin. In summary, the relatively large differences in the oxygen isotope ratios of brachiopods from different areas argue for significant differences in local climate conditions (e.g. evaporation/precipitation ratio). A change in the oxygen isotope composition of Permian sea water due to the waning of Gondwanan ice sheets should be documented in a decrease in δ18O of brachiopod calcite. However, no general trend is visible if brachiopod δ18O values from different areas are plotted together. Therefore, reconstructions of secular changes in the oxygen isotope ratio of Permian sea water and/or of the palaeotemperature history should not be performed by compiling oxygen isotope data from different areas or palaeocontinents (e.g. Korte et al., 2005), but have to be based on continuous oxygen isotope records from specific areas.

Fig. 4.4 Comparison of conodont δ18O record with published oxygen isotope data for brachiopod calcite. (A) Permian brachiopod calcite δ18O data from high latitude areas; (B) Permian brachiopod calcite δ18O data from low latitude areas; (C) Permian conodont δ18O from low latitude areas (this study). Locfit regression for conodont apatite δ18O as in Fig. 4.3. Absolute ages after Gradstein et al. (2004). Alternate ages (top) after Schmitz and Davydov (2012) and Shen et al. (2011).

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74 Chapter 4 Permian composite oxygen isotope record

4.4.2 Permian oxygen isotope record of conodont apatite

4.4.2.1 Palaeoecology of studied conodont taxa

Conodont animals are generally interpreted to have been nektonic organisms, and knowledge of the life habitat for analyzed taxa is important to the interpretation of the oxygen isotope ratios. Streptognathodus, Mesogondolella, Jinogondolella and Clarkina are the most abundant taxa in the Permian samples analyzed. Hindeodus occurred only rarely, and was analyzed separately in cases where the abundances were sufficient to compose mono-generic samples. Since no significant difference in δ18O was observed between ramiform and platform elements of Clarkina or Jinogondolella from the same samples (taking into account an analytical reproducibility of ±0.2‰), both ramiform and platform elements were used for analysis. Streptognathodus is the predominant genus recovered from the earliest Permian samples in this study, and is often considered to be a Carboniferous holdover (e.g. Mei et al., 2002). Based on the distribution of conodont taxa in specific lithologies of U.S. Midcontinent cyclothems, Streptognathodus is widely accepted to have lived in normal open-marine, warm surface waters, whereas coeval Hindeodus is considered to have inhabited more shallow environments (Heckel, 1994, 2002; Boardman et al., 1998). Mesogondolella, Jinogondolella and Clarkina are usually found in offshore, predominately basinal facies (e.g. Wardlaw and Collinson, 1984; Mei and Henderson, 2001; Kozur and Wardlaw, 2010). Analyses of coeval Early Permian Mesogondolella and Streptognathodus elements from the Naqing section (this study) yield comparable δ18O values, suggesting both taxa could live in waters with comparable temperatures. Since Streptognathodus is considered as a surface-water dwelling taxon, a similar habitat is unexpected for co-occurring Mesogondolella.

75 Chapter 4 Permian composite oxygen isotope record

The comparison of the oxygen isotope ratios of Pennsylvanian Streptognathodus and Hindeodus indicated no significant differences in δ18O, but a 0.6‰ higher oxygen isotope value of Gondolella relative to Hindeodus (Joachimski et al., 2006). A similar difference in δ18O between Hindeodus and coeval Jinogondolella or Clarkina was observed in a study on conodonts from two Guadalupian–Lopingian boundary sections (Chen et al., 2011), suggesting a shallower and warmer water habitat for Hindeodus, and potentially a deeper, cooler water habitat for Jinogondolella and Clarkina. Interestingly, oxygen isotope analyses measured on co-occuring Hindeodus and Clarkina from the PTB section of Meishan yield comparable δ18O values in samples below the PT event horizon, but a 0.8‰ difference in latest Permian and earliest Triassic samples (Joachimski et al., 2012). The comparable oxygen isotope ratios of Clarkina and Hindeodus indicate that both taxa were able to thrive in comparable shallow water depths during low sea level (e.g. <50m; Joachimski et al., 2012). However, as sea level was rising in the latest Permian and Early Triassic, Clarkina migrated to deeper and slightly cooler waters. This interpretation is consistent with the early hypothesis that Clarkina was a free-swimming taxon thriving in offshore or even deeper waters (Tian, 1993a, b; Orchard, 1996; Lai et al., 2001). In summary, the oxygen isotope analyses of Permian conodonts suggests that Streptognathodus, Hindeodus, and probably Early Permian Mesogondollella as well, were thriving in near-surface sea water and recorded surface water temperature, whereas the habitats of Jinogondolella and Clarkina are more difficult to constrain. The isotope results suggest that both surface and deeper water masses could be potential life habitat for gondolellid conodonts.

4.4.2.2 Permian conodont apatite oxygen isotope record

In comparison to the brachiopod calcite δ18O record, the composite

76 Chapter 4 Permian composite oxygen isotope record conodont apatite δ18O record shows generally minor variations (1–2‰) for samples from specific time intervals. However, Middle Permian conodont samples from the Guadalupe Mountains have generally higher δ18O values (21 to 22‰) compared to conodont samples from South China (20 to 21‰; Fig. 4.3). This is consistent with observations made by Korte et al. (2005) showing higher δ18O values for brachiopods from the Guadalupe Mountains in comparison to brachiopods from other areas. These higher oxygen isotope ratios can be explained by higher evaporation rates and thus higher δ18O values of sea water in the Delaware basin. The evaporate component of the Artesia Group (Guadalupe shelf deposits) are well known (e.g. Kendall, 1969). Lopingian conodonts from Iran and South China also show a major difference in δ18O, with Changhsingian conodonts from South China having values ranging from 21 to 22.3‰, whereas coeval conodont samples from Iran are between 19.8 and 21‰ (Fig. 4.3). This difference can not be attributed to their different palaeolatitudes, since Iran was situated at subtropical latitudes (20–30°S), whereas South China was located in an equatorial position. Since a higher evaporation/precipitation ratio is expected in subtropical latitudes, sea water δ18O values are expected to have been higher in Iran. These examples confirm that, as indicated by the oxygen isotope ratios of brachiopod shell calcite, sea water in different areas of the Permian oceans was characterized by relatively large differences in δ18O as consequence of regional differences in the evaporation/precipitation ratio. In order to exclude this effect, the following discussion will focus on the oxygen isotope record measured only on conodonts from the South China block. Conodont apatite oxygen isotope values are generally between 22 and 23‰ throughout the Asselian, Sakmarian and Artinskian. No significant differences are observed in the oxygen isotope ratios of coeval Mesogondolella and Streptognathodus elements, suggesting that both taxa were thriving in

77 Chapter 4 Permian composite oxygen isotope record comparable water depths (Fig. 4.4). Since early Permian conodont samples were collected from a slope setting representing fully marine conditions, evaporation and thus higher salinities cannot explain the relatively high oxygen isotope ratios. This interpretation is supported by palaeobotanical evidence that indicates humid climatic conditions for south China during most of the Permian (Cleal and Thomas, 2005; Hilton and Cleal, 2007). Consequently, the high δ18O values are interpreted as expression of the LPIA, with the formation of isotopically light continental ice masses resulting in high δ18O values of Early Permian sea water. The LPIA was suggested to have had ice volumes comparable to the Pleistocene glacial maxima (Buggisch et al., 2008; Elrick 18 and Scott, 2010). Assuming a δ Ow of +1‰ VSMOW for Early Permian seawater, sea water temperatures calculated from the oxygen isotope ratios of Early Permian conodonts would have been between 26 and 30 °C. These temperature estimates seem realistic for tropical seas and are in agreement with the sedimentary record of the South China block, where the Early Permian is represented by carbonate sedimentation with abundant ‘warm-water’ taxa such as fusulinacean foraminifera, rugose corals and the occurrence of reefal buildups (Shi, 2001; Wang et al., 2006). Additionally, oncoid-bearing deposits are widely developed in Asselian to Sakmarian strata across the entire South China platform, indicating warm and shallow waters, and which are potentially comparable to the modern tropical environments on the Great Bahama Bank or the west Florida shelf (Shi and Chen, 2006). Conodont apatite δ18O values began to decrease in the early to middle Kungurian with minimum values registered in the late Capitanian. The gradual 2‰ decrease as recorded by gondolellids corresponds to 8 °C warming if explained exclusively by a change in surface water temperature. If gondolellids lived in deeper waters, the 2‰ change in δ18O would preferentially correspond to a change in the oxygen isotope composition of Permian sea water and

78 Chapter 4 Permian composite oxygen isotope record

Fig. 4.5 Comparison of Permian conodont δ18O record (circles–hindeodids, triangles –streptognathodids, squares–gondollelids) from South China with glacial intervals in Gondwana (WA: Western Australia, AN: Antarctic; EA: Eastern Australia) and Siberia (SI) after Fielding et al. (2008c), and the magnitude of eustatic sea level change during the Permian (modified after Rygel et al. 2008). Palaeotemperature calculated from δ18O of conodont apatite, assuming a δ18O value for Permian sea water of −1‰ VSMOW (ice-free–black scale) and +1‰ VSMOW (glacial–red scale). Absolute ages after Gradstein et al. (2004). Alternate ages (top) after Schmitz and Davydov (2012) and Shen et al. (2011). Eruption ages of Emeishan and Siberian Trap volcanism after Sun et al. (2010) and Mundil et al. (2004), respectively. translate into a 200 m sea-level rise as consequence of the waning of continental ice masses, assuming a change in δ18O of seawater of 0.1‰ per 10 m glacio-eustatic sea-level change (Schrag et al., 1996). For comparison,

79 Chapter 4 Permian composite oxygen isotope record

Zachos et al. (2001) compiled oxygen isotope values of benthic foraminifera and calculated a 2.3‰ change in the oxygen isotope composition of sea water as consequence of the Cenozoic growth of ice sheets. Alternatively, a combination of climatic warming and waning of high-latitude ice masses could account for the relatively large decrease in δ18O. The 2.0‰ decrease is comparable to the δ18O change measured on surface-dwelling foraminifers from the equatorial Pacific (1.76‰) and Atlantic (1.80‰) during the Last Glacial Maximum (LGM) to the Holocene (Ignacio Martinet et al., 1997). This would correspond to a lowering of sea level by 120 m (Fairbanks, 1989; Peltier and Fairbanks, 2006) or 130 m (Yokoyama et al., 2000; Mitrovica, 2003), and a cooling of low latitudes by 1–3 °C (Bard et al., 1997; Pelejero et al., 1999; Lea et al., 2002; Rosell-Melé et al., 2004; Barker et al., 2005; Hippler et al., 2006), or 4–6 °C (Beck et al., 1992; Guilderson et al., 1994; Stute et al., 1995; Thompson et al., 1995; McCulloch et al., 1999). This interpretation suggests that the main deglaciation of the LPIA would have occurred in the Kungurian to Capitanian. This is in contrast to the widely accepted interpretation that, based on the abundance and distribution of glacial deposits, the LPIA reached its climax during the Asselian and early Sakmarian (Isbell et al., 2003; Fielding et al., 2008c; Zeng et al., 2012), and also that high latitude ice sheets melted during the later Sakmarian. Other evidence for post-Sakmarian glaciers were reported from eastern and Western Australia (Fielding et al., 2008c) and Siberia (Chumakov, 1994) until the late Capitanian (Fig. 4.5). Interestingly, the glacial record from eastern Australia suggests that glaciation P2 (late Sakmarian to mid-Artinskian) was even larger than the preceding glacial phase P1 (Asselian to early Sakmarian), a glacial interval generally thought to represent the climax of the LPIA (Fielding et al., 2008c). Furthermore, non-glacial sediments between glacials P2 and P3 indicate repeated and large-scale relative sea-level changes up to 70 m, suggesting that

80 Chapter 4 Permian composite oxygen isotope record ice sheets were waning and waxing during this time period (Fielding et al., 2008a). For example, small and transient Artinskian to Kungurian glacial deposits were reported from Western Australia (Mory et al., 2008). The succeeding glacial intervals P3 (late Kungurian to latest Roadian) and P4 (late Wordian to late Capitanian) reported from eastern Australia are considered to represent only regionally alpine-style glaciations containing limited ice volume (Fielding et al., 2008a, b; Rygel et al., 2008; Birgenheier et al., 2010), which would have only a minor influence on the oxygen isotopic composition of Permian seawater. Interestingly, a late Kungurian termination of LPIA as indicated by the conodont δ18O data agrees well with a significant decrease in the magnitude of glacioeustatic sea level fluctuations as suggested by Ross and Ross (1987; Fig. 4.5). In contrast, a recently published compilation of glacio-eustatic sea-level by Rygel et al. (2008) suggests sea level fluctuations of up to 120 m during the late Pennsylvanian to middle Sakmarian, and significantly minor fluctuations of 30–70 m and 10–60 m during the Middle Sakmarian and the Roadian through Capitanian, respectively (Fig. 4.5). However, the precise correlation between the conodont δ18O record from South China and the glacial record in Gondwana remains difficult due the lack of a well-constrained biostratigraphic correlations as well as the potential stratigraphic incompleteness of the Gondwanan successions as consequence of glacial erosion. The conodont oxygen isotope record across the Guadalupian–Lopingian (G–L) transition suggests warming of about 4 °C in the late Capitanian, followed by cooling of about 6 to 8 °C across the G–L boundary and in the earliest Wuchiapingian, and again significant warming in the early Wuchiapingian. The rapid 4 °C warming and succeeding 6 to 8 °C cooling coincide with the main and final phase of Emeishan volcanism, respectively, and with relative sea level changes suggesting that the variations in the oxygen

81 Chapter 4 Permian composite oxygen isotope record isotope ratios may be the combined effect of changes in water depth and volcanism-related climate changes (Chen et al., 2011). Starting from the early Wuchiapingian and lasting until the late Changhsingian, conodont δ18O values increase by 2‰. This increase in δ18O corresponds to cooling of 9 °C, with absolute temperatures decreasing from 32 to 23 °C if a δ18O value of -1‰ VSMOW (ice-free world) is assumed for Late Permian sea water. Cool climatic conditions in the Changhsingian were previously proposed by Beauchamp and Baud (2002), Kozur (1998), Mei et al. (2002) and Mertmann (2003), based on faunal diversity, conodont distribution, and lithological evidence as well as oxygen isotope ratios of carbonates (for summary see Table 3 in Racki and Wignall, 2005). In addition, cooling is indicated by the occurrence of distinct peri-Gondwanan cool water brachiopod elements and the absence of compound corals in Changhsingian strata in southwestern Tibet (Shen et al., 2006, 2010). However, 9 °C cooling in low-latitudes seems unrealistic during non-glacial times, taking into account that the tropics were only 1 to 6 °C cooler during the Last Glacial maximum in comparison to the Holocene. Interestingly, inferred cooling as indicated by the conodont apatite δ18O record coincides with the Wuchiapingian to Changhsingian transgression observed in the Tethys area and the South China block (Chen et al., 1998). As a consequence, Clarkina (or Neogondolella) may have moved to deeper and thus cooler waters. Taking a modern water temperature/water depth gradient for low-latitudinal seas, 9 °C cooling would correspond to deepening of approximately 130 to 150 m (Levitus and Boyer, 1994). Since this estimate is inconsistent with the minor sea level rise in the Wuchiapingian and Changhsingian, the Lopingian increase in the oxygen isotope ratios can best be explained as the combined effect of climate cooling and Clarkina moving to deeper/cooler waters following the sea level rise. The interpretation that Clarkina was able to live in different water depths

82 Chapter 4 Permian composite oxygen isotope record is confirmed by the oxygen isotope ratios of Clarkina and Hindeodus from the PTB interval (Joachimski et al., 2012). Latest Permian Hindeodus and Clarkina have comparable oxygen isotope ratios, indicating that both taxa lived in identical water depths. Both taxa show decreasing δ18O values across the Permian–Triassic extinction horizon, translating into significant climate warming. However, the decrease in gondollelid δ18O is smaller in comparison to the increase measured on Hindeodus, which suggests that Clarkina was moving to deeper waters in conjunction with a sea-level rise in the latest Changshingian (Joachimski et al., 2012). Climate warming in the latest Permian coincides with both the significant negative carbon isotope excursion and the eruption of the Siberian Traps, and may be interdependent and explained by the emission of volcanic CO2 (Renne et al., 1995; Sobolev et al., 2011), thermal decomposition of pre-existing sediments induced by the volcanism (Svensen et al., 2009; Shen et al., 2011), and the destabilization of 12C-enriched methane hydrates contained in permafrost soils (e.g. Retallack and Jahren, 2008). All of these processes will have resulted in higher greenhouse gas levels, and may have contributed to the significant climate warming observed in the latest Permian.

4.5 Conclusions

The compilation of brachiopod calcite as well as conodont apatite oxygen isotope data shows relatively large differences in δ18O for datasets derived from different palaeocontinents. These differences are attributed to local hydrological conditions associated with regional aridity and differences in the evaporation/precipitation ratio. As a consequence, the reconstruction of the oxygen isotope composition of Permian sea water and/or of sea water temperature has to be based on samples from a specific area. The conodont apatite oxygen isotope record from the South China block is

83 Chapter 4 Permian composite oxygen isotope record used to interpret the palaeoclimate history of the Permian. The conodont taxa Streptognathodus and Hindeodus are interpreted to represent surface water dwellers recording surface water δ18O and temperature. Gondollelids, which represent the most abundant taxa in the sections studied, may have lived in different parts of the water column, and may be used to reconstruct secular changes in the oxygen isotope composition of Permian sea water and to a lesser degree to reconstruct palaeotemperature variations. This interpretation is based on the observation that the oxygen isotope ratios of gondollelids seem to co-vary with sea-level changes. During a high or rising sea level, Permian gondollelids show higher δ18O values in comparison to hindeodids. In contrast, during lower sea level and thus minimum absolute water depths, gondollelids and hindeodids show comparable isotope ratios suggesting that both taxa were living in comparable water depths. The conodont apatite oxygen isotope record supports glaciated conditions in the Early Permian by showing high values between 22 and 23‰, but argues for a later deglaciation phase than indicated by sedimentological evidence. The δ18O decrease of 2‰ during the Kungurian to Capitanian was measured on gondollelid conodonts and is interpreted to reflect predominately the waning of high-latitude ice masses and subordinately climate warming. This suggests that post-Sakmarian ice sheets may have been widespread in Gondwana until the Kungurian. The Guadalupian–Lopingian transition interval is characterized by significant temperature fluctuations with 4 °C warming succeeding by 6 to 8 °C cooling, interpreted as combined climate change initiated by Emeishan volcanism and changes in habitat depth of gondollelid conodonts as consequence of sea-level changes. The significant increase in δ18O in the Wuchiapingian to Changhsingian to values around 22.0‰ is ascribed to the combined effect of climate cooling and Clarkina migrating to deeper/cooler waters as sea level was rising. The rapid δ18O decrease of 2 ‰ across the PTB

84 Chapter 4 Permian composite oxygen isotope record indicates a dramatic global warming in the latest Permian and earliest Triassic, potentially caused by massive injection of greenhouse gases to the ocean-atmosphere system as a consequence of Siberian Trap volcanism, volcanism-induced thermal decomposition of organic carbon-rich sediments and/or destabilization of methane hydrates contained in permafrost soils.

85

86 Chapter 5 Conclusions

Chapter 5: Conclusions

Separate oxygen isotope analyses to different Permian conodont taxa confirm the previous interpretation that Streptognathodus and Hindeodus were surface water dwellers, whereas the habitat of gondollelids is more difficult to constrain. Based on the data presented in this study, it is suggested that gondollelids were able to live in different water depths depending on changes in sea level and water depth. The conodont apatite oxygen isotope record from China supports glaciated conditions in the Early Permian, but argues for a later deglaciation phase than indicated by sedimentological evidence (Kungurian vs. Sakmarian). The oxygen isotope record across the Guadalupian–Lopingian transitional interval measured on gondolellid conodonts indicates a 4 °C warming succeeding by 6 to 8 °C cooling. These temperature fluctuations coincide with the main and final phase of Emeishan volcanism. However, parallel changes in sea-level suggest that changes in water depth and thus water temperature in combination with superimposed volcanism-induced climatic changes may have been responsible for the observed temperature changes. In addition, a 1.5‰ increase in inorganic δ13C was found to occur immediately after the warming episode, indicating that increased temperature may have resulted in an intensified hydrological cycle, fertilization of the oceans and enhanced primary productivity in the latest Capitanian. High oxygen isotope values around 22.0‰ (VSMOW) are reported for the Changhsingian, supporting previous suggestions based on faunal diversity, conodont distribution, and lithological evidence that cool climatic conditions may have prevailed during the Changhsingian. Warming of 4 to maximum 8 °C as indicated by a pronounced decrease in δ18O of latest Permian conodonts

87 Chapter 5 Conclusions seems to coincide with the eruption of Siberian Trap volcanism. In addition, the change in the oxygen isotope ratios in the latest Permian is coincident with the significant negative carbon isotope shift as well as the major phase of the end Permian mass extinction, thus in favor of the interpretation that the addition of isotopically light carbon to the ocean-atmosphere system by Siberian trap volcanism and related processes resulted in higher greenhouse gas levels and global warming, with climate warming contributing to the collapse of both marine and terrestrial ecosystems during the end-Permian mass extinction. Reconstructed very warm climate conditions in the Early Triassic may have played an important role in the timing and pattern of the slow recovery of marine and terrestrial ecosystems in the aftermath of the Permian–Triassic crisis. Oxygen isotope analyses of conodonts from Iran and the USA as well as the comparison with oxygen isotope records based on the analysis of brachiopod calcite document that, in case of the Permian, oxygen isotopes records from different palaeocontinents are characterized by significantly different values, probably as consequence of the local hydrological conditions (evaporation/precipitation). In order to confirm the global significance of the oxygen isotope trends reported from China, additional oxygen isotope records from other regions are needed.

88 References

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116 Tables

Tables

Table 2.1 Oxygen isotope values of conodonts from Penglaitan and Tieqiao sections. Palaeotemperatures calculated assuming an oxygen isotopic composition of -1‰ VSMOW for Permian sea water

δ18O Temp. Sample Bed no. Conodont taxa analysed apatite Conodont Zone (‰ VSMOW) (°C)

Penglaitan section

PLTS 19.7 16 Clarkina.sp. 21.5 23.8 C. asymmetrica PLTS 14.05 16 Clarkina.sp. 21.8 22.5 C. asymmetrica PLTS 11.4 17 Clarkina.sp. 21.1 25.4 C. dukouensis PLTS 10.36 17 Clarkina.sp. 20.8 26.6 C. dukouensis PLTS 9.67 17 Clarkina.sp. 20.9 26.2 C. dukouensis PLTS 6.59 17 Clarkina.sp. 21.0 25.9 C. dukouensis PLTS 6.19 17 Clarkina.sp. 20.5 28.0 C. dukouensis PLTS 5.8 17 Clarkina.sp. 19.9 30.5 C. dukouensis PLTS 4.43 17 Clarkina.sp. 21.1 25.4 C. dukouensis PLTS 1.08 17 Clarkina.sp. 20.6 27.6 C. postbitteri postbitteri PL1 18/6k C. postbitteri postbitteri 20.7 27.0 C. postbitteri postbitteri PL7 18/4k ramiform elements 19.9 30.3 J. granti PL7 18/4k J. granti, J. xuanhanensis 20.3 29.0 J. granti PL8 18/4i J. granti, J. xuanhanensis 19.8 31.0 J. granti PL9 18/4f J. granti 20.1 29.5 J. granti PL10 18/4b J. granti, J. xuanhanensis 20.3 28.8 J. granti PL11 18/3c J. granti, 20.1 29.7 J. granti PL13 18/3a Hindeodus sp., ramiform elements 19.1 33.8 J. granti PL14 18/2c Hindeodus sp., ramiform elements 19.0 34.4 J. xuanhanensis PL15 18/2a J. xuanhanensis 20.3 28.9 J. xuanhanensis PL18 19 ramiform elements 20.3 28.7 J. xuanhanensis PL18 19 J. xuanhanensis, J. altudensis 20.5 28.1 J. xuanhanensis J. xuanhanensis, J. prexuanhanensis, PL19 19 20.3 28.6 J. xuanhanensis ramiform elements

Jindogondolella sp., ramiform PL26 20 20.9 26.2 J. shannoni elements

Tieqiao section TQ 63 H 129 Clarkina.sp. 18.9 34.9 C. liangshanensis TQ 62 H 129 Clarkina.sp. 18.6 36.0 C. liangshanensis TQ 61 H 128 Clarkina.sp. 19.3 33.0 C. liangshanensis TQ 60 H 126 Clarkina.sp. 19.1 33.9 C. liangshanensis TQ 58 H 122 Clarkina.sp. 18.7 35.7 C. leveni TQ 57 H 122 C. asymmetrica, C. doukouensis 19.1 34.0 C. leveni

117 Tables

Continued

δ18O Temp. Sample Bed no. Conodont taxa analysed apatite Conodont Zone (‰ VSMOW) (°C) TQ 56 H 122 C. asymmetrica, C. doukouensis 19.7 31.2 C. leveni TQ 54 H 120 C. postbitteri., C. asymmetrica 19.9 30.3 C. dukouensi TQ 1 H 119/8 C. postbitteri postbitteri 20.9 26.3 C. postbitteri postbitteri TQ 2 H C. postbitteri postbitteri 20.6 27.5 C. postbitteri postbitteri TQ 3 H 119/7e C. postbitteri hongshuiensis 20.1 29.7 C. postbitteri postbitteri TQ 5 H 119/6j C. postbitteri hongshuiensis 20.3 29.0 C. postbitteri,hongshuiensis TQ 7 H 119/6g J. granti 19.7 31.4 J. granti TQ 8 H 119/6f J. granti 20.2 29.4 J. granti TQ 8 H 119/6f ramiform elements 20.1 29.7 J. granti TQ 9 H 119/6d J. granti 20.0 29.9 J. granti TQ 10 H 119/6c J. granti 20.3 28.6 J. granti TQ 11 H 119/6b J. granti 19.6 31.6 J. granti TQ 12 H 119/6a J. granti 20.0 30.2 J. granti TQ 13 H 119/5 J. granti 19.9 30.4 J. granti TQ 14 H 119/5 Hindeodus sp., ramiform elements 19.1 33.8 J. granti J. xuanhanensis, Hindeodus sp., TQ 16 H 119/2 19.9 30.7 J. xuanhanensis ramiform elements J. xuanhanensis, J. TQ 17 H 118 20.1 29.8 J. xuanhanensis prexuanhanensis TQ 18 H 118 J. xuanhanensis 20.2 29.1 J. xuanhanensis TQ 31 H 116 J. postserrata, J. shannoni? 20.9 26.4 J. postserrata TQ 33 H 116 J. postserrata, Hindeodus sp. 20.9 26.4 J. postserrata J. postserrata, J. shannoni, TQ 35 H 116 20.6 27.6 J. postserrata ramiform elements J. postserrata, Hindeodus sp., TQ 37 H 116 20.8 26.7 J. postserrata ramiform elements J. postserrata, Hindeodus sp., TQ 38 H 116 20.7 27.1 J. postserrata ramiform elements J. postserrata, Hindeodus sp., TQ 40 H 116 20.3 28.7 J. postserrata ramiform elements J. postserrata, Hindeodus sp., TQ 41 H 116 20.8 26.8 J. postserrata ramiform elements TQ 42 H 116 J. postserrata, ramiform elements 20.5 28.0 J. postserrata TQ 42 H 116 J. postserrata 20.5 28.0 J. postserrata TQ 43 H 116 J. postserrata 20.7 27.0 J. postserrata TQ 43 H 116 J. postserrata, ramiform elements 20.5 28.2 J. postserrata TQ 44 H 116 J. postserrata, ramiform elements 20.8 26.6 J. postserrata J. aserrata, Hindeodus sp., TQ 45 H 115 20.2 29.4 J. aserrata ramiform elements TQ 46 H 115 J. aserrata, ramiform elements 20.9 26.4 J. aserrata J. aserrata, Hindeodus sp., TQ 47 H 115 20.6 27.5 J. aserrata ramiform elements

118 Tables

Table 3.1 Oxygen isotope ratios of conodonts from the Meishan and Shangsi sections (South China). Paleotemperatures are calculated using the equation given by Pucéat et al. (2010) assuming a δ18O value for Late Permian sea water of -1‰ VSMOW

Sample Bed no. δ18O (‰VSMOW) T (°C) Clarkina Hindeodus ramiforms

Meishan B & C sections

MSC 33 MSC 29 18.8 35.1 MSC 32 MSC 28 19.3 32.9 MSC 31 MSC 27d 20.4 19.4 28.5 / 32.8 MSC 30 MSC 27c 20.4 19.4 28.3 / 32.5 MSC 29 MSC 27b 19.5 32.1 MSC 28 MSC 27a 20 19.4 29.9 / 32.6 MSC 27 MSC 26 21 25.9 MSC 26 MSC 24f-1 20.9 26.3 MSB 25 MSB 24f 21.3 24.6 MSC 23 MSC 24e Top 21.3 24.6 MSC 24 MSC 24e 21.6 21.2 23.2 / 25.2 MSC 22 MSC 24e 21 25.9 MSC 21 MSC 24d 21.1 21.2 20.8 25.3 / 25.0 / 26.8 MSC 20 MSC 24d-4 21.4 24.2 MSC 19 MSC 24d-3 21.4 24.2 MSC 18 MSC 24d-2 21.6 23.3 MSC 17 MSC 24d-1 21 25.9 MSC 16 MSC 24c 21.2 25 MSC 15 MSC 24c-2 21.2 25.9 MSC 14 MSC 24c-1 21 25 MSC 13 MSC 24c 20.7 27.1 MSC 12 MSC 24b 20.7 20.6 27.0 / 27.6 MSC 11 MSC 24a 21.2 21.2 21 25.0 / 25.0 / 26.1 MSC 10 MSC 24a 21.3 24.6 MSC 9 MSC 24a base 21.5 23.8 MSC 8 MSC 23-7 22 21.6 MSC 7 MSC 23-4 22 21.6 MSB 6 MSB 23-3 21.9 22.1 MSB 5 MSB 23-1 22.3 20.4 MSC 4a MSC 23 22.3 MSB 4 MSB 23-1 base 21.3 24.6 MSB 3 MSB 22 top 21.8 22.5 MSC 2 MSC 22-15 21.6 23.3 MSC 1a MSC 22c 21.3 MSC 1 MSC 22-14 21.6 23.3

119 Tables

Continued

Sample Bed no. δ18O (‰VSMOW) T (°C)

Clarkina Hindeodus ramiforms

Shangsi section

33 18.7 35.5 32 19.3 33 31b 19 34.3 31a 19 34.3 30d 18.9 34.8 28c 19.5 32 27c 20 30.2 27a 21.4 24.1

26 21.6 23.3

120 Tables

Table 4.1 Stratigraphic age, analysed conodont taxa and oxygen isotope ratios of conodont apatite

Sample Age (Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

SS33 Shangsi section Sichuan China 250,1 Induan Isarcicella isarcica Hindeodus sp., Ramiform 18.7 SS32 Shangsi section Sichuan China 250,3 Induan Isarcicella isarcica Hindeodus sp., Ramiform 19.3 SS31b Shangsi section Sichuan China 250,4 Induan Isarcicella isarcica Hindeodus sp., Ramiform 19.0 SS31a Shangsi section Sichuan China 250,7 Induan Isarcicella lobata Hindeodus sp., Ramiform 19.0 Hindeodus sp., Ramiform SS30d Shangsi section Sichuan China 250,7 Induan Isarcicella lobata 18.9 Hindeodus sp. P1 MSD28 Meishan D section ZheJiang China 251,0 Induan Isarcicella isarcica Hindeodus sp., Ramiform 19.3 MSD29 Meishan D section ZheJiang China 251,0 Induan Isarcicella isarcica Hindeodus sp., Ramiform 18.8 Hindeodus changxingensis MSD27b Meishan D section ZheJiang China 251.0 Changhsingian Hindeodus sp., Ramiform 19.5 Clarkina meishanensis MSD27c Meishan D section ZheJiang China 251,0 Induan Hindeodus parvus Clarkina sp. P1 20.4 MSD27c Meishan D section ZheJiang China 251,0 Induan Hindeodus parvus Hindeodus sp., Ramiform 19.4 MSD27d Meishan D section ZheJiang China 251,0 Induan Hindeodus parvus Clarkina sp. P1 20.4 MSD27d Meishan D section ZheJiang China 251,0 Induan Hindeodus parvus Hindeodus sp., Ramiform 19.4 Hindeodus changxingensis MSD26 Meishan D section ZheJiang China 251,0 Changhsingian Clarkina sp. P1 21.0 Clarkina meishanensis Hindeodus changxingensis MSD27a Meishan D section ZheJiang China 251,0 Changhsingian Clarkina sp. P1 20.0 Clarkina meishanensis

Hindeodus changxingensis MSD27a Meishan D section ZheJiang China 251,0 Changhsingian Hindeodus sp., Ramiform 19.4 Clarkina meishanensis

MSC24e top Meishan C section ZheJiang China 251.0 Changhsingian C. meishanensis Clarkina sp. 21. 3

121 Tables 122

Sample Age (Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

MSD24e Meishan D section ZheJiang China 251.0 Changhsingian C.changxingensis yini Clarkina sp. P1 21.6 MSD24e Meishan D section ZheJiang China 251.0 Changhsingian C.changxingensis yini Hindeodus sp., Ramiform 21.2 MSC24F-1 Meishan C section ZheJiang China 251.0 Changhsingian C. meishanensis Clarkina sp. 20.9 C. changxingensis yini/C. MSC24e Meishan C section ZheJiang China 251.0 Changhsingian Clarkina sp. 21.0 meishanensis zhangi MSD24d Meishan D section ZheJiang China 251.0 Changhsingian C. changxingensis yini Clarkina sp. P1 21.1 MSD24d Meishan D section ZheJiang China 251.0 Changhsingian C. changxingensis yini Hindeodus sp. P1 21.2 MSD24d Meishan D section ZheJiang China 251.0 Changhsingian C. changxingensis yini Hindeodus sp., Ramiform 20.8 C. changxingensis yini/ MSC24d2 Meishan C section ZheJiang China 251.1 Changhsingian Clarkina sp. 21.6 C. meishanensis zhangi MSD24c Meishan D section ZheJiang China 251.1 Changhsingian C. changxingensis yini Hindeodus sp., Ramiform 21.2 C. changxingensis yini/C. MSC24d3 Meishan C section ZheJiang China 251.1 Changhsingian Clarkina sp. 21.4 meishanensis zhangi C. changxingensis yini/C. MSC24d-4 Meishan C section ZheJiang China 251.1 Changhsingian Clarkina sp. 21.4 meishanensis zhangi MSD24a Meishan D section ZheJiang China 251.1 Changhsingian C. changxingensis yini Clarkina sp. P1 21.2 MSD24a Meishan D section ZheJiang China 251.1 Changhsingian C. changxingensis yini Hindeodus sp. P1 21.2 C. changxingensis yini/C. MSC24C-2 Meishan C section ZheJiang China 251.1 Changhsingian Clarkina sp. 21.2 meishanensis zhangi C. changxingensis yini/C. MSC24d1 Meishan C section ZheJiang China 251.1 Changhsingian Clarkina sp. 21.0 meishanensis zhangi C. changxingensis yini/C. MSC24d-1 Meishan C section ZheJiang China 251.1 Changhsingian Clarkina sp. 20.4 meishanensis zhangi

MSD24a Meishan D section ZheJiang China 251.1 Changhsingian C. changxingensis yini Hindeodus sp., Ramiform 21.0

Tables

Sample Age (Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

MSD24b Meishan D section ZheJiang China 251.1 Changhsingian C. changxingensis yini Clarkina sp. P1 20.7 MSD24b Meishan D section ZheJiang China 251.1 Changhsingian C. changxingensis yini Hindeodus sp., Ramiform 20.6 C. changxingensis yini/C. MSC24C Meishan C section ZheJiang China 251.1 Changhsingian Clarkina sp. 20.7 meishanensis C. changxingensis yini/C. MSC24C-1 Meishan C section ZheJiang China 251.1 Changhsingian Clarkina sp. 21.0 meishanensis zhangi

MSC24a C. changxingensis yini/C. Meishan C section ZheJiang China 251.1 Changhsingian Clarkina sp. 21.3 925-29(37) meishanensis zhangi

MSC24a C. changxingensis yini/C. Meishan C section ZheJiang China 251.1 Changhsingian Clarkina sp. 21.5 base meishanensis zhangi

C. changxingensis yini/C. MSC23-7 Meishan C section ZheJiang China 251.1 Changhsingian Clarkina sp. 22.0 meishanensis zhangi

C. changxingensis yini/C. MSC23-4 Meishan C section ZheJiang China 251.2 Changhsingian Clarkina sp. 22.0 meishanensis zhangi

C. changxingensis yini/C. MSC23-3 Meishan C section ZheJiang China 251.2 Changhsingian Clarkina sp. 21.1 meishanensis zhangi C. changxingensis yini/C. MSC23-3 Meishan C section ZheJiang China 251.2 Changhsingian Clarkina sp. 21.8 meishanensis zhangi Ab 1.7 Abadeh section Iran 251.2 Changhsingian Clarkina sp. 20.0 C. bachmanni MSC23-1 Meishan C section ZheJiang China 251.2 Changhsingian Clarkina sp. 22.3 /C.parasubcarinata MSC22-14 Meishan C section ZheJiang China 251.2 Changhsingian C. bachmanni Clarkina sp. 21.6 MSC22-15 Meishan C section ZheJiang China 251.2 Changhsingian C. bachmanni Clarkina sp. 21.6 MSC 22-12 Meishan C section ZheJiang China 251.2 Changhsingian C. bachmanni Clarkina sp. 22.3

SS28c Shangsi section Sichuan China 251.2 Changhsingian Hindeodus changxingensis Hindeodus sp., Ramiform 19.5

123 Tables 124

Sample Age (Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

Ab 2.5 Abadeh section Iran 251.3 Changhsingian Clarkina sp. 20.3 JY 5 Jiangya section Hunan China 251.3 Changhsingian Clarkina sp. 21.2 Ab 2.8 Abadeh section Iran 251.3 Changhsingian Clarkina sp. 20.1 MSC21-9 Meishan C section ZheJiang China 251.4 Changhsingian C. changxingensis Clarkina sp. 21.9 SS27c Shangsi section Sichuan China 251.4 Changhsingian C. meishanensis Clarkina sp. 20.0 SS27a Shangsi section Sichuan China 251.4 Changhsingian C. meishanensis Clarkina sp. 21.4 SS26 Shangsi section Sichuan China 251.4 Changhsingian C. yini Clarkina sp. 21.6 MSC 21-2b Meishan C section ZheJiang China 251.4 Changhsingian C. changxingensis Clarkina sp. 21.9 Ab 3.7 Abadeh section Iran 251.5 Changhsingian Clarkina sp. 20.1 Ab 3.9 Abadeh section Iran 251.5 Changhsingian Clarkina sp. 20.5 MSc 20-8 Meishan C section ZheJiang China 251.5 Changhsingian C. changxingensis Clarkina sp. 21.2 Ab 4.25 Abadeh section Iran 251.6 Changhsingian Clarkina sp. 20.9 Ab 4.34 Abadeh section Iran 251.6 Changhsingian Clarkina sp. 20.9 Ab 4.5 Abadeh section Iran 251.6 Changhsingian Clarkina sp. 20.1 Ab 4.75 Abadeh section Iran 251.6 Changhsingian Clarkina sp. 20.5 MSC 19-20 Meishan C section ZheJiang China 251.6 Changhsingian C. changxingensis Clarkina sp. 21.6 Ab 5.0 Abadeh section Iran 251.7 Changhsingian Clarkina sp. 20.6 Ab 5.75 Abadeh section Iran 251.8 Changhsingian Clarkina sp. 20.4 MSc 16-11 Meishan C section ZheJiang China 251.9 Changhsingian C. changxingensis Clarkina sp. 21.9 MSC 16-9 Meishan C section ZheJiang China 251.9 Changhsingian C. changxingensis Clarkina sp. 21.8 Ab 6.3 Abadeh section Iran 251.9 Changhsingian Clarkina sp. 20.3 Ab 6.7 Abadeh section Iran 251.9 Changhsingian Clarkina sp. 19.9 MSC 16-3 Meishan C section ZheJiang China 252.0 Changhsingian C. changxingensis Clarkina sp. 22.3

Tables

Sample Age(Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

Ab 7.2 Abadeh section Iran 252.0 Changhsingian Clarkina sp. 20.1 Ab 8.23 Abadeh section Iran 252.2 Changhsingian Clarkina sp. 20.4 Ab 8.7 Abadeh section Iran 252.2 Changhsingian Clarkina sp. 20.2 Ab 9.4 Abadeh section Iran 252.3 Changhsingian Clarkina sp. 20.0 MSD 13-24 Meishan D section ZheJiang China 252.4 Changhsingian C. changxingensis Clarkina sp. 20.7 Ab 10 Abadeh section Iran 252.4 Changhsingian Clarkina sp. 20.3 MSD 13-18 Meishan D section ZheJiang China 252.5 Changhsingian C. changxingensis Clarkina sp. 21.6 Ab 10.4 Abadeh section Iran 252.5 Changhsingian Clarkina sp. 20.6 MSD 13-13a Meishan D section ZheJiang China 252.5 Changhsingian C. changxingensis Clarkina sp. 21.4 MSD 13-9 Meishan D section ZheJiang China 252.6 Changhsingian C. changxingensis Clarkina sp. 21.3 Ab 11.2 Abadeh section Iran 252.6 Changhsingian Clarkina sp. 20.1 MSD 13-6b-6c Meishan D section ZheJiang China 252.7 Changhsingian C. changxingensis Clarkina sp. 22.0 Ab 11.95 Abadeh section Iran 252.7 Changhsingian Clarkina sp. 19.9 MSD 12-9d-9e Meishan D section ZheJiang China 252.8 Changhsingian C. changxingensis Clarkina sp. 21.8 MSD 12-4 Meishan D section ZheJiang China 252.8 Changhsingian C. subcarinata Clarkina sp. 21.3 Ab 12.45 Abadeh section Iran 252.8 Changhsingian Clarkina sp. 19.8 MSD 11-7 Meishan D section ZheJiang China 253.0 Changhsingian C. subcarinata Clarkina sp. 21.5 MSD 11-2b Meishan D section ZheJiang China 253.1 Changhsingian C. subcarinata Clarkina sp. 21.3 Ab 14.9 Abadeh section Iran 253.2 Changhsingian Clarkina sp. 19.9 MSD 9-14 Meishan D section ZheJiang China 253.2 Changhsingian C. wangi Clarkina sp. 20.9 Ab 15.3 Abadeh section Iran 253.2 Changhsingian Clarkina sp. 19.8 MSD 9-10 Meishan D section ZheJiang China 253.3 Changhsingian C. wangi Clarkina sp. 22.1 MSD 9-6 Meishan D section ZheJiang China 253.3 Changhsingian C. wangi Clarkina sp. 21.5

125 Tables 126

Sample Age(Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

Ab 16 Abadeh section Iran 253.3 Changhsingian Clarkina sp. 20.3 MSD 8-4 Meishan D section ZheJiang China 253.4 Changhsingian C. wangi Clarkina sp. 22.0 MSD 6-10 Meishan D section ZheJiang China 253.4 Changhsingian C. wangi Clarkina sp. 21.3 MSD 6-6 Meishan D section ZheJiang China 253.5 Changhsingian C. wangi Clarkina sp. 22.0 MSD 6-3 Meishan D section ZheJiang China 253.5 Changhsingian C. wangi Clarkina sp. 21.4 MSD 6-2 Meishan D section ZheJiang China 253.5 Changhsingian C. wangi Clarkina sp. 20.9 Ab 17 Abadeh section Iran 253.5 Changhsingian Clarkina sp. 20.9 MSD 5-14 Meishan D section ZheJiang China 253.5 Changhsingian C. wangi Clarkina sp. 21.2 MSD 5-12 Meishan D section ZheJiang China 253.5 Changhsingian C. wangi Clarkina sp. 20.7 MSD 5-11 Meishan D section ZheJiang China 253.5 Changhsingian C. wangi Clarkina sp. 20.9 MSD 5-10 Meishan D section ZheJiang China 253.5 Changhsingian C. wangi Clarkina sp. 19.7 MSD 5-7 Meishan D section ZheJiang China 253.6 Changhsingian C. wangi Clarkina sp. 21.1 MSD 5-8 Meishan D section ZheJiang China 253.6 Changhsingian C. wangi Clarkina sp. 21.2 MSD 5-2 Meishan D section ZheJiang China 253.6 Changhsingian C. wangi Clarkina sp. 21.6 MSD 5-1 Meishan D section ZheJiang China 253.6 Changhsingian C. wangi Clarkina sp. 20.9 Ab 18 Abadeh section Iran 253.6 Changhsingian Clarkina sp. 20.4 MSD 4b-11 Meishan D section ZheJiang China 253.7 Changhsingian C. wangi Clarkina sp. 20.6 MSD 4b-9 Meishan D section ZheJiang China 253.7 Changhsingian C. wangi Clarkina sp. 21.4 MSD 4b-4 Meishan D section ZheJiang China 253.7 Changhsingian C. wangi Clarkina sp. 20.9 MSD 4b-6 Meishan D section ZheJiang China 253.7 Changhsingian C. wangi Clarkina sp. 21.0 MSD 4a-5 Meishan D section ZheJiang China 253.8 Changhsingian C. wangi Clarkina sp. 20.6 Ab 19 Abadeh section Iran 253.8 Changhsingian Clarkina sp. 19.8 MSD 3-2 Meishan D section ZheJiang China 253.8 Wuchiapingian C. orientails Clarkina sp. 20.1 Ab 21.2 Abadeh section Iran 254.3 Wuchiapingian Clarkina sp. 21.6

Tables

Sample Age(Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

Ab 21.9 Abadeh section Iran 254.5 Wuchiapingian Clarkina sp. 20.3 JY 8 Jiangya section Hunan China 254.5 Wuchiapingian Clarkina sp. 21.6 JY 9 Jiangya section Hunan China 254.5 Wuchiapingian Clarkina sp. 21.9 JY 10 Jiangya section Hunan China 254.6 Wuchiapingian Clarkina sp. 22.0 JY 11-R Jiangya section Hunan China 254.7 Wuchiapingian Ramiform 20.4 JY 11 Jiangya section Hunan China 254.7 Wuchiapingian Clarkina sp. 21.1 JY 12 Jiangya section Hunan China 254.9 Wuchiapingian Clarkina sp. 20.5 GYD 192.7 Dukou section Sichuan China 255.3 Wuchiapingian Clarkina sp. 20.2 GYD 189.7 Dukou section Sichuan China 255.4 Wuchiapingian Clarkina sp. 19.5 Ab 27 Abadeh section Iran 255.6 Wuchiapingian Clarkina sp. 20.8 GYD 181.0 Dukou section Sichuan China 255.6 Wuchiapingian Clarkina sp. 20.6 GYD 177.6 Dukou section Sichuan China 255.7 Wuchiapingian Clarkina sp. 20.8 GYD 172.45 Dukou section Sichuan China 255.8 Wuchiapingian Clarkina sp. 21.1 Ab 28 Abadeh section Iran 255.8 Wuchiapingian Clarkina sp. 20.3 JY 16 Jiangya section Hunan China 255.9 Wuchiapingian Clarkina sp. 20.4 GYD 166.8 Dukou section Sichuan China 256.0 Wuchiapingian Clarkina sp. 20.8 GYD 162.8 Dukou section Sichuan China 256.1 Wuchiapingian Clarkina sp. 20.9 JY 19 Jiangya section Hunan China 256.2 Wuchiapingian Clarkina sp. 20.0 GYD 155.6 Dukou section Sichuan China 256.3 Wuchiapingian Clarkina sp. 21.1 Ab 30 Abadeh section Iran 256.3 Wuchiapingian Clarkina sp. 19.8 GYD 153.5 Dukou section Sichuan China 256.3 Wuchiapingian Clarkina sp. 19.2 JY 20 Jiangya section Hunan China 256.4 Wuchiapingian Clarkina sp. 19.9 JY 21 Jiangya section Hunan China 256.5 Wuchiapingian Clarkina sp. 19.8 GYD 145 Dukou section Sichuan China 256.6 Wuchiapingian Clarkina sp. 20.8

127 Tables 128

Sample Age(Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

GYD 142.3 Dukou section Sichuan China 256.6 Wuchiapingian Clarkina sp. 20.0 JY 22 Jiangya section Hunan China 256.7 Wuchiapingian Clarkina sp. 19.8 JY 23 Jiangya section Hunan China 256.7 Wuchiapingian Clarkina sp. 20.0 GYD 138.3 Dukou section Sichuan China 256.7 Wuchiapingian Clarkina sp. 20.2 JY 24 Jiangya section Hunan China 256.9 Wuchiapingian Clarkina sp. 19.8 GYD 132 Dukou section Sichuan China 256.9 Wuchiapingian Clarkina sp. 20.4 GYD 131.5 Dukou section Sichuan China 256.9 Wuchiapingian Clarkina sp. 19.8 GYD 130.0 Dukou section Sichuan China 257.0 Wuchiapingian Clarkina sp. 19.8 GYD 126.6 Dukou section Sichuan China 257.1 Wuchiapingian Clarkina sp. 19.5 GYD 123 Dukou section Sichuan China 257.2 Wuchiapingian Clarkina sp. 20.7 JY 26 Jiangya section Hunan China 257.3 Wuchiapingian Clarkina sp. 19.5 GYD 114.5 Dukou section Sichuan China 257.4 Wuchiapingian Clarkina sp. 20.5 GYD 114.2 Dukou section Sichuan China 257.4 Wuchiapingian Clarkina sp. 20.1 JY 27 Jiangya section Hunan China 257.4 Wuchiapingian Clarkina sp. 19.8 GYD 110 Dukou section Sichuan China 257.5 Wuchiapingian Clarkina sp. 19.5 GYD 104 Dukou section Sichuan China 257.7 Wuchiapingian Clarkina sp. 20.3 GYD 104 R+M Dukou section Sichuan China 257.7 Wuchiapingian Ramiform 19.9 GYD 99.65 Dukou section Sichuan China 257.8 Wuchiapingian Clarkina sp. 20.3 GYD 93.5 Dukou section Sichuan China 257.9 Wuchiapingian Clarkina sp. 20.2 GYD 93.5 Dukou section Sichuan China 257.9 Wuchiapingian Clarkina sp. 20.5 GYD 93.5-R Dukou section Sichuan China 257.9 Wuchiapingian Ramiform 20.0 GYD 91.2 Dukou section Sichuan China 258.0 Wuchiapingian Clarkina sp. 20.3 GYD 87.7 Dukou section Sichuan China 258.1 Wuchiapingian Clarkina sp. 20.0 GYD 87.0 Dukou section Sichuan China 258.1 Wuchiapingian Clarkina sp. 20.0

Tables

Sample Age(Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

PLTN(19.1-19.23) PenglaitanN section Guangxi China 258.2 Wuchiapingian C. orientalis Clarkina orientalis 19.9 TQ 63 Tieqiao section Guangxi China 258.2 Wuchiapingian C. liangshanensis Clarkina sp. 18.9 GYD 84 Dukou section Sichuan China 258.2 Wuchiapingian Clarkina sp. 20.4 GYD 82.0 Dukou section Sichuan China 258.2 Wuchiapingian Clarkina sp. 20.4 TQ 62 Tieqiao section Guangxi China 258.3 Wuchiapingian C. liangshanensis Clarkina sp. 18.6 PLTNC-15(7.97-8.10) PenglaitanN section Guangxi China 258.5 Wuchiapingian C. liangshanensis Clarkina liangshanensis 19.6 TQ 61 Tieqiao section Guangxi China 258.5 Wuchiapingian C. liangshanensis Clarkina sp. 19.3 GYD 71.95 Dukou section Sichuan China 258.5 Wuchiapingian Clarkina sp. 20.3 PLTNC-11(5.63-5.73) PenglaitanN section Guangxi China 258.5 Wuchiapingian C. liangshanensis Clarkina liangshanensis 20.0 PLTNC-8(3.75-3.98) PenglaitanN section Guangxi China 258.5 Wuchiapingian C. liangshanensis Clarkina liangshanensis 18.9 PLTNC-9(4.68-4.83) PenglaitanN section Guangxi China 258.6 Wuchiapingian C. liangshanensis Clarkina liangshanensis 19.7 PLTNC-6(2.27-2.52) PenglaitanN section Guangxi China 258.6 Wuchiapingian C. liangshanensis Clarkina liangshanensis 19.7 PLTNC-1(0.14-0.55) PenglaitanN section Guangxi China 258.7 Wuchiapingian C. liangshanensis Clarkina liangshanensis 20.0 TQ 60 Tieqiao section Guangxi China 258.7 Wuchiapingian C. liangshanensis Clarkina sp. 19.1 GYD 60.2 Dukou section Sichuan China 258.8 Wuchiapingian Clarkina sp. 19.5 TQ 58 Tieqiao section Guangxi China 259.0 Wuchiapingian C. leveni Clarkina sp. 18.7 TQ 56 Tieqiao section Guangxi China 259.2 Wuchiapingian C. leveni C. asymmetrica, C. dukouensis 19.7

TQ 57 Tieqiao section Guangxi China 259.2 Wuchiapingian C. leveni C. asymmetrica, C. dukouensis 19.1 GYD 45 Dukou section Sichuan China 259.2 Wuchiapingian Clarkina sp. 20.7 GYD 41.9 Dukou section Sichuan China 259.3 Wuchiapingian Clarkina sp. 20.6 GYD 39.9 Dukou section Sichuan China 259.3 Wuchiapingian Clarkina sp. 19.6 GYD 37.9 Dukou section Sichuan China 259.4 Wuchiapingian Clarkina sp. 20.4 GYD 35.6 Dukou section Sichuan China 259.5 Wuchiapingian Clarkina sp. 20.1

129 Tables 130

Sample Age(Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

GYD 33.8 Dukou section Sichuan China 259.5 Wuchiapingian Clarkina sp. 20.4 GYD 31.9 Dukou section Sichuan China 259.6 Wuchiapingian Clarkina sp. 20.8 GYD 30.2 Dukou section Sichuan China 259.6 Wuchiapingian Clarkina sp. 20.0 GYD 28.4 Dukou section Sichuan China 259.7 Wuchiapingian Clarkina sp. 21.2 GYD 26.45 Dukou section Sichuan China 259.7 Wuchiapingian Clarkina sp. 20.3 GYD 23.6 Dukou section Sichuan China 259.8 Wuchiapingian Clarkina sp. 20.7 GYD 20.8 Dukou section Sichuan China 259.9 Wuchiapingian Clarkina sp. 19.8 PLTS 19.7 Penglaitan section Guangxi China 259.9 Wuchiapingian C. asymmetrica Clarkina sp. 21.5 GYD 18.3 Dukou section Sichuan China 259.9 Wuchiapingian Clarkina sp. 20.4 GYD 15.6 Dukou section Sichuan China 260.0 Wuchiapingian Clarkina sp. 20.3 PLTS 14.05 Penglaitan section Guangxi China 260.0 Wuchiapingian C. asymmetrica Clarkina sp. 21.8 GYD 13.6 Dukou section Sichuan China 260.0 Wuchiapingian Clarkina sp. 20.5 PLT 11.4 Penglaitan section Guangxi China 260.1 Wuchiapingian C. dukouensis Clarkina sp. 21.1 PLTS 10.36 Penglaitan section Guangxi China 260.1 Wuchiapingian C. dukouensis Clarkina sp. 20.8 PLTS 9.67 Penglaitan section Guangxi China 260.1 Wuchiapingian C. dukouensis Clarkina sp. 20.9 TQ 54 Tieqiao section Guangxi China 260.2 Wuchiapingian C. dukouensis C. postbitteri, C. asymmetrica 19.9 PLTS 6.59 Penglaitan section Guangxi China 260.2 Wuchiapingian C. dukouensis Clarkina sp. 21.0 PLTS 6.19 Penglaitan section Guangxi China 260.2 Wuchiapingian C. dukouensis Clarkina sp. 20.5 PLTS 5.8 Penglaitan section Guangxi China 260.2 Wuchiapingian C. dukouensis Clarkina sp. 19.9 GYD 5 Dukou section Sichuan China 260.3 Wuchiapingian Clarkina sp. 21.0 PLTS 4.43 Penglaitan section Guangxi China 260.3 Wuchiapingian C. dukouensis Clarkina sp. 21.1 GYD 2.7 Dukou section Sichuan China 260.3 Wuchiapingian Clarkina sp. 20.6 PLTS 1.08 Penglaitan section Guangxi China 260.4 Wuchiapingian C. postbitteri postbitteri Clarkina sp. 20.6 TQ 2 Tieqiao section Guangxi China 260.4 Wuchiapingian C. postbitteri postbitteri C. postbitteri postbitteri 20.6

Tables

Sample Age(Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

TQ 1 Tieqiao section Guangxi China 260.4 Wuchiapingian C. postbitteri postbitteri C. postbitteri postbitteri 20.9 TQ 3 Tieqiao section Guangxi China 260.4 Wuchiapingian C. postbitteri postbitteri C. postbitteri hongshuiensis 20.1 PL1 Penglaitan section Guangxi China 260.4 Wuchiapingian C. postbitteri postbitteri C. postbitteri postbitteri 20.7 TQ 5 Tieqiao section Guangxi China 260.4 Capitanian C. postbitteri,hongshuiensis C. postbitteri, hongshuiensis 20.3 PL7 Penglaitan section Guangxi China 260.5 Capitanian J. granti Ramiform elements 19.9 PL7 Penglaitan section Guangxi China 260.5 Capitanian J. granti J. granti, J. xuanhanesis 20.3 PL8 Penglaitan section Guangxi China 260.5 Capitanian J. granti J. granti, J. xuanhanesis, 19.8 PL9 Penglaitan section Guangxi China 260.5 Capitanian J. granti J. granti 20.1 PL10 Penglaitan section Guangxi China 260.5 Capitanian J. granti J. granti, J. xuanhanensis 20.3 PL11 Penglaitan section Guangxi China 260.6 Capitanian J. granti J. granti 20.1 Hindeodus sp., Ramiform PL13 Penglaitan section Guangxi China 260.7 Capitanian J. granti 19.1 elements TQ 7 Tieqiao section Guangxi China 260.7 Capitanian J. granti J. granti 19.7 TQ 8 Tieqiao section Guangxi China 260.7 Capitanian J. granti J. granti 20.2 TQ 8 Tieqiao section Guangxi China 260.7 Capitanian J. granti Ramiform elements 20.1 TQ 9 Tieqiao section Guangxi China 260.7 Capitanian J. granti J. granti 20.0 TQ 10 Tieqiao section Guangxi China 260.7 Capitanian J. granti J. granti 20.3 TQ 11 Tieqiao section Guangxi China 260.7 Capitanian J. granti J. granti 19.6 TQ 12 Tieqiao section Guangxi China 260.7 Capitanian J. granti J. granti 20.0 Hindeodus sp., Ramiform PL14 Penglaitan section Guangxi China 260.7 Capitanian J. xuanhanensis 19.0 elements PL15 Penglaitan section Guangxi China 260.8 Capitanian J. xuanhanensis J. xuanhanensis 20.3 TQ 13 Tieqiao section Guangxi China 260.8 Capitanian J. granti J. granti 19.9 Hindeodus, Ramiform TQ 14 Tieqiao section Guangxi China 260.8 Capitanian J. granti 19.1 elements PL18 Penglaitan section Guangxi China 261.0 Capitanian J. xuanhanensis Ramiform elements 20.3 PL18 Penglaitan section Guangxi China 261.0 Capitanian J. xuanhanensis J. xuanhanensis, J. altudensis 20.5

131 Tables 132

Sample Age(Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

J. xuanhanensis, Ramiform TQ 16 Tieqiao section Guangxi China 261.1 Capitanian J. xuanhanensis 19.9 elements, Hindeodus sp. J. xuanhanensis, PL19 Penglaitan section Guangxi China 261.1 Capitanian J. xuanhanensis J. prexuanhanensis, Ramifrom 20.3 elements J. xuanhanensis, J. TQ 17 Tieqiao section Guangxi China 261.2 Capitanian J. xuanhanensis 20.1 prexuanhanensis TQ 18 Tieqiao section Guangxi China 261.4 Capitanian J. xuanhanensis J. xuanhanensis 20.2 GM49 EF Roadcut Guadalupe Mts. USA 262.0 Capitanian J. postserrata? Jinogondolella sp. 21.9 PL26 Penglaitan section Guangxi China 262.2 Capitanian J. shannoni J. sp., Ramiform elements 20.9 GM 38 Williams Ranch Road 1 Guadalupe Mts. USA 262.2 Capitanian J. postserrata? Jinogondolella sp. 20.8 South Boundary Section GM 33 Guadalupe Mts. USA 262.2 Capitanian J. postserrata? Jinogondolella sp. 21.4 II GM46 EF Roadcut Guadalupe Mts. USA 262.4 Capitanian J. postserrata Jinogondolella sp. 20.9 GM 37 Williams Ranch Road 1 Guadalupe Mts. USA 263.2 Capitanian J. postserrata Jinogondolella sp. 21.3 TQ 29 Tieqiao section Guangxi China 263.3 Capitanian J. shannoni J. shannoni 19.2 TQ 31 Tieqiao section Guangxi China 263.7 Capitanian J. postserrata J. postserrata, J. shannoi? 20.9 TQ 33 Tieqiao section Guangxi China 263.9 Capitanian J. postserrata J. postserrata, Hindeodus sp. 20.9 J. postserrata, J. shannoni, TQ 35 Tieqiao section Guangxi China 264.2 Capitanian J. postserrata 20.6 Ramiform elements J. postserrata, Hindeodus sp., TQ 37 Tieqiao section Guangxi China 264.5 Capitanian J. postserrata 20.8 Ramiform elements Rader slide Roadcut GM 13 Guadalupe Mts. USA 264.6 Capitanian J. postserrata Jinogondolella sp. 21.5 on US 62/180 J. postserrata, Hindeodus sp., TQ 38 Tieqiao section Guangxi China 264.7 Capitanian J. postserrata 20.7 Ramiform elements Rader Roadcut on US 62/180, little outcrop GM 29 Guadalupe Mts. USA 264.8 Capitanian J. postserrata Jinogondolella sp. 21.8 NE of Rader slide roadcut on US 62/180

Tables

Sample Age(Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

First cliff above lunch rock GM 34 on Josh's section in Guadalupe Mts. USA 264.9 Capitanian J. postserrata Jinogondolella sp. 21.4 Patterson Hills; Rader slide Roadcut on GM 28 Guadalupe Mts. USA 264.9 Capitanian J. postserrata Jinogondolella sp. 21.3 US 62/180 J. postserrata, Hindeodus sp., TQ 40 Tieqiao section Guangxi China 265.1 Capitanian J. postserrata 20.3 Ramiform elements J. postserrata, Hindeodus sp., TQ 41 Tieqiao section Guangxi China 265.1 Capitanian J. postserrata 20.8 Ramiform elements TQ 42 Tieqiao section Guangxi China 265.4 Capitanian J. postserrata J. postserrata, Ramiform elements 20.5 TQ 42 Tieqiao section Guangxi China 265.4 Capitanian J. postserrata J. postserrata 20.5 TQ 43 Tieqiao section Guangxi China 265.5 Capitanian J. postserrata J. postserrata 20.7 TQ 43 Tieqiao section Guangxi China 265.5 Capitanian J. postserrata J. postserrata, Ramiform elements 20.5 TQ 44 Tieqiao section Guangxi China 265.8 Wordian J. aserrata J. aserrata, Ramiform elements 20.8 East Patterson Hills - East GM 26 Slope (section of Mary & Guadalupe Mts. USA 265.8 Wordian J. aserrata Jinogondolella sp. 21.5 Brandon) East Patterson Hills - East GM 25 Slope (section of Mary & Guadalupe Mts. USA 265.9 Wordian J. aserrata Jinogondolella sp. 21.4 Brandon) East Patterson Hills - East GM 24 Slope (section of Mary & Guadalupe Mts. USA 266.0 Wordian J. aserrata Jinogondolella sp. 21.9 Brandon) J. aserrata, Hindeodus sp., TQ 45 Tieqiao section Guangxi China 266.1 Wordian J. aserrata 20.2 Ramiform elements Manzanita Patterson Hill GM 8 Guadalupe Mts. USA 266.2 Wordian J. aserrata Jinogondolella sp. 21.8 Roadcut on US 62/177 Manzanita Patterson Hill GM 7 Guadalupe Mts. USA 266.2 Wordian J. aserrata Jinogondolella sp. 21.6 Roadcut on US 62/179

133 Tables 134

Sample Age(Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

Manzanita Patterson GM7 Hill Roadcut on US Guadalupe Mts. USA 266.2 Wordian J. aserrata Ramiform 21.6 62/178 Lunch Rock, access GM 35 Guadalupe Mts. USA 266.3 Wordian J. aserrata Jinogondolella sp. 21.2 ravine on Patterson Hill Manzanita Patterson GM 5 Hill Roadcut on US Guadalupe Mts. USA 266.4 Wordian J. aserrata Ramiform 21.6 62/180 Manzanita Patterson GM 5 Hill Roadcut on US Guadalupe Mts. USA 266.4 Wordian J. aserrata Jinogondolella sp. 22.0 62/181 East Patterson Hills - GM 23 East Slope (section of Guadalupe Mts. USA 266.4 Wordian J. aserrata Jinogondolella sp. 21.4 Mary & Brandon) J. aserrata, Ramiform TQ 46 Tieqiao section Guangxi China 266.4 Wordian J. aserrata 20.9 elements Manzanita Patterson GM 4 Hill Roadcut on US Guadalupe Mts. USA 266.4 Wordian J. aserrata Jinogondolella sp. 21.8 62/182 East Patterson Hills - GM 22 East Slope (section of Guadalupe Mts. USA 266.4 Wordian J. aserrata Jinogondolella sp. 21.4 Mary & Brandon) J. aserrata, Hindeodus sp., TQ 47 Tieqiao section Guangxi China 267.1 Wordian J. aserrata 20.6 Ramiform elements TQ 48 Tieqiao section Guangxi China 267.6 Wordian J. aserrata J. aserrata 20.8 NS 63 Naqing section Guizhou China 269.6 Roadian J. nanjingensis Jinogondolella sp. 20.7 NS 65 Naqing section Guizhou China 270.1 Roadian J. nanjingensis Jinogondolella sp. 21.7 GM 3 Type Pipeline Shale Guadalupe Mts. USA 270.2 Roadian J. nanjingensis Jinogondolella sp. 21.0 GM 2 Type Pipeline Shale Guadalupe Mts. USA 270.3 Roadian J. nanjingensis Jinogondolella sp. 21.1 GM 1 Type Pipeline Shale Guadalupe Mts. USA 270.3 Roadian J. nanjingensis Jinogondolella sp. 21.3 NSC 714.7 Naqing section Guizhou China 270.5 Roadian J. nanjingensis Jinogondolella sp.Ramiform 21.5

Tables

Sample Age(Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

Spinosa Roadcut GM 10 Guadalupe Mts. USA 270.5 Roadian J. nanjingensis Jinogondolella sp. 21.2 Quarry on US 62/181 Stratotype Canyon GM 16 (Kungurian/Roadian Guadalupe Mts. USA 270.6 Kungurian J. nanjingensis Jinogondolella sp. 21.0 GSSP) Stratotype Canyon GM 19 (Kungurian/Roadian Guadalupe Mts. USA 270.6 Kungurian M. Idahoensis Jinogondolella sp. 21.1 GSSP) NS 60 Naqing section Guizhou China 270.9 Kungurian S. hanzhougensis-M. idahoensis Ramiform 21.1 NS 57 Naqing section Guizhou China 271.7 Kungurian S. subsymmertricus Ramiform 21.9 NS 53 Naqing section Guizhou China 272.7 Kungurian S. subsymmertricus-M. Siciliensis Mesogondolella sp. 22.0 NSC 540.7 Naqing section Guizhou China 273.1 Kungurian S. subsymmertricus-M. Siciliensis Mesogondolella sp. 20.9 NS 51 Naqing section Guizhou China 273.2 Kungurian S. subsymmertricus-M. Siciliensis Mesogondolella sp. 22.0 NSC 520.1 Naqing section Guizhou China 273.4 Kungurian S. subsymmertricus-M. Siciliensis Mesogondolella sp. 21.5 Mesogondolella sp., NSC 511.3 Naqing section Guizhou China 273.6 Kungurian S. subsymmertricus-M. Siciliensis 22.1 Ramiform NS 49 Naqing section Guizhou China 273.7 Kungurian S. subsymmertricus-M. Siciliensis Mesogondolella sp. 22.0 Mesogondolella sp., NSC 490 Naqing section Guizhou China 273.9 Kungurian S. subsymmertricus-M. Siciliensis 21.9 Ramiform NSC 486 Naqing section Guizhou China 274.0 Kungurian S. subsymmertricus-M. Siciliensis Mesogondolella sp. 21.9 NSC 485.6 Naqing section Guizhou China 274.0 Kungurian S. subsymmertricus-M. Siciliensis Mesogondolella sp. 21.7 NS 48 Naqing section Guizhou China 274.0 Kungurian S. subsymmertricus-M. Siciliensis Mesogondolella sp. 22.2 NS 48 R Naqing section Guizhou China 274.0 Kungurian S. subsymmertricus-M. Siciliensis Ramiform 22.7 Mesogondolella sp., NSC 478.5 Naqing section Guizhou China 274.1 Kungurian S. subsymmertricus-M. Siciliensis 22.1 Ramiform NSC 453.8 Naqing section Guizhou China 274.4 Kungurian S. guizhouensis Sweetognathus.sp 21.7 NSC 436.7 Naqing section Guizhou China 274.7 Kungurian S. guizhouensis Sweetognathus.sp 22.2

135 Tables 136

Sample Age(Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

NS 40 Naqing section Guizhou China 275.5 Kungurian N. pnevi Mesogondolella sp. 22.0 NSC 376.2 Naqing section Guizhou China 275.6 Kungurian N. pnevi Mesogondolella sp. 21.4 Mesogondolella sp., NSC 378.5 Naqing section Guizhou China 275.6 Kungurian N. pnevi 21.4 ramiform NSC 375.8 Naqing section Guizhou China 275.7 Artinskian S. whitei-M. bisselli Sweetognathus sp. 21.9 NS 39 Naqing section Guizhou China 276.9 Artinskian S. whitei-M. bisselli Mesogondolella sp. 21.9 NS 38 Naqing section Guizhou China 278.5 Artinskian S. whitei-M. bisselli Mesogondolella sp. 22.7 LD 369 Naqing section Guizhou China 280.0 Artinskian S. whitei-M. bisselli Mesogondolella sp. 21.9 NS 37 Naqing section Guizhou China 280.0 Artinskian S. whitei-M. bisselli Mesogondolella sp. 22.3 NS 37 Naqing section Guizhou China 280.0 Artinskian S. whitei-M. bisselli Ramiform 21.6 NS 36 Naqing section Guizhou China 281.8 Artinskian S. whitei-M. bisselli Ramiform 21.6 NS 35 Naqing section Guizhou China 283.1 Artinskian S. whitei-M. bisselli Mesogondolella sp. 22.2 NS 34 Naqing section Guizhou China 284.4 Sakmarian S. binodosus Mesogondolella sp. 22.8 NSC 362.0 Naqing section Guizhou China 284.4 Sakmarian S. binodosus Mesogondolella sp. 22.0 NS 33 Naqing section Guizhou China 285.2 Sakmarian S. binodosus Mesogondolella sp. 22.2 Mesogondolella sp., NS 31 Naqing section Guizhou China 286.4 Sakmarian S. binodosus 22.7 ramiform NSC 353.5 Naqing section Guizhou China 286.7 Sakmarian S. binodosus Mesogondolella sp. 21.9 NSC 350.9 Naqing section Guizhou China 287.4 Sakmarian S. binodosus Mesogondolella sp. 21.0 NS 28 Naqing section Guizhou China 288.5 Sakmarian S. binodosus Mesogondolella sp. 22.1 NS 28mixed Naqing section Guizhou China 288.5 Sakmarian S. binodosus Ramiform 22.6 NS 25 Naqing section Guizhou China 290.3 Sakmarian M. monstra Ramiform 22.4 NS 25M Naqing section Guizhou China 290.3 Sakmarian M. monstra Mesogondolella sp. 22.6 NS 24 Naqing section Guizhou China 290.7 Sakmarian M. monstra Mesogondolella sp. 22.7

Tables

Sample Age(Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

Mesogondolella sp., NS 71 Naqing section Guizhou China 290.9 Sakmarian M. monstra 23.0 Streptognathodus sp. NS 23 Naqing section Guizhou China 291.4 Sakmarian M. monstra Mesogondolella sp. 22.5 NS 22 Naqing section Guizhou China 291.9 Sakmarian M. monstra Mesogondolella sp. 22.8 NSC 330.8 Naqing section Guizhou China 292.8 Sakmarian M. monstra Mesogondolella sp. 22.0 NSC 330.6 Naqing section Guizhou China 292.8 Sakmarian M. monstra Mesogondolella sp. 22.3 NS 19 Naqing section Guizhou China 293.5 Sakmarian M. monstra Mesogondolella sp. 20.3 NS 70 Naqing section Guizhou China 294.7 Asselian St. fusus Mesogondolella sp. 22.4 NSC 318.7 Naqing section Guizhou China 295.2 Asselian St. fusus Streptognathodus sp. 22.6 NSC 315.6 Naqing section Guizhou China 295.5 Asselian St. fusus Streptognathodus sp. 22.3 NS 14 Naqing section Guizhou China 295.6 Asselian St. fusus Streptognathodus sp. 22.3 NSC 313 Naqing section Guizhou China 295.8 Asselian St. fusus Streptognathodus sp. 21.2 NS 13 Naqing section Guizhou China 295.9 Asselian St. fusus Streptognathodus sp. 22.4 Streptognathodus sp., NS 12 Naqing section Guizhou China 296.2 Asselian St. fusus 22.2 ramform NS 11 Naqing section Guizhou China 296.4 Asselian St. fusus Streptognathodus sp. 22.2 NSC 307.8 Naqing section Guizhou China 296.4 Asselian St. fusus Streptognathodus sp. 22.8 NS 8 Naqing section Guizhou China 297.0 Asselian St. fusus St. sp. and M. sp. 23.0 NSC 302.3 Naqing section Guizhou China 297.0 Asselian St. fusus Streptognathodus sp. 22.5 NS 7 Naqing section Guizhou China 297.3 Asselian St. fusus Streptognathodus sp. 22.9 NSC 294.6 Naqing section Guizhou China 297.9 Asselian St. constrictus Streptognathodus sp. 22.1 NS 5 Naqing section Guizhou China 298.1 Asselian St. sigmoidalis Streptognathodus sp. 22.5 NSC 288.5 Naqing section Guizhou China 298.6 Asselian St. sigmoidalis Streptognathodus sp. 21.9 NSC 286 Naqing section Guizhou China 298.9 Asselian St. sigmoidalis Streptognathodus sp. 21.2

137 Tables 138

Sample Age(Gradstein Analysed conodont δ18O Location Area Stage Conodont biozone apatite ident et al. 2004) taxa (VSMOW)

NS 1 Naqing section Guizhou China 298.9 Asselian St. sigmoidalis Streptognathodus sp. 22.4 NSC 285.1 Naqing section Guizhou China 299.0 Asselian St. sigmoidalis Streptognathodus sp. 22.9 NSC 284.6 Naqing section Guizhou China 299.0 Gzhelian St. sigmoidalis Streptognathodus sp. 21.9 NSC 283 Naqing section Guizhou China 299.2 Gzhelian St. sigmoidalis Streptognathodus sp. 22.3 NSC 281.2 Naqing section Guizhou China 299.4 Gzhelian St. isolatus Streptognathodus sp. 21.8 NSC 281.2 Naqing section Guizhou China 299.4 Gzhelian St. isolatus Streptognathodus sp. 21.5

Acknowledgements

Acknowledgements

First and foremost, I would like to thank my supervisor Professor Michael M. Joachimski. Michael gave his valuable guidance and suggestions in every phase of this work with patience and encouragement. This work would not have been possible without his help. Michael’s supervision during the past three years was a pleasure and unforgettable experience for me, which not just provided me the best scientific training, but also gave me a chance to learn a different culture. Both will be significant to my future scientific career and life. And I am looking forward to work with Michael again in the future. I gratefully acknowledge my colleagues at the Nanjing Institute of Geology and Palaeontology, Chinese Academy of Sciences. Prof. Wang Xiangdong provided significant assistance in field work and samples shipping. Prof. Shen Shuzhong generously contributed conodont samples from Iran and South China and gave constructive suggestions which considerably improved the thesis. Prof. Cao Changqun, Prof. Wang Wei, Prof. Wang Yue, Prof. Qi Yupin, Dr. Chen Jun, Dr. Zheng QuanFeng, Dr. Zhang Yiqiang and Mr. Yuan Dongxun are greatly thanked for providing support during the field work in China. I am also would like to express my sincere thanks to colleagues from Wuhan (China University of Geoscience). Prof. Lai Xulong and Sun Yadong supported me during the field work and indentified the conodonts from Laibin area. Prof. Lai Xulong, Dr. Jiang Haishui, Dr. Luo Genming provided P–T boundary samples and critically reviewed the P–T boundary part of this thesis. Dr. Lance Lambert from the University of Texas at San Antonio is greatly acknowledged for assisting the collection of samples in the Guadalupe Mountains, improving English style as well as giving valuable comments on the palaeocology of Permian conodonts.

139 Acknowledgements

Ms. Daniele Lutz is sincerely acknowledged for generously teaching me the detailed procedure of preparing conodont samples for oxygen isotope analysis, and Claudio Filomena, Henning Shulz, Inga Osbahr, Stefan Schöbel and many other colleagues at the GeoZentrum Nordbayern are warmly thanked for their kind help during my stay in Erlangen. Last but not least, I take this opportunity to sincerely thank my family; their unconditional support made me focus fully on my work. I apologize to them for every moment I should have been with them but I could not. This work was financially supported by the German Science Foundation DFG (project JO 219/9-1), and I was additionally funded by a scholarship (Stipendium der Universität aus Mitteln des STIBET-Doktorandenprogramm des DAAD), which I gratefully acknowledge.

140 Curriculum Vitae

Curriculum vitae

Chen Bo

Address GeoZentrum Nordbayern der Universität Erlangen-Nürnberg Schloßgarten 5 D-91054 Erlangen, Germany Email: [email protected] Tel: +49 (0) 9131 85 26066

Private Date of Birth: April 27th, 1981 Place of Birth: Hunan, China Nationality: China

Education 2001–2005 B.S., Civil Engineering Department, Hunan University of Science and Technology, major in

Prospecting Technology and Engineering 2005–2008 M.S., Nanjing Institute of Geology and Palaeontology, Chinese Academy of Sciences, major in Paleontology and Stratigraphy 2008–2009 Ph.D. candidate, Nanjing Institute of Geology and Palaeontology, Chinese Academy of Sciences, major in

Paleontology and Stratigraphy 2009–present Ph.D. candidate of isotope geochemistry in GeoZentrum Nordbayern Universität Erlangen-Nürnberg

141