Palaeoenvironmental Changes and Vegetation History During the Triassic-Jurassic Transition Nina R

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Palaeoenvironmental changes and vegetation history during the Triassic-Jurassic transition Nina R. Bonis Palaeoecology Institute of Environmental Biology Department of Biology Faculty of Science Utrecht University

Laboratory of Palaeobotany and Palynology Budapestlaan 4 3584 CD Utrecht The Netherlands [email protected] [email protected]

ISBN 978-90-393-5269-4 NSG publication No. 20100129 LPP Contribution Series No. 29

Graphic design by Nick Liefhebber www.liefhebber.biz [email protected]

Printed by GVO drukkers & vormgevers B.V. | Ponsen & Looijen, Ede Palaeoenvironmental changes and vegetation history during the Triassic-Jurassic transition

Palaeomilieu veranderingen en vegetatiegeschiedenis gedurende de Trias-Jura overgang

(met een samenvatting in het Nederlands)

Proefschrift ter verkrijging van de graad van doctor aan de Universiteit Utrecht op gezag van de rector magnificus, prof.dr. J.C. Stoof, ingevolge het besluit van het college voor promoties in het openbaar te verdedigen op vrijdag 29 januari 2010 des ochtends te 10.30 uur

door

Nina Rosa Bonis geboren op 13 februari 1983 te Oosterhout

Promotor: Prof. dr. A.F. Lotter

Co-promotor: Dr. W.M. Kürschner

The research reported in this thesis was funded by the ‘High Potential’ stimulation program of Utrecht University and financially supported by the LPP foundation.

voor Joeri Contents

General introduction and synopsis 10

Chapter 1 A detailed palynological study of the Triassic 16 - Jurassic transition in key sections of the Eiberg Basin (Northern Calcareous Alps, Austria) with W. M. Kürschner and L. Krystyn

Published in Review of Palaeobotany and Palynology 156, 376-400 (2009)

Chapter 2 Climate change driven black shale deposition 48 during the end-Triassic in the western Tethys with M. Ruhl and W. M. Kürschner

Published in Palaeogeography, Palaeoclimatology, Palaeoecology. Special Issue: Triassic climate (in press)

Chapter 3 Abrupt climate change during the 64 end-Triassic mass-extinction with W. M. Kürschner

To be submitted

Chapter 4 Atmospheric methane injection caused 78 end-Triassic mass extinction with M. Ruhl, G.J. Reichart, J.S. Sinninghe Damsté and W. M. Kürschner

To be submitted

Chapter 5 Milankovitch-scale palynological turnover 90 across the Triassic - Jurassic transition at St. Audrie’s Bay, SW UK with M. Ruhl and W. M. Kürschner

Submitted to Journal of the Geological Society Chapter 6 Changing CO2 conditions during the end-Triassic 116 inferred from stomatal frequency analysis on Lepidopteris ottonis (Goeppert) Schimper and Ginkgoites taeniatus (Braun) Harris with W. M. Kürschner and J. H. A. Van Konijnenburg-Van Cittert

Submitted to Palaeogeography, Palaeoclimatology, Palaeoecology

Chapter 7 Vegetation history, diversity patterns, and climate 140 change across the Triassic-Jurassic boundary with W. M. Kürschner

Submitted to Paleobiology

References 166

Algemene inleiding en samenvatting 188

Dankwoord/Acknowledgements 194

Curriculum Vitae 196

Publications 197

Color Figures 199 SYNOPSIS

General introduction and synopsis

The Triassic-Jurassic (T-J) boundary, ~201.58 Ma (Schaltegger et al., 2008), is generally known as one of the ‘big five’ mass extinction events in Earth’s history (e.g., Newell, 1963; Raup and Sepkoski, 1982; Benton, 1995; Taylor, 2004). It is one of the least studied because complete and well-dated sections are scarce. To study this end-Triassic mass extinction event, a multidisciplinary project was initiated: ‘Earth’s and life’s history: from core to biosphere (CoBi)’. With a combined use of palaeomagnetism, cyclostratigraphy, geochemistry, and biogeology, questions about the timing, cause and patterns of the extinction were addressed. This thesis will focus on the biogeological aspect by using two techniques: palynology, and stomatal frequency analysis on fossil leaves.

Explanations for the biotic turnover during the late Triassic have included both gradualistic and catastrophic mechanisms (e.g., Hallam and Wignall, 1997; Tanner et al., 2004; Hesselbo et al., 2007). Marine extinction could be related to shelf habitat loss during severe regression in the end Triassic (Hallam and Wignall, 1999), but this does not account for the extinctions in the terrestrial realm (e.g., tetrapods). A frequently proposed mechanism is massive volcanism of the Central Atlantic Magmatic Province (CAMP), one of the largest known Phanero- zoic flood basalt provinces, related to the breakup of Pangaea (Wignall, 2001; Hesselbo et al., 2002; Knight et al., 2004; Marzoli et al., 1999, 2004; Schaltegger et al., 2008). This CAMP basalt volcanism can be associated with the following effects: rapid global warming, ocean anoxia or increased oceanic fertilization (or both), calcification crises, and a sharp decrease in carbon isotope values (Wignall, 2005). The T-J transition interval is characterized by two major perturbations in carbon isotope records described from many sections within and outside the Tethys realm. An end-Triassic relatively short lived ‘initial’ negative carbon isotope excursion (CIE) precedes a more gradual excursion (‘main’ CIE) at the base of the Jurassic (e.g., Pálfy et al., 2001; Hesselbo et al., 2002; Guex et al., 2004; Kürschner et al., 2007; Ward

et al., 2007; Ruhl et al., 2009). The release of large amounts of carbon dioxide (CO2) in the atmosphere by CAMP volcanism induced climate change and could have caused biotic disturbance (McElwain et al., 1999; Hesselbo et al., 2002; Tanner et al., 2004). Methane release from gas hydrates represents another important event that has been suggested to be associated with the end-Triassic volcanism (e.g., Pálfy et al., 2001; Beerling and Berner, 2002; Wignall, 2005). An alternative catastrophic mechanism causing the extinction is an extraterrestrial impact (Olsen et al., 2002a, b) but to date, no convincing evidence has been found for an impact-caused mass extinction.

Examples of end-Triassic biotic disturbances are: an extinction among tetrapods (e.g., Olsen et al., 2002b; Lucas and Tanner, 2007a), abrupt and substantial changes in the composition of brachiopod and bivalve communities (Hallam, 1981; Kiessling et al., 2007; Tomašových and

10 SYNOPSIS

Siblík, 2007; Mander et al., 2008), the final extinction of conodonts with sporadic survivors persisting into the Hettangian (Pálfy et al., 2007), a global radiolarian faunal change (Long- ridge et al., 2007; Pálfy et al., 2007), the final disappearance of the already low-diversity ammonite assemblages (Simms and Ruffell, 1990; Hallam, 2002) and a collapse of reef ecosystems (e.g., Kiessling et al., 2007). However, the severity and patterns (i.e., abrupt, stepwise or gradual) of the end-Triassic extinctions are disputed (e.g., Hallam, 2002; Bambach et al., 2004; Tanner et al., 2004; Lucas and Tanner, 2008).

Although the T-J transition is characterized by extinctions in the marine realm, evidence for a floral turnover is ambiguous. A recent study from the Germanic Basin showed a severe vegetation shift across the T-J boundary, linked to CAMP volcanism (Van de Schootbrugge et al., 2009). A major extinction of 60% of sporomorph taxa followed by a sharp spore spike at the T-J boundary is claimed in the Newark Basin, USA (Fowell and Olsen, 1993; Fowell et al., 1994; Olsen et al., 2002a, b). This spore spike is followed by a conifer (Cheirolepidiaceae) pollen dominated palynoflora which is used by Cornet (1977) and Fowell et al. (1994) to mark the base of the Jurassic in the Newark Basin. By contrast, most palynological studies from Europe show gradual changes in assemblages at the T-J transition (e.g., Warrington, 1974; Morbey, 1975; Lund, 1977; Schuurman, 1979; Achilles, 1981; Kürschner et al., 2007). Also the T-J plant macrofossil record is equivocal. Quantitative macrobotanical data from East Greenland showed that Triassic forests with high-diversity communities were replaced by lower diversity forests and that there was a gradual extinction prior to the T-J boundary (McElwain and Punyasena, 2007; McElwain et al., 2007). On the contrary, the palynological record from Greenland shows no major diversity or assemblage changes, or conclusive evidence for an extinction event (Raunsgaard Pedersen and Lund, 1980; Koppelhus, 1997).

Palynological records across the T-J boundary are controversially discussed because of the paucity of sections with a sufficient time resolution and/or well established stratigraphic framework. Furthermore, many records are qualitative (i.e. absence/presence data). There- fore, it is important to carry out detailed quantitative studies. The presence of both well- preserved ammonites and palynomorphs in key T-J boundary sections from the Northern Calcareous Alps (Austria) and southern UK allow for an integration of terrestrial microfloral events in a marine biostratigraphic framework. The palynological records presented in this thesis are used to 1) obtain a firm palynostratigraphic framework for the T-J boundary interval, 2) make a reconstruction of past changes in vegetation and climate, and 3) understand the magnitude and nature of the floral turnover as evidenced by palynology.

The Kuhjoch section in the Karwendel syncline (Northern Calcareous Alps, Austria) has recently been approved as the Global boundary Stratotype Section and Point (GSSP) for the base of the Jurassic period. The first occurrence (FO) ofPsiloceras spelae tirolicum has been chosen as the primary boundary marker (Von Hillebrandt et al., 2007; Von Hillebrandt and

11 SYNOPSIS

Krystyn, 2009). Chapter 1 presents a detailed quantitative palynological and carbon isotope study across the T-J boundary from the nearby (5 km) Hochalplgraben section, with first data from the Kuhjoch section. Both sections were deposited in the Eiberg Basin in the western Tethys realm. The FO of Cerebropollenites thiergartii is useful as a palynological marker to identify the approximate position of the base of the Jurassic as it is the only palynomorph which has its FO close to Psiloceras spelae tirolicum. The end-Triassic initial negative CIE is demonstrated at the transition from the Kössen Formation to the Kendlbach Formation. The palynological and carbon isotope records from different sections within the Eiberg Basin (i.e., Hochalplgraben, Kuhjoch and Tiefengraben) correlate very well and the established palynostratigraphic scheme allows for very detailed local and regional correlations.

The focus of Chapter 2 is the nature and cause of the end-Triassic black shale deposition in the Eiberg Basin. This co-occurs with the initial CIE, a prasinophyte mass occurrence, and major vegetation changes. A dramatically increased influx of conifer (Cheirolepidiaceae) pollen and increased total organic carbon values are succeeded by an acme of green algae (Cymatiosphaera). A model is presented in which increased terrestrial organic matter influx is related to enhanced seasonality and increased erosion of the hinterland. Reduced salinity of the surface waters led to the mass occurrence of green algae. Stratification of the water column may have caused anoxic bottom water conditions and black shale deposition during the initial CIE at the base of the Kendlbach Formation.

The vegetation changes during the end-Triassic initial CIE are discussed in detail in Chapter 3. High resolution palynological data from Hochalplgraben and Kuhjoch have been used to infer relative changes in temperature and humidity by using multivariate statistical analyses. For the first time, relative climate changes are directly linked to the CIE. The sporomorph record indicates that a conifer dominated sclerophyllous hardwood vegetation is replaced by a mixed spore plant (e.g., ferns, fern allies, mosses and liverworts) and gymnosperm vegetation. Abrupt warming coincides with the negative carbon isotope shift. The CIE is also simultaneous with a trend from a relatively dry to a more humid climate. The abrupt

global warming during the initial CIE is probably the result of CO2 outgassing from CAMP volcanism and additional marine methane release which caused a northward shift of the tropical summerwet biome on land adjacent to the western Tethys realm.

Global significance of T-J carbon cycle changes might be disputed due to varying magnitudes of the CIE between localities and the possible influence of hydrocarbon-source changes on the bulk sedimentary carbon isotope composition. Therefore, in Chapter 4 the actual size of the T-J carbon cycle changes is estimated on the basis of compound-specific measurements of the carbon isotope composition of long-carbon-chain n-alkanes derived from epicuticular waxes of land plants. Data from Kuhjoch show a 5-6‰ negative CIE, marked by carbon isotope values that are even 2-3‰ lower than previously assumed. It appeared that previous explana-

12 SYNOPSIS

tions based solely on enhanced volcanism no longer provide a feasible mechanism. The magnitude and rate of change in carbon isotope values implies that up to ~6900-8200 Gt 13C depleted carbon must have been transferred from the methane-hydrate reservoir to the atmosphere and oceans. The palynological inferred climate changes are discussed with respect to changes in the carbon cycle during the initial CIE and it appeared that the end-Triassic carbon-cycle changes coincide with a strong warming event and with enhanced hydrological cycling. Hence, our data unambiguously confirm the causal link between massive methane release, climate change and the extinction event at the T-J boundary interval.

Another T-J boundary key section is the St. Audrie’s Bay section in Somerset, southwest UK. Chapter 5 presents a high resolution quantitative palynological study from this section, which revealed a palynofloral transition interval with four pronounced spore peaks during the end-Triassic. Regular cyclic increases in palynomorph concentrations can be linked with periods of increased runoff, and correspond to the orbital eccentricity cycle. The spore peaks can be related to precession-induced variations in monsoon strength. There is no compelling evidence of a global end-Triassic spore spike which, by analogy with the Cretaceous-Tertiary boundary fern spike, could be related to a catastrophic mass extinction event. Climate change is a more plausible mechanism for explaining the increased amount of spores.

Global warming induced by massive CO2 input during deposition of the CAMP is one of the major climate changes that has been suggested for the T-J boundary interval. Fossil leaves are indicators of palaeoatmospheric CO2 levels because of the negative relationship between atmospheric CO2 and stomatal development during growth. In Chapter 6 end-Triassic fluctuations in CO2 concentration are reconstructed by the use of stomatal frequency analysis on a single plant species: the seedfern Lepidopteris ottonis. The stomatal index shows no significant intra- and interpinnule variation which makes this species a suitable proxy for past relative CO2 changes. Records of decreasing stomatal index and density from the bottom to the top of the Wüstenwelsberg section (Bavaria, Germany) indicate rising CO2 levels during the T-J transition. Additionally, stomatal frequency data of fossil ginkgoalean leaves (Gink- goites taeniatus) suggest a palaeoatmospheric CO2 value of 2750 ppmv for the latest Triassic.

The high-resolution T-J boundary palynological datasets from Hochalplgraben and St. Audrie’s Bay (presented in Chapters 1 and 5, respectively) were used to reconstruct changes in vegetation, diversity and climate in Chapter 7. Hochalplgraben shows a change from gymnosperms to spore producing plants and an increased diversification of spore types during the latest Rhaetian. Multivariate statistical analyses revealed a trend to wetter conditions across the T-J boundary. By contrast, in St. Audrie’s Bay a mixed gymnosperm forest is replaced by a monotonous vegetation consisting mainly of Cheirolepidiaceae. This corresponds with a diversity decrease and a change to a warmer and more arid climate. Neither of the sections demonstrates a distinctive floral mass extinction. A compilation of T-J boundary

13 SYNOPSIS

sections across the world demonstrates the presence of Cheirolepidiaceae dominated forests in the Pangaean interior and an increase in spore producing plants near the Tethys Ocean. This floral differentiation reflected in the T-J palynological record is the indirect result of CAMP volcanism. The increase in greenhouse gases caused a warmer climate and an enhanced thermal contrast between the continent and the seas. Consequently, the monsoon system got stronger and induced a drier continental interior and more intensive rainfall near the margins of the Tethys Ocean.

N.B. The chapters of this thesis are or will be published as separate papers in scientific journals. Consequently, some repetition of statements could not be avoided. The datasets presented in this thesis are available upon request.

CHAPTER 1

A detailed palynological study of the Triassic-Jurassic transition in key sections of the Eiberg Basin (Northern Calcareous Alps, Austria)

The Triassic-Jurassic transition is characterized by a major extinction in the marine realm but evidence for floral turnover is ambiguous. Here we present the results of a detailed palynological and carbon isotope 13 (δ Corg) study across the Triassic-Jurassic boundary from the Hochalpl- graben section, with first data from the Kuhjoch section. Both sections are located in the Eiberg Basin (Northern Calcareous Alps, Austria) and they contain well-preserved palynomorphs and ammonites which allow an integration of terrestrial microfloral events in a marine biostratigraphic framework. Five palynomorph assemblages are recognized in 13 the Hochalplgraben section. The initial δ Corg shift occurs at the base of the Tiefengraben Member, the lower part of the Kendlbach Formation, and coincides with an acme of prasinophytes, mainly Cymatiosphaera polypartita. Typical Late Triassic pollen taxa (e.g., Lunatisporites rhaeticus, Rhaetipollis germanicus and Ovalipollis pseudoalatus) disappear at the top of the Schattwald beds (Tiefengraben Member). The first occurrence of the ammonite Psiloceras spelae n. ssp., which is proposed as a marker for the base of the Jurassic System, occurs in the Trachysporites- Heliosporites palynomorph assemblage zone. The base of this zone is marked by the first occurrence ofCerebropollenites thiergartii. Our results 13 show that palynological and δ Corg records from different sections within the Eiberg Basin correlate well and that the established palynostratigraphic scheme allows for very detailed local and regional correlations (e.g., with Danish, German and English basins).

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1. Introduction

The change in (palyno)floras during the end-Triassic period is controversial as some studies report abrupt changes and/or extinction while others report gradual transitions. A major extinction of 60% of terrestrial palynomorph taxa followed by a sharp spore spike at the Triassic-Jurassic (T-J) boundary is claimed in the Newark Basin, USA (Fowell and Olsen, 1993; Fowell et al., 1994; Olsen et al., 2002 a). This spore spike is followed by a Classopollis- dominated palynoflora which is used by Cornet (1977) and Fowell et al. (1994) to mark the base of the Jurassic in the Newark Basin. In the Sverdrup Basin in Arctic Canada, palynomorph assemblages show a significant and relatively abrupt change at the T-J transition from common Norian-Rhaetian palynomorphs to an influx of Jurassic bisaccates and the first appearance of typical Early Jurassic forms such asCerebropollenites thiergartii (Embry and Suneby, 1994). A recent macroecological study showed that Triassic forests in Greenland with high diversity communities were replaced by lower diversity forests and that there was a gradual extinction prior to the T-J boundary (McElwain and Punyasena, 2007; McElwain et al., 2007). In contrast to the palynological record from the Newark Basin, palynological studies from Europe show gradual changes in assemblages at the T-J transition (e.g., War- rington, 1974; Morbey, 1975; Lund, 1977; Schuurman, 1979; Achilles, 1981; Hounslow et al, 13 2004; Kürschner et al., 2007). Two distinctive negative δ Corg excursions, known as the ‘initial’ and the ‘main’, have been recognized from various T-J boundary sections within and outside the Tethys realm (e.g., Pálfy et al., 2001; Hesselbo et al., 2002; Guex et al., 2004; Kürschner et al., 2007; Ward et al., 2007; Ruhl et al., 2009). The Eiberg Basin, located in the Northern Calcareous Alps in Austria, is a key area to study the end-Triassic biotic crisis. This basin contains a relatively complete marine T-J boundary succession because of continuous subsidence during the Late Rhaetian (Krystyn et al., 2005). The sediments from the Eiberg Basin are rich in both marine and terrestrial palynomorphs, but previous palynological records across the T-J boundary are mainly qualitative and often of low temporal resolution (Morbey, 1975; Schuurman, 1979). A high-resolution quantitative palynological record was obtained from the Tiefengraben section in the eastern part of the Eiberg Basin (Kürschner et al., 2007), but a constrained biostratigraphic age assessment of the T-J boundary was limited because of the absence of earliest Jurassic ammonites. However, at the western end of the Eiberg Basin (the Karwendel Syncline) several sections were found with earliest Jurassic ammonites (Von Hillebrandt et al., 2007; Von Hillebrandt and Krystyn, 2009). One of these sections, Kuhjoch, has recently been proposed as a Global boundary Stratotype Section and Point (GSSP) for the base of the Jurassic (von Hillebrandt et al., 2007). The first occurrence (FO) of Psiloceras spelae n. ssp. has been proposed as the primary boundary marker. The GSSP proposal is currently in the ratification process. The Hochalplgraben section is positioned 5 km west of Kuhjoch and it is an important new section as it contains well-preserved ammonites (e.g., Psiloceras spelae n. ssp., P. cf. pacificum) and palynomorphs. This allows the

17 CHAPTER 1

integration of terrestrial microfloral events in a marine biostratigraphic framework. The aims of the present study are: (1) to obtain a detailed quantitative palynological record from Hochalplgraben with first data from Kuhjoch to document detailed palynological changes across the T-J boundary; (2) to assess regional palynofloral patterns by comparing our palynological record with the palynomorph assemblage zones of the Tiefengraben section (Kürschner et al, 2007); (3) to determine whether there are palynological events that coincide with the FO of the ammonite which is proposed as a marker for the base of the Jurassic; and (4) to understand the magnitude and nature of the floral turnover as evidenced by palynology.

2. (Palaeo) geographical setting

The Northern Calcareous Alps are one of the rare regions where the marine Triassic-Jurassic sedimentary record is relatively continuous in many sections. The Hochalplgraben and Kuhjoch sections are located in Northern Tyrol (Fig. 1a) and belong to the southern flank of the western Karwendel Syncline. The Hochalplgraben section (Fig. 1a) is located at 47°28’20’’N/11°24’42’’E, about 25 km north-north-east of Innsbruck and 4 km west-northwest of the village of Hinterriss (1:50.000 scale topographic map of Austria, sheet 118 – Inns- bruck). The coordinates of the nearby candidate GSSP site at Kuhjoch are 47°29’02’’N/11°31’50’’E. Detailed location maps of the sections are given in Von Hillebrandt et al. (2007). The fully marine sedimentary succession exposed in the Hochalplgraben and Kuhjoch sections in the Eiberg Basin was deposited on the western margin of the Tethys Ocean during the T-J transition. The Eiberg Basin is an intraplatform depression that can be traced over 200 km from the Salzkammergut (Upper Austria) in the east, to the Lahnenwiesgraben valley (northwest of Garmisch-Partenkirchen, Germany) in the west (Fig. 1a). To the south, the basin was bordered by the Dachstein lagoon with locally fringing reefs and to the north by the Oberrhaet limestone lagoon (Fig. 1b) (Krystyn et al., 2005, von Hillebrandt et al., 2007). The sections are more distally positioned within the basin than the previously investigated Tiefen- graben and Kendlbachgraben sections (e.g., Morbey, 1975; Golebiowski and Braunstein, 1988; Kürschner et al., 2007).

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8˚ 10˚ 12˚ 14˚ 16˚ 18˚ 50˚ 50˚ km Czech Republic 0 50 100

49˚ Germany 49˚

Vienna Slovakia

48˚ Austria 48˚ Garmisch- Hochalplgraben N Salzburg Partenkirchen Tiefengraben Kuhjoch S Innsbruck Salzkammergut 47˚ area Hungary 47˚ Switzerland

Italy Slovenia Croatia 46˚ 46˚ 8˚ 10˚ 12˚ 14˚ 16˚ 18˚ a

Figure 1 (a) Location of the Hochalplgraben, Kuhjoch and Tiefengraben sections. The shaded area represents the Eiberg Basin (b) Position of the Eiberg Basin in relation to late Norian and Rhaetian facies in a north-south cross-section of the Northern Calcareous Alps

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3. Biostratigraphy

The recovery of Psiloceras spelae n. ssp. (Von Hillebrandt et al., 2007, Von Hillebrandt and Krystyn, 2009) in the Hochalplgraben and Kuhjoch sections is important as its FO significantly narrows the interval between the LO of the latest Triassic and the FO of the earliest Jurassic guide fossils in the Eiberg Basin. P. spelae n. ssp. is closely related to P. spelae known from Nevada (USA). In a voting of the Triassic-Jurassic Boundary Working Group in 2008 P. spelae has been selected to define the first Jurassic ammonite biohorizon, and allows for an intercontinental Tethys-Panthalassa correlation of the future T-J boundary. The slight differences between the subspecies may be due to biogeographic separation (Von Hillebrandt and Krystyn, 2009). 13 Kürschner et al. (2007) presented detailed palynostratigraphy and δ Corg stratigraphy of the Tiefengraben section, located in a more marginal setting in the eastern part of the Eiberg Basin. Five palynomorph assemblage zones have been described and the main characteristics of these zones are summarised below. Rhaetipollis - Limbosporites zone (RL zone): This zone is characterized by a low diversity assemblage with high numbers of Classopollis meyeriana and Classopollis torosus accompanied by Ovalipollis pseudoalatus. Cingulizonates rhaeticus, Limbosporites lundbladii and Rhaetipollis germanicus are present. Rhaetipollis - Porcellispora zone (RPo zone): At the base of this zone spore diversity and abundance, particularly those of Porcellispora longdonensis and Calamospora tener show an acme. There is a brief decline in Classopollis and an increase in Vitreisporites in the lower part of the zone. A distinct increase in spore numbers (e.g., C. tener, Deltoidospora and Polypodiisporites polymicroforatus) denotes the top of the zone. Rhaetipollis germanicus disappears in the upper part of this zone. Trachysporites - Porcellispora zone (TPo zone): This zone is characterised by a continuous decline in Classopollis, and Classopollis torosus even temporarily disappears. Spore assemblages show a decrease in C. tener, Deltoidospora and P. polymicroforatus, accompanied by an increase of Carnisporites, Concavisporites, P. longdonensis and Trachysporites fuscus. Trachysporites - Heliosporites zone (TH zone): Classopollis torosus reappears, while numbers of Classopollis meyeriana remain stable. There is an increase upsection of Porcellispora longdonensis, which is followed by an acme of Heliosporites reissingeri. In the upper part of the zone, H. reissingeri and P. longdonensis decline. Of note is the FO of Cerebropollenites thiergartii at the base of the zone and the peaks in echinate and baculate trilete spores (e.g., Acanthotri- letes varius and Conbaculatisporites) at the top of the zone. Trachysporites - Pinuspollenites zone (TPi zone): Main characteristic of this zone is that Trachysporites fuscus dominates the spore assemblage. Of note is also the common occurrence of Pinuspollenites minimus.

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4. Materials and methods 4.1 Palynology

Fifty-one rock samples from the Hochalplgraben section and fifteen additional samples from the Kuhjoch section were selected for palynological analysis. Between 10-15 g of sediment was processed according to standard procedures as described in Kürschner et al. (2007). The slides are stored in the collection of the Section Palaeoecology, Laboratory of Palaeobotany and Palynology, Utrecht University, The Netherlands. The palynomorphs are very well preserved with a palynomorph colour of 1-2 on the thermal alteration scale (TAS) of Batten (2002). Pollen and spore identification was mainly based on Schulz (1967), Morbey (1975), Lund (1977) and Schuurman (1976, 1977, 1979). The identified morphotaxa of spores, pollen and aquatic palynomorphs are listed in Appendix A and photographs of typical selected palynomorphs are shown in Plates I-IV. About 300 terrestrial palynomorphs were counted per sample (see exact palynomorph sums in Figs. 4-7). Relative abundances were calculated and plotted using the Tilia/TiliaGraph and TGView computer programs (Grimm, 1991-2001). Palynomorph assemblage zones were established by constrained cluster analysis using CONISS (Grimm, 1987) within Tilia. A subsequent qualitative analysis, scanning two complete slides per sample, was carried out to check if rare taxa were present which could be important for biostratigraphy. Rare taxa are indicated in Appendix A, the complete absence/ presence dataset is available on request. The qualitative dataset from the Hochalplgraben section has also been used to carry out a diversity analysis with the computer program PAST (Hammer et al., 2001). Data were subjected to the range-through assumption (absences between first and last appearance are treated as presences) and reworked taxa were rejected from the diversity analysis.

13 4.2 δ Corg analysis

13 Seventy samples from Hochalplgraben were analyzed for bulk δ Corg isotope compositions. 0.9 g of sediment was crushed and treated with 15 ml of 1 M HCl to remove the carbonate. The C-isotope ratio was measured on homogenized samples by Elemental Analyzer Continu- ous Flow Isotope Ratio Mass Spectrometry using a Fisons 1500 NCS Elemental Analyzer coupled to a Finnigan Mat Delta Plus mass spectrometer. The analytical precision was 0.026‰ based on duplicate measurements. Results are reported using standard delta notation relative to Vienna PDB.

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Plate I

22 CHAPTER 1

Plate II

23 CHAPTER 1

Plate III

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Plate IV

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Plate I: pollen Plate II: spores

Following the taxon name is the sample number, followed by the Following the taxon name is the sample number, followed by the slide number in brackets, followed by the stage coordinate for a slide number in brackets, followed by the stage coordinate for a Leica DM LB2 microscope (color version on p. 201) Leica DM LB2 microscope (color version on p. 202)

1. Alisporites radialis, Hin18 (1), 30.4/95.4 1. Aratrisporites parvispinosus, Hin12 (2), 43.4/106.7 2. Alisporites diaphanus, Hin16 (1), 41.2/95.0 2. Aratrisporites minimus, HinA(-1) (1), 39.5/95.3 3. Pinuspollenites minimus, Hin18 (2), 44.4/104.5 3. Calamospora tener, Hin16 (1), 28.3/101.1 4. Quadraeculina anellaeformis, Hin16 (1), 32.7/111.9 4. Todisporites sp., HinA(-1) (1), 41.0/103.5 5. Platysaccus sp., Hin18 (1), 37.4/103.8 5. Deltoidospora sp., HinA(-1) (2), 44.6/106.3 6. Ovalipollis pseudoalatus, Hin16 (1), 32.3/104.6 6. Concavisporites sp., Hin18 (2), 29.6/110.8 7. Vesicaspora fuscus, Hin18 (2), 38.5/110.7 7. Polypodiisporites ipsviciensis, Kuhjoch 050926/4 (1), 8. Perinopollenites elatoides, Hin18 (1), 42.2/99.9 33.3/109.1 9. Vitreisporites bjuvensis, HinB4 (1), 38.0/95.0 8. Polypodiisporites polymicroforatus, Hin18 (2), 40.9/111.9 10. Vitreisporites pallidus, Hin18 (1), 36.3/111.4 9. Stereisporites seebergensis, HinA(-1) (1), 36.8/109.5 11. Classopollis meyeriana (tetrad), HinA(-1) (1), 37.8/102.9 10. Annulispora folliculosa, Hin12 (2), 30.1/95.5 12. Araucariacites australis, Hin18 (1), 28.8/101.5 11. Stereisporites australis, Kuhjoch 051025/3 (1), 33.7/98.4 13. Cycadopites sp., Hin18 (1), 44.6/102.6 12. Rogalskaisporites cicatricosus, HinA10 (1), 45.2/94.9 14. Lagenella martini, Kuhjoch 050926/5 (1), 24.6/96.8 13. Thymospora canaliculata, Hin6 (1), 35.4/107.6 15. Rhaetipollis germanicus, Hin18 (2), 36.4/104.6 14. Baculatisporites sp., Hin18 (2), 40.5/111.9 16. Cerebropollenites thiergartii, HinA7 (1), 36.0/110.0 15. Conbaculatisporites sp., Hin18 (2), 39.5/96.4 17. Cerebropollenites thiergartii, Kuhjoch 051025/4 (2), 16. Acanthotriletes varius, HinA(-1) (1), 44.7/105.0 32.9/102.8 17. Lophotriletes verrucosus, Hin18 (2), 33.9/107.0 18. Cerebropollenites thiergartii, Kuhjoch 050926/4 (2), 18. Carnisporites spiniger, Hin18 (1), 38.6/110.2 42.0/103.3 19. Porcellispora longdonensis, Ho 20 (1), 15.5/97.9 19. Cerebropollenites thiergartii, Kuhjoch 051025/4 (1), 20. Carnisporites anteriscus, Hin16 (2), 32.0/98.0 38.0/97.8 21. Carnisporites lecythus, Hin17 (2), 41.3/101.0 22. Carnisporites leviornatus, HinA(-1) (2), 38.0/98.0

26 CHAPTER 1

Plate III: spores Plate IV: aquatic palynomorphs

Following the taxon name is the sample number, followed by the Following the taxon name is the sample number, followed by the slide number in brackets, followed by the stage coordinate for a slide number in brackets, followed by the stage coordinate for a Leica DM LB2 microscope (color version on p. 203) Leica DM LB2 microscope (color version on p. 204)

1. Trachysporites fuscus, Kuhjoch 050926/4 (1), 1. Rhaetogonyaulax rhaetica, Hin18 (1), 39.5/103.0 43.3/106.8 2. Dapcodinium priscum, HinA(-1) (1), 36.7/94.1 2. Cornutisporites seebergensis, HinA10 (1), 42.3/103.0 3. cf. Beaumontella type A, Kuhjoch 051030/2 (1), 29.8/96.9 3. Perinosporites thuringiacus, Hin18 (1), 34.1/108.4 4. Suessia swabiana, HinB4 (2), 31.3/111.7 4. Triancoraesporites ancorae, Hin16 (2), 40.4/99.9 5. Cleistosphaeridium mojsisovicsii, HinA10 (1), 35.0/94.8 5. Triancoraesporites reticulatus, Hin18 (1), 40.2/102.6 6. Beaumontella langii, HinA(-1) (1), 41.3/105.9 6. Camarozonosporites laevigatus, Hin18 (2), 44.6/104.6 7. Valveodinium koessenium, HinA(-1) (1), 34.2/94.1 7. Camarozonosporites rudis, Hin12 (1), 42.1/106.8 8. Veryhachium sp., Hin16 (2), 39.2/105.5 8. Lycopodiacidites rhaeticus, HinA10 (1), 40.8/96.0 9. Micrhystridium sp. (short spines), HinA10 (1), 37.0/103.0 9. Zebrasporites laevigatus, Hin18 (2), 30.0/108.3 10. Micrhystridium sp. (long spines), Hin18 (1), 41.8/99.3 10. Zebrasporites interscriptus, Hin18 (2), 33.5/110.8 11. cf. Leiosphaeridia sp., HinA(-1) (1), 40.0/94.0 11. Densoisporites nejburgii, Hin18 (1), 46.0/96.3 12. Tasmanites sp., Hin16 (2), 34.8/110.0 12. Cingulizonates rhaeticus, Hin18 (2), 44.8/105.8 13. Cymatiosphaera polypartita (fine walls), Ho15 (1), 28.5/98.8 13. Kyrtomisporis laevigatus, Hin18 (2), 43.2/96.4 14. Cymatiosphaera polypartita (coarse walls), Hin18 (1), 14. Kyrtomisporis speciosus, HinA11 (2), 44.0/111.7 41.8/99.3 15. Densosporites fissus,Hin16 (2), 93.4/105.8 15. Pterospermella sp., Hin18 (1), 40.6/103.8 16. Limbosporites lundbladii, Kuhjoch 051025/4 (1), 33.5/104.6 16. Botryococcus sp., Kuhjoch 051025/4 (2), 35.0/110.0 17. Ischyosporites variegatus, HinA15 (1), 37.7/96.4 17. Foraminiferal test lining, Kuhjoch 050926/4 (1), 31.7/105.7 18. Heliosporites reissingeri, HinA10 (2), 45.0/98.3 18. Foraminiferal test lining, HinA(-1) (1), 25.0/94.1 19. Nevesisporites bigranulatus, HinA10 (1), 46.0/97.8 19. Tytthodiscus cf. faveolus, Hin6 (1), 42.5/97.8

27 CHAPTER 1

5. Results and discussion 5.1 Lithostratigraphy and depositional environment

Figure 2 shows the lithology of the Hochalplgraben profile. The lower part consists of limestones of the Kössen Formation, interpreted to have been deposited in an open marine intraplatform basin (Golebiowski and Braunstein, 1988). The Kendlbach Formation overlies the Kössen Formation and is represented in this section by the Tiefengraben Member which formed in a restricted, shallow marine environment (Golebiowski and Braunstein, 1988). It comprises a thin unit of olive-grey marly sediments that is overlain by red marls (Schattwald beds) that are succeeded by a thicker unit of olive-grey sediments with thin silt- and sandstone layers towards the top (Tiefengraben Member, Fig. 2). A small-scale fault occurs in the upper part of the Schattwald beds at 550 cm depth, and is indicated by a 20 cm thick deformed clay horizon (Fig. 2, this study; Fig. 12 in Von Hillebrandt et al., 2007). We estimate from comparison with other sections that maximum 2-3 m of the sedimentary sequence is missing. The overlying Breitenberg Member has not been included in this study; it represents an open marine shelf environment (Golebiowski and Braunstein, 1988). The lithology of the Kuhjoch section (Fig. 3, this study; Von Hillebrandt et al., 2007; Von Hillebrandt and Krystyn, 2009) is very similar to that at Hochalplgraben. In the Kuhjoch section, a 20 cm thick deformed clay horizon occurs at the top of the Schattwald beds (Fig. 3). However, this disturbance is absent from a recently discovered outcrop on the east side of the Kuhjoch hill, ~ 10 m from where the original outcrop is located. A lowering of sea level at the top of the Kössen Formation resulted in many surrounding shallow water areas becoming emergent, and leading to erosion and discontinuous T-J boundary successions. The Hochalpl- graben and Kuhjoch sites were situated more distally within the basin and were therefore less affected by sea-level drop.

13 5.2 δ Corg Hochalplgraben

13 The bulk δ Corg values from the sedimentary organic material show distinct changes throughout the section (Fig. 2). Samples from the Kössen Formation have values of about -28‰. At 13 the top of the Kössen Formation the δ Corg values show an abrupt initial negative shift to -31‰. This is followed by a sharp increase to higher values of around -25‰ in the Schattwald beds. 13 The onset of the main negative δ Corg excursion (to values of about -28‰) is present at the top 13 of the Schattwald beds. Despite the possibility that the δ Corg signal could be (partly) influ- 13 enced by a change in the source of organic matter, there is supporting evidence that the δ Corg data can be used for stratigraphic correlation at least at an infra-basin scale (Ruhl et al., 2009).

28 CHAPTER 1

5.3 Palynology

5.3.1 Terrestrial vs. aquatic palynomorphs

Most of the Hochalplgraben section is dominated by terrestrial palynomorphs (Fig. 2). There is an increase in aquatic palynomorphs towards the top of the Kössen Formation where a sharp peak in aquatic palynomorphs of 90% occurs, that mainly consists of the prasinophyte 13 Cymatiosphaera polypartita. This peak coincides with the initial negative δ Corg excursion. The aquatic palynomorphs abundance is reduced to 10% at the base of the Schattwald beds. From the top of the Schattwald beds to 1050 cm from the base of the section, the aquatic palynomorphs make up around 60% of the total palynomorph assemblage. The change in the terrestrial:aquatic palynomorph ratio in the Kuhjoch section is similar to that of the Hochalpl-graben section (Fig. 3).

5.3.2 Terrestrial palynology

5.3.2.1 Hochalplgraben

Significant changes occur in the terrestrial palynomorph assemblages through the T-J boundary interval (Fig. 4, see foldout in the back of the thesis). Within this section there is a decreasing trend in the pollen:spore ratio (Fig. 2). In the Kössen Formation, the terrestrial palynomorph fraction consists of about 90% pollen. At the base of the Kendlbach Formation spore abundance increases to values of about 60%. At 550 cm from the base of the record, 13 coinciding with the onset of the main negative δ Corg excursion, a peak abundance of 80% of pollen is present. From 700 cm depth to the top of the section spores are common, with consistent values of around 90%. Based on cluster analysis, four different assemblages (H1-H4) were distinguished. The uppermost assemblage (H4) was subdivided into H4a and H4b, where assemblage H4b is distinguished from H4a by the increase of Pinuspollenites minimus from H4a to H4b and the decrease of Acanthotriletes varius and Conbaculatisporites spp. in H4b. Five palynomorph assemblages can be recognized:

Assemblage H1. This is a low diversity pollen assemblage dominated by Classopollis meyeriana. Other abundant pollen taxa are Alisporites, Ovalipollis pseudoalatus, Rhaetipollis germanicus, Tsugaepollenites pseudomassulae and Vitreisporites. Spores are almost absent. Two samples in the upper part of the assemblage contain mainly Classopollis. In the lower one, Classopollis torosus shows an acme while in the upper sample C. meyeriana dominates. Addition- ally, the abundance of Baculatisporites, Conbaculatisporites, Deltoidospora, Polypodiisporites polymicroforatus and Ricciisporites tuberculatus increases towards the top of this assemblage. This assemblage correlates with the RL zone from the Tiefengraben section.

29 CHAPTER 1

Hochalplgraben

Diversity Depth (cm) Terrestrial palynomorphsAquaticPollen palynomorphs Spores

1800

1700

1600

1500

1400 b 1300

1200 Jurassic 1100

1000

900 Kendlbach Formation Tiefengraben Member Tiefengraben 800

700 a

600

500

400

Schattwald beds 300 Triassic 200 No data 100 Eiberg Mb Kössen Fm 0 -32 -30 -28 -26 -24 20 40 60 80 100 20 40 60 80 100 0 20 40 60 80 100 Fault δ 13Corg [‰] Percentage [%] Percentage [%] # pollen and spore taxa

Grey marl and siltstone Red marl and siltstone Limestone

13 Figure 2: Lithology, δ Corg composition, terrestrial (pollen and spores):aquatic (marine and freshwater) palynomorph ratio, pollen: spore ratio, and terrestrial palynomorph diversity through the Triassic-Jurassic transition in the Hochalplgraben section. Ammonite occurrences: a) Psiloceras spelae n. ssp. b) Psiloceras cf. pacificum. Note that levels are relative to the base of the measured section

30 CHAPTER 1

Kuhjoch

s en re ll o Terrestrial palynomorphsAquatic palynomorphsPo Sp Depth (cm)

1900

1800

1700

1600

1500

1400

1300

1200 b Jurassic 1100

1000

900

800 Kendlbach Formation Tiefengraben Member Tiefengraben 700

600 a 500

400

300

200

100 Schattwald

Triassic 0

-100

-200 No data

Eiberg Mb -300 Kössen Fm

-400

20 40 60 80 100 20 40 60 80 100 Grey marl and siltstone Percentage [%] Percentage [%]

Red marl and siltstone Limestone Fault

Figure 3: Lithology, terrestrial (pollen and spores):aquatic (marine and freshwater) palynomorph ratio, and pollen: spore ratio through the Triassic-Jurassic transition in the Kuhjoch section. Ammonite occurrences: a) Psiloceras spelae n. ssp. b) Psiloceras cf. pacificum. Note that levels are relative to the base of the Kendlbach Formation

31 CHAPTER 1

Assemblage H2. This assemblage is characterized by a decrease in Classopollis meyeriana to values between 20% and 30%. Vitreisporites decreases upsection from values of about 20% to 10%. Ovalipollis pseudoalatus is still present, although in lower abundance than in the previous assemblage. In this assemblage spores are more diverse (>20 taxa), very well preserved and much more abundant than in the preceding assemblage H1. Dominant spores are Concavi- sporites, Deltoidospora, Polypodiisporites polymicroforatus and Ricciisporites tuberculatus. Some of the spore taxa which appear in H2 are Acanthotriletes varius, Camarozonosporites rudis, Cingulizonates rhaeticus, Densosporites fissus, Kyrtomisporis laevigatus, Limbosporites lundbladii, Polypodiisporites ipsviciensis and Triancoraesporites ancorae. Given the marked decline in the abundances of Classopollis, the increase of Vitreisporites and the increase of spore taxa, this assemblage correlates with the RPo zone from the Tiefengraben section.

Assemblage H3. The amount of Classopollis is generally higher than in the previous assemblage, although it shows an upward decreasing trend. There is a virtual absence of pollen taxa that were common in H1 and H2 (e.g., Ovalipollis pseudoalatus, Rheatipollis germanicus, Tsugaepollenites pseudomassulae, Vitreisporites). Additionally, there is a change in spore taxa composition, the most prominent being the upsection increase of Heliosporites reissingeri (10-15%) and Trachysporites fuscus (40%). Furthermore, Ricciisporites tuberculatus has a very low abundance and Polypodiisporites polymicroforatus decreases from about 10% to 1-3%. Concavisporites and Deltoidospora are still present, but several of the taxa that first appeared in Assemblage H2 (e.g., Camarozonosporites rudis, Densosporites fissus, Kyrtomisporis laevigatus and Triancoraesporites ancorae) show a decline in assemblage H3, or are absent. There is a peak abundance of Classopollis meyeriana (90%) immediately above the Schattwald beds. However, a hiatus may have influenced its distribution. Of stratigraphic importance are the FOs of Cerebropollenites thiergartii and Ischyosporites variegatus. Based on the strong increase of H. reissingeri and T. fuscus this assemblage correlates with the lower part of the TH zone from the Tiefengraben section.

Assemblage H4a. This assemblage is mainly characterised by very low abundances of Classopollis species (<10%) and high abundances of Deltoidospora, Ricciisporites tuberculatus and Trachysporites fuscus. It differs from the previous assemblage in that it displays a decline of Heliosporites reissingeri and a marked increase in R. tuberculatus. A remarkable feature is the sudden increase of Acanthotriletes varius and Conbaculatisporites in the sample at 1250 cm. These features are present in the upper part of the TH zone from the Tiefengraben section.

Assemblage H4b. This assemblage shows an increase of Pinuspollenites minimus. The most abundant taxa are Deltoidospora, Ricciisporites tuberculatus and Trachysporites fuscus. Further- more, Acanthotriletes varius and Conbaculatisporites decrease upsection. This assemblage correlates with the TPi zone from the Tiefengraben section.

32 CHAPTER 1

5.3.2.2 Kuhjoch

Overall, the pollen:spore ratio from Kuhjoch (Fig. 3) is very similar to that of Hochalplgraben (Fig. 2). In contrast to the results from the Hochalplgraben section, the Classopollis meyeriana peak registered at the top of the Schattwald beds has not been recognized, likely because of the lower sampling resolution. Four different assemblages (K1-K4) could be distinguished in the Kuhjoch section based on cluster analysis (Fig. 5, see foldout in the back of the thesis). While the two lower assemblages from Kuhjoch correspond with the two lowermost ones from Hochalplgraben (K1=H1 and K2=H2), the upper assemblages are different (K3=H3+H4a, K4=H4b) (Fig. 9). Note that the boundaries between the assemblages are tentative and may still be revised after a future high-resolution study.

Assemblage K1. The oldest sample from the lowermost assemblage shows a high amount of Classopollis meyeriana, Classopollis torosus and Ovalipollis pseudoalatus. The upper sample consists mainly of Classopollis meyeriana. This assemblage correlates with the RL zone from the Tiefengraben section.

Assemblage K2 is characterized by a decrease in Classopollis meyeriana and a marked increase in the number of spore taxa. Furthermore, Vitreisporites bjuvensis is common. Given these features, this assemblage correlates with the RPo zone from the Tiefengraben section.

Assemblage K3. The amount of Classopollis meyeriana shows a decreasing trend with two minor peaks. The FOs of Cerebropollenites thiergartii and Ischyosporites variegatus are of stratigraphic importance. Furthermore, there are changes in spore taxa composition; most prominent are the upsection increase of Trachysporites fuscus and Heliosporites reissingeri and the decrease of Polypodiisporites polymicroforatus. This assemblage correlates with the TH zone from the Tiefengraben section.

Assemblage K4. In the uppermost assemblage, Classopollis meyeriana is rare. However, the amount of indeterminate deformed pollen (probably Classopollis) increases. There is an upsection increase of Pinuspollenites minimus pollen. The most abundant spores are Deltoi- dospora, Ricciisporites tuberculatus and Trachysporites fuscus. Assemblage K4 correlates with the TPi zone from the Tiefengraben section.

5.3.2.3 Palynofloral diversity

It is remarkable that almost all of the spore taxa present in the Hochalplgraben section appear for the first time above the Kössen Fm (Fig. 4). The Schattwald beds show the maximum diversity (Fig. 2) which is in conflict with their red colour, as this would suggest a highly oxidized depositional environment not favourable for palynomorph preservation. Schuurman

33 CHAPTER 1

(1979) mentioned that the assemblages from the Kössen Formation contain relatively few taxa, while the assemblages from the Schattwald beds contain a rich and diverse palynoflora. He suggested a gradual nature of the progressive trend in species diversification. The present study shows that there is indeed a gradual diversification of spores after the Kössen Fm (Fig. 2 and 4) and that maximum diversity occurs in an interval where the marine fauna is impover- ished. However, most of these spores are known to have much longer stratigraphic ranges than recorded in these sections. Previous records of Acanthotriletes varius (Morbey, 1975; Achilles, 1981), Camarozonosporites rudis (Schulz, 1967; Visscher and Brugman, 1981; Dybkjær, 1991), Cingulizonates rhaeticus (Morbey, 1975; Kürschner et al., 2007), Limbosporites lundbladii (Morbey, 1975; Achilles, 1981; Kürschner et al., 2007;), Triancoraesporites ancorae (Dybkjær, 1991), Zebrasporites laevigatus (Morbey, 1975; Achilles, 1981) and Zebrasporites interscriptus (Achilles, 1981; Dybkjær, 1991) show that these taxa were already present earlier in the Rhaetian. The virtual lack of spores in the Kössen Formation is probably caused by non- favourable environmental conditions limiting spore-producing plant growth during deposition. Nevertheless, the present study shows an almost fourfold increase in pollen and spore diversity across the boundary (Fig. 2). This is in sharp contrast to floral diversity data from other regions. Although the position of the T-J boundary in the Newark Basin is still under discussion there is a clear palynological turnover and a strong palynomorph diversity decrease of about 60% in the Passaic Formation below the Jacksonwald Basalt (Fowell and Olsen, 1993). Also a recent quantitative macrobotanical study from Greenland shows a decrease in standing species richness by about 85% (McElwain et al., 2007). The contrasting trends in floral diversity likely reflect regional differences in environmental stress, climatic changes and different palaeogeographic positions during the end-Triassic biotic crisis.

5.3.3 Aquatic palynomorphs

A regressive-transgressive couplet from the Kössen Formation to the upper part of the Tiefengraben Member can be inferred from lithology. This transgression has been described from the UK and the Tethys realm by Hallam (1990a). Batten and Koppelhus (1996) report the beginning of a widespread late Rhaetian marine transgression in Europe that was maintained throughout much of the Jurassic Period. In the Northern Calcareous Alps, evidence for a major end-Triassic regression include signs of karstification (Satterley et al., 1994), a sedimentary break on the carbonate platforms, and a lack of deep water fossils such as ammonites (Krystyn et al., 2005). Red siltstone and mudstone horizons with mud cracks above the Kössen Formation in the Lorüns section (Voralberg, Austria) confirm an end-Triassic sea-level drop (McRoberts et al., 1997). Lindström and Erlström (2006) showed that the end-Triassic regression was preceded by a latest Rhaetian maximum flooding event and succeeded by an earliest Jurassic transgression. In the Tiefengraben section the regression- transgression pattern is indicated by the change in dinoflagellate cyst species fromRhaeto- gonyaulax rhaetica to Dapcodinium priscum at the transition from the Kössen Formation to

34 CHAPTER 1

the Kendlbach Formation (Kürschner et al., 2007). This same pattern can be observed in the Hochalplgraben and Kuhjoch sections although less pronounced (Fig. 6 and 7). This change could be interpreted as a regression as R. rhaetica was adapted to open marine conditions at greater water depth while D. priscum was an opportunistic, euryhaline species that occupied various ecological niches (Courtinat and Piriou, 2002; Holstein, 2004). R. rhaetica is mainly found in low energy sediments (mostly bay deposits) while D. priscum occurs in high to low energy deposits that reflect near shore and restricted marine environments (Courtinat and Piriou, 2002). Immediately above the Kössen Formation there is a conspicuous peak in aquatic palynomorphs (Fig. 2), mainly caused by the increase of the prasinophyte Cymatiosphaera polypartita (Fig. 6 and 7, see foldout in the back of the thesis). Fossil prasinophytes were suggested to be “disaster species” that survived the widespread extinctions in the Middle Palaeozoic and that they are most abundant in the absence of other phytoplankton (Tappan, 1980). Prasinophytes are highly successful in adapting to changes in their environment and this ability permits them to proliferate under conditions hazardous to other organisms (Guy-Ohlson, 1996). Prasinophyte acmes may have been favoured by lower temperatures and reduced salinities in the surface water layer (Prauss and Riegel, 1989). Thus, the peak abundance of Cymatiosphaera polypartita can be interpreted as indicating ecological disturbances rather than a transgression. The prasinophyte abundance peak was also recorded in the Tiefengraben section (Kürschner et al., 2007) where there are actually two separate peaks that coincide with the absence of foraminiferal linings. In both the Hochalplgraben and Kuhjoch sections, the prasinophyte acme is contemporaneous with the absence of foraminifera (Fig. 6 and 7). Prasinophyte abundance appears to be sensitive to reduced grazing which could explain their abundance when foraminifera are absent (Metaxas and Scheibling, 1996). The rarity of foraminiferal linings in the Schattwald beds coincides with a decrease of foraminifera species (Von Hillebrandt et al., 2007). The decrease in foraminifera is likely caused by a calcification crisis resulting from acidification of the atmosphere and oceans by volcanism and methane release (Kürschner et al., 2007). Recently, Van de Schootbrugge et al. (2007a) suggested that green algal phytoplankton acmes may be symptomatic of elevated carbon dioxide levels in the atmosphere and oceans. Within the Schattwald beds from the Hochalplgraben section there is an increase (to 25% peak abundance) in Botryococcus (Fig. 6), an indicator of freshwater environment or influence (Guy-Ohlson, 1992). Therefore, the Schattwald beds were likely deposited in a shallow marine environment with possibly reduced salinity caused by a climate-change driven increased input of freshwater. A high percentage of quartz grains in the Schattwald beds from the nearby Lorüns section suggests a nearby terrestrial source (McRoberts et al., 1997). This may also explain the absence of ammonites. Above the Schattwald beds, the relative abundance of aquatic palynomorphs in the Hochalplgraben and Kuhjoch sections increases to around 60% (Fig. 2 and 3) which is due to the high amount of cf. Leiosphaeridia (Fig. 6 and 7). The occurrence of the first Tethyan Jurassic ammonitePsiloceras spelae n. ssp. lies within

35 CHAPTER 1

this interval. This implies that after the deposition of the Schattwald beds there were more open marine conditions with normal salinity. What could have caused the sea level changes during the T-J transition? According to Hallam and Wignall (1999) the regression-transgression couplet is likely to be global in extent with a rate of sea-level change of at least 1 cm in 0.2 ka (Hallam, 1997). As the Late Triassic was a non-glacial interval, glacio-eustasy can not be expected (Satterley, 1996). Hallam and Wignall (1999) proposed that the rate of sea level change was too rapid for ocean ridge volume change. According to Krystyn et al. (2005), high amplitude of sea level changes during a greenhouse period can only be explained by short-term tectonic uplift and slow rebound within a small area. Fluctuations in relative sea level could also influence terrestrial palynomorph assemblages. For the Tiefengraben section it was demonstrated that variations within terrestrial palynomorph assemblages were independent of sea level fluctuations, as both the absolute concentration and the relative abundance of aquatic versus terrestrial palynomorphs remains stable throughout the Tiefengraben Member (Kürschner et al., 2007). As changes in the palynomorph assemblages are consistent throughout the Eiberg Basin (e.g., shallower Tiefengraben in the eastern part, and deeper Hochalplgraben and Kuhjoch in the western part) it can be assumed that variations of terrestrial palynomorph assemblages represent vegetation changes in the hinterland.

5.4 Biostratigraphic correlations

5.4.1 Correlation within the Eiberg Basin

13 There is a good correlation between palynomorph assemblages and the δ Corg records from Hochalplgraben and Tiefengraben, and several bio-events can be recognized in both sections (Fig. 8). These include the mass occurrence of Cymatiosphaera polypartita, the FO of Cerebro- pollenites thiergartii, the main increase of Heliosporites reissingeri, the peak abundance of the echinate and baculate spores Acanthotriletes varius and Conbaculatisporites, and the occurrence of Pinuspollenites minimus. The palynomorph assemblages from the Hochalplgraben and Kuhjoch sections can directly be correlated with the palynomorph assemblage zones from the Tiefengraben section (Figs. 4, 5, 8 and 9). There are minor differences: the TH zone in Hochalplgraben is divided in two subzones due to an increase in Ricciisporites tuberculatus and a decrease in H. reissingeri (Fig. 4). In the Hochalplgraben section, a 20 cm thick deformed clay unit marks a fault at 550 cm in the lower part of the Tiefengraben Member. Our palynological results indicate that the TPo zone is lacking in the Hochalplgraben section due to this fault. For example, this level is characterized by an abrupt increase of Trachysporites fuscus and H. reissingeri in Hochalplgraben (Fig. 4) while the TPo zone is marked by a gradual increase of these two taxa in Tiefengraben. Moreover, the gradual increase of T. fuscus precedes that of H. reissingeri (Fig. 7 from Kürschner et al., 2007). Both features are not present in the palynologi-

36 CHAPTER 1

cm Hochalplgraben Tiefengraben 2000

cm 3000 4 TPi (H4b) 1500 TPi 3

2500

TH H4a 1000 TH

2 2000 TPo H3

500 RPo RPo (H2) 1500 1

Eiberg RL (H1) 0 -32 -30 -28 -26 -24

δ 13Corg [‰] 1 Choristoceras marshi 1000 2 Psiloceras spelae n. ssp. 3 Psiloceras cf. pacificum RL 4 Psiloceras calliphyllum Cerebropollenites thiergartii 500 Ischyosporites variegatus Heliosporites reissingeri Acanthotriletes varius Conbaculatisporites spp.

Pinuspollenites minimus 0 -34 -32 -30 -28 -26 -24 Cymatiosphaera polypartita δ 13Corg [‰]

Figure 8: Correlation of the Hochalplgraben section with the Tiefengraben section (Kürschner et

13 al., 2007) based on δ Corg stratigraphy, palynomorph assemblage zones and bio-events. RL: Rhaetipollis - Limbosporites zone, RPo: Rhaetipollis - Porcellispora zone, TPo: Trachysporites - Por- cellispora zone, TH: Trachysporites - Heliosporites zone, TPi: Trachysporites - Pinuspollenites zone. The grey lines indicate the missing part of the Tiefengraben zonation scheme within the Hochalplgraben section.

37 CHAPTER 1

13 cal record of Hochalplgraben. This hiatus is confirmed by the sharpδ Corg shift at the transition from the Schattwald beds to the Tiefengraben Member while in the Tiefengraben section this zone is characterized by a gradual negative trend from -26‰ to -28‰ (Fig. 8). 13 The δ Corg data from the Kuhjoch section also show a distinct shift at the same level (Ruhl et al., 2009), which makes it plausible that the TPo zone is also missing in Kuhjoch. Our preliminary palynological results support a missing TPo zone in Kuhjoch (e.g., the abrupt decrease of Vitreisporites directly followed by the FO of Cerebropollenites thiergartii), but a higher resolution of palynological sampling is needed to confirm this. We estimate that the thickness of the missing interval is about 2-3 m based on comparison of the lithology and 13 δ Corg records between the Hochalplgraben and Tiefengraben sections (Fig. 8). Cyclostrati- 13 graphic analysis of δ Corg records suggest that the hiatus encompasses a time interval of max. 50000 yrs (Ruhl, pers. comm.). However, in a more recently discovered outcrop on the east side of the Kuhjoch hill, ~ 10 m from the original outcrop, the transition from the Schattwald beds to the Tiefengraben Member appears to be undisturbed and we think there is a complete section here. A detailed geochemical and palynological study on this outcrop is in progress. Nevertheless, it is clear that a general palynostratigraphic scheme can be established for the entire Eiberg Basin, as the zones are clearly recognized in the Hochalplgraben and Tiefengra- ben sections, and also in the preliminary data from the Kuhjoch section. The zonation scheme for the Eiberg Basin correlates very well with previous zonation schemes for the Alpine realm (Fig. 9) (Morbey, 1975; Schuurman, 1977, 1979). However, the present scheme is more detailed because of the higher sampling resolution. The RL zone shows similarities with the MI miospore subzone as described by Morbey (1975). In both zones some of the main constituents are Classopollis torosus, Classopollis meyeriana and Ovalipollis pseudoalatus. Ricciisporites tuberculatus and Trachysporites fuscus are present while Granuloperculatipollis rudis is virtually absent. The base of the consecutive FG miospore subzone (Morbey, 1975) is defined by the FO ofDensosporites fissus, Retitriletes gracilis and Triancoraesporites ancorae which correlates with the base of the RPo zone. In the recent studies from the Eiberg Basin four zones could be distinguished within this FG miospore subzone. The RL zone also correlates with phase 3 as described by Schuurman (1977, 1979) based on the dominance of Classopollis together with Ovalipollis pseudoalatus and Rhaetipollis germanicus. Furthermore, Densosporites fissus, Heliosporites reissingeri, Polypodiisporites polymicroforatus, Quadraeculina anellaeformis and Ricciisporites tuberculatus are present, which supports the correlation. The main difference is that in the Hochalplgraben section Granuloperculatipollis rudis is absent, while Alisporites, Tsugaepollenites pseudomassulae and Vitreisporites are important constituents of the assemblage (just as in the Tiefengraben section). The RPo zone correlates with phase 4 based on the disappearance of Lunatisporites rhaeticus, Ovalipollis pseudoalatus and Ricciisporites germanicus. In contrast to the present study, Schuurman (1977, 1979) described that in phase 4 H. reissingeri is a dominant constituent. This is not the case either in the Hochalplgraben or in the Tiefengraben section. The TH and TPi zones correspond to phase 5 that was assigned to the Jurassic by Schuurman (1977, 1979). Similarities are the absence of

38 CHAPTER 1

O. pseudoalatus and R. germanicus, the dominance of some smooth trilete spores and the presence of H. reissingeri and Q. anellaeformis. Classopollis is not dominant in the TH and TPi zones in the Hochalplgraben section, which is in contrast to Schuurman’s study. While Schuurman (1977, 1979) stated that many trilete spores (e.g., Cingulizonates rhaeticus, Densosporites fissus, Perinosporites thuringiacus, Triancoraesporites reticulatus and Zebrasporites laevigatus) are absent, these are still present in Hochalplgraben.

Tethyan Northern Calcareous Alps Germany - UK biochronology (Austria) Scandinavia Chrono- Poulsen & stratigraphy Hochalpl. Kuhjoch Kürschner et al. macro micro Schuurman Morbey Lund Orbell Riding 2007 fossils fossils This study 1979 1975 1977 1973 2003 dinocysts terr. palyno dinocysts

4 P. calliphyllum

Neophyllites H4b K4 zone

Pl. gigantea Zone 3 P. cf. pacifi- TPi cum Zone

H4a Phase 5

P. ex gr. P. Pinuspollenites JURASSIC Hettangian tilmanni

K3 zone Trachysporites

2 P. spelae H3 TH n. ssp. Dapcodinium priscum TPo Dapcodinium priscum Heliosporites

zone FG Miospore Subzone Zone latest Rhaetian

H2 K2 Phase 4 Triassic RPo zone Ricciisporites calcareous Polypodiisporites microfossils

M. ultima 1 C. marshi TRIASSIC Zone M. rhaetica H1 K1 RL Zone zone Rhaetipollis Phase 3 Subzone

Rhaetian V. stuerzen M. posth./ Mid - Late Rhaetogonyaulax rhaetica Rhaetogonyaulax rhaetica Rhaetipollis baumi hernsteini MI Miospore Limbosporites

Figure 9: Comparison of the biostratigraphy and palynological zonation schemes from the Northern Calcareous Alps with those from Germany, Scandinavia and Britain. The Tethyan biochronology is based on Von Hillebrandt et al. (2007) and Kürschner et al. (2007). Macrofossil abbreviations: P. calliphyllum = Psiloceras calliphyllum, Pl. gigantea = Plagiostoma gigantea, P. cf. pacificum = Psilocerascf. pacificum, P.ex gr. P. tilmanni = Psiloceras ex gr. Psiloceras tilmanni, P. spelae n. ssp. = Psiloceras spelae n. ssp., C. marshi = Choristoceras marshi, V. stuerzenbaumi = Vandaites stuerzenbaumi. Microfossil abbreviations: M. ultima = Misikella ultima, M. rhaetica = Misikella rhaetica, M. posth./hernsteini = Misikella posthernsteini/hernsteini.

39 CHAPTER 1

5.4.2 Correlation with Germany, Denmark and the UK

Latest Triassic and earliest Jurassic palynological assemblages are well documented from the German and Danish Triassic basins (e.g., Schulz, 1967; Lund, 1977; Achilles, 1981; Dybkjær, 1988; Lund, 2003; Lindström and Erlström, 2006) and from the UK (e.g., Orbell, 1973; Warrington, 1974; Hounslow et al., 2004). Figure 9 shows the palynological correlation of these areas with the Eiberg Basin. The similarity between the palynomorph assemblage zones from the Eiberg Basin and the zones described by Lund (1977) from the North Sea Basin is striking. The RL zone from the Eiberg Basin corresponds to the Rhaetipollis Limbosporites zone sensu Lund, which is defined by the combined occurrence ofRhaetipollis germanicus and Limbosporites lundbladii. The succeeding Ricciisporites Polypodiisporites zone sensu Lund is characterized by the presence of common Polypodiisporites polymicroforatus or of Semiretisporis together with a wide variety of other spore species, but R. germanicus is practically absent (Lund, 1977). Despite the presence of R. germanicus in the RPo zone from the Eiberg Basin, there is a strong correspondence between the RPo zone and the Ricciisporites Polypodiisporites zone sensu Lund. The Pinuspollenites Trachysporites zone sensu Lund correlates with the TH zone and TPi zone, and is characterised by the presence of common Pinuspollenites minimus, Trachysporites fuscus and Heliosporites reissingeri. Orbell (1973) described two palynomorph zones from the British T-J transition. The boundary between the older Rhaetipollis zone and the younger Heliosporites zone is defined by the rapid decline in proportions of Ovalipollis ovalis, Rhaetipollis germanicus and Ricciisporites tuberculatus. In the Heliosporites zone, Heliosporites reissingeri occurs in very high numbers together with increased numbers of Quadraeculina anellaeformis and Camarozonosporites rudis. Both zones can be recognized in the Eiberg Basin: the RL zone correlates with the Rhaetipollis zone, and the RPo zone up to the TPi zone correlates with the Heliosporites zone. The exact position of the boundary defined by Orbell (1973) is questionable as the decline ofOvalipollis pseudoalatus, Rhaetipollis germanicus and R. tuberculatus is usually stepwise (Warrington, 2005). The dinoflagellate cyst zonation of Poulsen and Riding (2003) shows a correlation between Austria and Germany, UK, and Scandinavia. This transition between the Rhaeto- gonyaulax rhaetica zone and the Dapcodinium priscum zone is based on the common/abundant presence of Rhaetogonyaulax rhaetica (Poulsen and Riding, 2003).

5.5 Palynological events at the Triassic-Jurassic boundary

Typical Triassic palynomorph assemblages (e.g., with Lunatisporites rhaeticus, Ovalipollis pseudoalatus and Rhaetipollis germanicus) are still present in the Kössen Formation and Schattwald beds from the Hochalplgraben section. The records of these taxa show that they disappear at the end of the RPo zone (Fig. 4). At 700 cm above the base of the section, a recently found ammonite horizon with Psiloceras spelae n. ssp. (some 1500 cm below the

40 CHAPTER 1

Psiloceras calliphyllum interval) documents the oldest found Jurassic ammonite of the Tethys realm (Fig. 2). 150 cm below this ammonite level, Cerebropollenites thiergartii has its FO (at 553 cm). This pollen taxon has been suggested as a palynological marker for the base of the Jurassic (Fisher and Dunay, 1981; Kürschner et al., 2007). It also appears to have its FO near the base of the Jurassic in the Sverdrup Basin in Canada (Embry and Suneby, 1994), the Alborz mountains in Iran (Achilles et al., 1984), the Germanic Basin (Schulz, 1967) and in the Danish subbasin (Dybkjær, 1988). Another taxon which has previously been recorded from the Jurassic is Ischyosporites variegatus (Schulz, 1967; Achilles et al., 1984), which has its FO in Hochalplgraben at 573 cm. In the earlier palynological pilot study of the Kuhjoch section, the FO of C. thiergartii coincides with the FO of P. spelae n. ssp. (Von Hillebrandt et al., 2007). In our new detailed study, C. thiergartii and I. variegatus are registered below the first Jurassic ammonite (Fig. 5) which is similar to the situation in Hochalplgraben. Although the FO of C. thiergartii is slightly lower than the FO of P. spelae n. ssp. in the Austrian sections, it is still useful as a palynological marker to identify the approximate position of the base of the Jurassic as it is the only palynomorph which has its FO close to the earliest Jurassic ammonite.

5.6 A comment on Kuhjoch as a GSSP candidate

The present study shows that there are palynological events across the T-J boundary in different sections of the Eiberg Basin in the western Tethys, that can be correlated with palynological records outside the Alpine realm. In fact, the studied sections are unique because the terrestrial palynomorph record can be directly correlated to a complete record of the marine Rhaetian to the Hettangian micro- and macro-faunal sequence, particularly with respect to the main biostratigraphic markers, the ammonites (Von Hillebrandt et al, 2007; Von Hillebrandt and Krystyn, 2009). Moreover, the sedimentary sequence is studied 13 for δ Corg stratigraphy and cyclostratigraphy (Ruhl et al., 2009; Ruhl 2010) that enables a precise age assessment beyond biostratigraphic resolution. There are no other T-J boundary sections presently known that allow such study of the transition of Rhaetian to Hettangian fauna and microflora. The British GSSP candidate (St. Audrie’s Bay) does not contain a Late Triassic ammonite fauna while the North American GSSP candidate (New York Canyon) does not contain palynomorphs. Already at the present stage, we are able to recognize small differences in the sedimentary successions between individual sections, such as the minor hiatus in the palynological record at the top of the Rhaetian caused by local faulting in the Hochalplgraben and Kuhjoch sections. The hiatus is also visible as a sharp shift of 3‰ and 13 2.5‰ in the δ Corg record of the Hochalplgraben and Kuhjoch section, respectively (Ruhl et al., 2009). From cyclostratigraphic data we estimate the length of this transitional zone to be of max. 50000 years, as one half of an eccentricity cycle is missing (Ruhl et al., pers. comm.). The missing TPo zone represents the latest Rhaetian pollen assemblage zone which means that the sedimentary record of the proposed base of the Hettangian, as indicated by the FO of

41 CHAPTER 1

Psiloceras spelae n ssp., is undisturbed. Note that in the candidate GSSP section at Ferguson Hill, New York Canyon, a fault is present in uppermost Triassic sediments (Guex et al., 2009). Additionally, newly discovered outcrops are currently under investigation, such as a section at the east side of the Kuhjoch hill, ~10 m from the original outcrop which does not show the disturbance at the transition from the Schattwald beds to the Tiefengraben Member. Detailed geochemical and palynological studies to verify this are in progress.

N.B. After writing of this chapter, Kuhjoch was approved as the GSSP for the base of the Jurassic.

6. Conclusions

Five palynomorph assemblages are recognized in the Hochalplgraben section. The initial 13 δ Corg shift, known from other T-J boundary sections, is demonstrated at the transition from the Kössen Formation to the Tiefengraben Member and it coincides with a mass occurrence of the green alga Cymatiosphaera polypartita. The disappearance of typical Late Triassic pollen taxa as Lunatisporites rhaeticus, Ovalipollis pseudoalatus and Rheatipollis germanicus occurs at the top of the Schattwald beds. Psiloceras spelae n. ssp., proposed as a marker for the base of the Jurassic, has its FO in the lower part of the Tiefengraben Member. This level corresponds with the Trachysporites-Heliosporites zone which is characterized by the FO of Cerebropollenites thiergartii and an increase in Trachysporites fuscus and Heliosporites reissingeri. The FO of C. thiergartii appears to be a biostratigraphically useful marker that enables the correlation between terrestrial and marine realms. Palynomorph assemblages are changing rapidly at the T-J boundary interval and there is an increase in pollen and spore diversity at the transition from the Kössen Formation to the Schattwald beds. The maximum in palynofloral diversity coincides with the extinction interval in the marine realm. The palynomorph assemblages from the Hochalplgraben section correlate with the zones from the Tiefengraben section. Data from the Kuhjoch section shows similar changes in palynomorph assemblages. Therefore, the established palynostratigraphic scheme allows for high-resolution correlations within the Eiberg Basin.

42 CHAPTER 1

Acknowledgements

NB and WMK acknowledge funding from the “High Potential” stimulation program of Utrecht University. LK has been financially supported by the Austrian National Committee for IGCP (Project 458). We thank the Austrian Bundesforste for access to forest roads. A. von Hillebrandt is thanked for his collaboration in the field and for providing the Kuhjoch samples. We thank J. van Tongeren and N. Welters for assistance in the laboratory and A. van Dijk 13 (Int. Geoch. Lab., UU-NITG) and M. Ruhl for performing the δ Corg analyses. A. Lotter, H. Visscher and H. van Konijnenburg-van Cittert are thanked for useful discussions. Construc- tive comments by M. Stephenson, S. Lindström and two anonymous reviewers substantially improved the manuscript.

43 CHAPTER 1

Appendix A:

Alphabetical list of palynomorphs identified in the Hochalplgraben and Kuhjoch sections, * = encountered after qualitative analysis, R = maybe reworked

Pollen Alisporites diaphanus (Pautsch) Lund 1977 Alisporites radialis (Leschik) Lund 1977 Alisporites robustus Nilsson 1958 Araucariacites australis Cookson 1947 cf. Callialasporites dampieri * (Balme) Dev 1961 Cerebropollenites thiergartii Schulz 1967 Chasmatosporites apertus (Rogalska) Nilsson 1958 cf. Chasmatosporites type A * Classopollis meyeriana (Klaus) Venkatachala & Góczán 1964 Classopollis murphyi Cornet & Traverse 1975 Classopollis torosus (Reissinger) Klaus 1960, emend. Cornet & Traverse 1975 Cycadopites Wodehouse 1933 Cycadopites type A * Ephedripites Bolchovitina 1953 ex. Potonié 1958 Ephedripites tortuosus * Maedler 1964 Eucommiidites major * Schulz 1967 Eucommiidites troedssonii Erdtman 1948 Florinites pellucidus R (Wilson & Coe) Wilson 1958 Granuloperculatipollis rudis Venkatachala & Góczán 1964 Lagenella martini (Leschik) Klaus 1960 Lunatisporites rhaeticus (Schulz) Warrington 1974 Ovalipollis pseudoalatus (Thiergart) Schuurman 1976 Perinopollenites elatoides Couper 1958 Pinuspollenites minimus (Couper) Kemp 1970 Platysaccus Naumova 1937 Quadraeculina anellaeformis Maljavkina 1949 Rhaetipollis germanicus Schulz 1967 Striatoabieites aytugii R Visscher 1966 Triadispora Klaus 1964 Tsugaepollenites pseudomassulae (Maedler) Morbey 1975 Vesicaspora fuscus (Pautsch) Morbey 1975 Vitreisporites bjuvensis Nilsson 1958 Vitreisporites pallidus (Reissinger) Nilsson 1958

44 CHAPTER 1

Spores Acanthotriletes varius Nilsson 1958 Annulispora folliculosa (Rogalska) de Jersey 1959 Aratrisporites crassitectatus * Reinhardt 1964 Aratrisporites minimus Schulz 1967 Aratrisporites parvispinosus (Leschik) Playford & Dettmann 1965 Aratrisporites sp. Asseretospora gyrata (Playford & Dettman) Schuurman 1977 Baculatisporites Thomson & Pflug 1953 Calamospora tener (Leschik) Maedler 1964 Camarozonosporites aulosenensis Schulz 1967 Camarozonosporites laevigatus Schulz 1967 Camarozonosporites rudis (Leschik) Klaus 1960 Carnisporites anteriscus Morbey 1975 Carnisporites lecythus Morbey 1975 Carnisporites leviornatus (Levet-Carette) Morbey 1975 Carnisporites Maedler 1964 Carnisporites megaspiniger Morbey 1975 Carnisporites spiniger (Leschik) Morbey 1975 Cingulizonates rhaeticus (Reinhardt) Schulz 1967 Conbaculatisporites Klaus 1960 Concavisporites Pflug 1953 Converrucosisporites luebbenensis Schulz 1967 Cornutisporites rugulatus Schulz 1967 Cornutisporites seebergensis * Schulz 1962 Cosmosporites elegans Nilsson 1958 cf. Cyclotriletes R Deltoidospora Miner 1935 Densoisporites nejburgii (Schulz) Balme 1970 Densosporites fissus (Reinhardt) Schulz 1967 Densosporites irregularis Hacquebard & Barss 1957 Echinitosporites iliacoides R Schulz & Krutzsch 1961 Foveosporites Balme 1957 Foveosporites foveoreticulatus * Doering 1965 Foveosporites multifoveolatus * Doering 1965 cf. Guthoerlisporites magnificus* Bharadwaj 1954 Heliosporites reissingeri (Harris) Muir & van Konijnenburg - van Cittert 1970 Ischyosporites variegatus (Couper) Schulz 1967 Kyrtomisporis laevigatus Maedler 1964 Kyrtomisporis speciosus Maedler 1964b

45 CHAPTER 1

Leptolepidites Couper 1953 emend. Schulz 1967 Limbosporites lundbladii Nilsson 1958 Lophotriletes gibbosus * (Ibrahim) Potonié & Kremp 1954 Lophotriletes verrucosus Schulz 1967 Lycopodiacidites rhaeticus Schulz 1967 Lycopodiacidites rugulatus (Couper) Schulz 1967 Neochomotriletes triangularis (Bolchovitina) Reinhardt 1962 Nevesisporites bigranulatus (Levet-Carette) Morbey 1975 Paraklukisporites foraminis * Maedler 1964 Perinosporites thuringiacus Schulz 1962 Platypetera trilingua (Horst) Schulz 1967 Polycingulatisporites bicollateralis (Rogalska) Morbey 1975 Polypodiisporites ipsviciensis (de Jersey) Playford & Dettman 1965 Polypodiisporites polymicroforatus (Orlowska-Zwolinska) Lund 1977 Porcellispora longdonensis (Clarke) Scheuring 1970 emend. Morbey 1975 Retitriletes (v.d. Hammen ex Pierce) Doering, Krutzsch, Mai & Schulz 1963 Retitriletes gracilis (Nilsson) Doering, Krutzsch, Mai & Schulz 1963 Retitriletes semimuris (Danzé-Corsin & Laveine) Reiser & Williams 1969 Retitriletes subrotundus (Kara-Murza in Koll.) Doering, Krutzsch, Mai & Schulz 1963 Ricciisporites tuberculatus Lundblad 1954 Rogalskaisporites cicatricosus (Rogalska) Danzé-Corsin & Laveine 1963 cf. Saccizonati sp. * Bharadwaj 1957 Selagosporis mesozoicus * Schulz 1967 cf. Sellaspora rugoverrucata * R van der Eem 1983 Semiretisporis gothae Reinhardt 1962 Spore indet A * Spore indet B Stereisporites Pflug 1953 Stereisporites australis (Cookson) Schulz 1970 Stereisporites infrapunctus Schulz 1970 Stereisporites punctatus * Schulz 1970 Stereisporites seebergensis Schulz 1966 Stereisporites sinuosa * Schulz 1970 Thymospora canaliculata Schuurman 1977 Tigrisporites microrugulatus Schulz 1967 Todisporites Couper 1958 Trachysporites fuscus Nilsson 1958 Trachysporites sp. A Triancoraesporites ancorae (Reinhardt) Schulz 1967 Triancoraesporites reticulatus Schulz 1962

46 CHAPTER 1

Uvaesporites argentaeformis (Bolchovitina) Schulz 1967 Vallasporites ignacii Leschik 1956 Verrucosisporites cheneyi Cornet & Traverse 1975 Verrucosisporites sp. * Zebrasporites interscriptus (Thiergart) Klaus 1960 Zebrasporites laevigatus (Schulz) Schulz 1967

Dinoflagellate cysts Beaumontella langii (Wall) Below 1987 cf. Beaumontella type A Cleistosphaeridium mojsisovicsii Morbey 1975 Dapcodinium priscum Evitt 1961 emend. Below 1987 Rhaetogonyaulax rhaetica (Sarjeant) Loeblich & Loeblich 1968 Suessia swabiana Morbey 1975 emend. Below 1987 Valveodinium koessenium (Morbey) Below 1987

Acritarchs Micrhystridium Deflandre 1937 emend. Sarjeant 1967 Veryhachium Deunff 1958

Prasinophytes Cymatiosphaera polypartita Morbey 1975 cf. Leiosphaeridia Eisenack 1958 Pterospermella Eisenack 1972 Tasmanites Newton 1875 Tytthodiscus cf. faveolus Morbey 1975 Schizocystia cf. rara * Playford & Dettmann 1965

Chlorococcales Botryococcus Kuetzing 1849

Foraminiferal test linings

47 CHAPTER 2

Climate change driven black shale deposition during the end-Triassic in the western Tethys

Several new Triassic-Jurassic boundary sections from the Eiberg Basin (Northern Calcareous Alps, Austria) have been studied at high resolution. We present integrated geochemical and biological proxy data from this western Tethys shelf basin. High-resolution correlation of Kuhjoch, the Global boundary Stratotype Section and Point (GSSP) for the base of the Jurassic, Hochalplgraben and Tiefengraben shows that the initial and main Carbon Isotope Excursions (CIE) are contemporaneous with first and last occurrences of boundary defining macro- and microfossils. We focus on the end-Triassic initial CIE at the transition from the limestones of the Kössen Formation to the marls of the Kendlbach Formation. This change coincides with a dramatically increased influx of conifer (Cheiro- lepidiaceae) pollen and increased total organic carbon (TOC) values, succeeded by an acme of green algae (Cymatiosphaera). We present a model in which increased terrestrial organic matter influx is related to enhanced seasonality and increased erosion of the hinterland. Reduced salinity of the surface waters led to the mass occurrence of green algae. Stratification of the water column may have caused anoxic bottom water conditions and black shale deposition during the initial CIE at the base of the Kendlbach Fm.

48 CHAPTER 2

1. Introduction

The Triassic-Jurassic (T-J) transition period is marked by major environmental changes (e.g., McElwain et al., 1999; Beerling, 2002a). Marine and terrestrial extinction events and floral turnovers during the end-Triassic (Hallam, 2002; Tanner et al., 2004; Kiessling et al., 2007; Lucas and Tanner, 2008) coincide with possible perturbations of the global carbon cycle

(Hesselbo et al., 2002; Hesselbo et al., 2004). Climatic changes (T and CO2 concentration) in the end-Triassic (Cleveland et al., 2008) may be linked to the onset of volcanic activity of the Central Atlantic Magmatic Province (CAMP) that is related to the break-up of Pangaea. The T-J transition is characterized by two distinct negative organic carbon isotope excursions (CIE) in many sections within and outside the Tethys Ocean (Pálfy et al., 2001; Hesselbo et al., 2002; Guex et al., 2004; Ward et al., 2004; Ward et al., 2007; Williford et al., 2007;

Kürschner et al., 2007; Ruhl et al., 2009). Negative excursions in Ccarb- (Galli et al., 2005,

2007) and Coyster -isotope records (Van de Schootbrugge et al., 2007a; Korte et al., 2009) confirm the reality of global carbon cycle changes in the end-Triassic and transition to the Jurassic. The high resolution biological and organic geochemical proxy records of several well preserved sections including Kuhjoch, recently approved by the ICS as the Global boundary Stratotype Section and Point (GSSP) for the base of the Jurassic (Von Hillebrandt et al., 2007; Ogg et al., 2008), enable a good correlation within the Eiberg Basin. These records show several remarkable changes in the palaeoenvironment in the western Tethys realm at the transition from the Triassic to the Jurassic. We focus on the end-Triassic relatively short lived initial negative C-isotope excursion that precedes the main CIE at the base of the Jurassic. The initial CIE directly succeeds the transition from limestones of the Kössen Formation (Fm) to marls of the Kendlbach Fm and coincides with the deposition of organic rich, black and finely laminated sediments throughout the Eiberg Basin. Prasinophytes are often associated with black shale deposition throughout geological history (Riegel, 2008). The present study shows high resolution organic geochemical and palynological proxy records at the transition from the Kössen Fm to the Kendlbach Fm. We present a model to link the deposition of black shales to (regional) climate change.

2. Material and methods

Detailed lithostratigraphic, palaeogeographical and geological descriptions of the studied sections were previously reported in Kürschner et al. (2007), Von Hillebrandt et al. (2007), Ruhl et al. (2009) and Von Hillebrandt and Krystyn (2009).

49 CHAPTER 2

2.1 Geographical and geological setting

The Northern Calcareous Alps (NCA) is one of the few regions where continuous marine T-J boundary records are preserved. The Kuhjoch, Hochalplgraben and Tiefengraben outcrops are located in the Karwendel Syncline. This is an east-west trending synclinal structure within the Lechtal Nappe in the western NCA (Von Hillebrandt et al., 2007). Hochalplgraben and Kuhjoch are situated in the increasingly steep to overturned western and southwestern flank of the syncline. Kuhjoch is located at 47°29’02”N, 11°31’50”E, Hochalplgraben at 47°28’20”N, 11°24’42”E and Tiefengraben at 47°41’45’’N/ 13°21’00’’E (Fig. 1a).

10˚ 12˚ 14˚ 16˚

km A 0 50 100 49˚ 49˚

Vienna Munich

48˚ Salzburg 48˚ Tiefengraben Hochalplgraben Kuhjoch 47˚ 47˚

46˚ 46˚ 10˚ 12˚ 14˚ 16˚

NORTH SOUTH

Kuhjoch Hochalplgraben B Tiefengraben Kendlbach Fm Jurassic

Oberrh. Eiberg Basin Oberrhaet Limestone Hallstatt Basin Limest. Eiberg Mb Kössen Fm Hochalm Mb Rhaetian

Triassic Hauptdolomite Fm Dachstein Fm Norian

Figure 1: a) Location of the Kuhjoch, Hochalplgraben and Tiefengraben sections, b) North-South directed late Norian to Rhaetian facies cross-section of the north western Tethys shelf margin with the position of Kuhjoch, Hochalplgraben and Tiefengraben in the Eiberg Basin.

50 CHAPTER 2

2.2 Palaeogeography

During the Late Triassic the NCA together with the Southern Alps and the Dinarides formed an up to 300 km wide and approximately 500 km long shelf strip at the western end of the Tethys Ocean (Kürschner et al., 2007). Along this Tethyan passive margin extensive carbonate platforms developed, which were flanked by reefs rimming open shelf basins. At the very end of the Middle Norian, the Kössen Basin formed as a result of extensional tectonics (Hetenyi, 2002). By Rhaetian time, prograding siliciclastic sediments of the Kössen Fm over the Hauptdolomite Fm strongly modified and reduced the carbonate shelf (Haas, 2002; Krystyn et al., 2005). During deposition of the late Rhaetian Eiberg Member (Mb), which succeeds the Hochalm Mb within the Kössen Fm, the Eiberg Basin formed between the newly growing carbonate platforms of the Oberrhaet Limestone (Golebiowski, 1990). All three sections were deposited in the intra-platform Eiberg Basin (Krystyn et al., 2005) (Fig. 1b). The Kuhjoch and Hochalplgraben sections are located in the deeper central part of the Eiberg Basin in the western NCA. The Tiefengraben section was deposited in a more restricted shallow environment at the eastern end of the basin. The Eiberg Basin underwent continuous subsidence and reached water depths of up to few hundred meters in the latest Rhaetian. As a result, deposition was less affected by the end-Triassic sea level fall and fully marine conditions prevailed continuously in the deeper parts of the basin. A distinct lithological transition from limestones of the Kössen Fm (Eiberg Mb) to marly sediments of the Kendlbach Fm (Tiefengraben Mb) characterizes the sedimentary sequences in the Eiberg Basin during the T-J transition. Dark brown to black sediments of 2-3 cm directly succeed the Kössen limestones in Kuhjoch and Hochalplgraben (Fig. 3 and Fig. 5). On top of these beds follow the olive-grey marly to silty sediments of the Tiefengraben Mb. By contrast, the black sediments are missing in the Tiefengraben outcrop and olive-grey sediments directly follow the Kössen limestones. These sediments could be missing in the Tiefengra- ben section either due to a short depositional hiatus possibly related to the sea level low-stand or due to continued limestone deposition in the more proximal part of the basin time equivalent with shale and marl deposition in the deeper central part of the Eiberg Basin. Reddish oxidized, silty marls of the Schattwald beds succeed the 30-50 cm of grey Tiefengraben Mb sediments in Kuhjoch and Hochalplgraben, but are not developed in the Tiefengraben section.

2.3 Materials

The transition from the Kössen Fm (Eiberg Mb) to the Kendlbach Fm (Tiefengraben Mb) was sampled on high resolution (2-10 cm) at three localities (Kuhjoch, Hochalplgraben and Tiefengraben) in the Northern Calcareous Alps, Austria. Bulk stable C-isotope ratios and total organic carbon (TOC) content were measured on sedimentary organic matter from marly to clayey sediments. A high resolution palynological study was performed on the same samples.

51 CHAPTER 2

2.4 Methods: C-isotope and TOC measurements

Carbonate was removed from the sediments by rinsing 0.9 g of powdered sediment twice with 15 ml of 1 M HCl. Neutral pH values were reached by rinsing the residue twice with 22.5 ml demineralized water. After freeze drying, around 9 mg of homogenized de-carbonated sample residue was analyzed online for carbon content with a CNS-analyzer (NA 1500) following standard procedures. The TOC content of the sediment from Hochalplgraben and Tiefengra- ben was obtained by multiplying the carbon content of the de-carbonated sample by the ratio between the weight of the decarbonated sample and the original weight of the sample. TOC values from Kuhjoch were obtained by pyrolysis of bulk sediment on a Rock-Eval VI apparatus using lab procedures as described in Behar et al. (2001). Pyrolysis measurements were performed at the integrated laboratory of the Faculty of Geosciences (Utrecht University) 13 and The Netherlands Institute of Applied Geoscience TNO (NITG-TNO). The δ Corg values were measured on homogenized de-carbonated sample residue, containing 25µg of carbon, by Elemental Analyzer Continuous Flow Isotope Ratio Mass Spectrometry using a Fisons 1500 NCS Elemental Analyzer coupled to a Finnigan Mat Delta Plus mass spectrometer at the Geo- chemistry group of the Department of Earth Sciences, Utrecht University. Isotope ratios are reported in standard delta notation relative to Vienna PDB. Average analytical precision based on routine analysis of internal laboratory reference materials showed an error margin of 0.04‰ and lower.

2.5 Methods: palynological processing and analysis

Between 10 and 20 g of sediment was crushed into small fragments and dried for 24 h at 60 °C. A Lycopodium spore tablet was added to each sample. The samples were treated twice alternately with cold HCl (30%) and cold HF (40%) to remove the carbonates and silicates, respectively. The residue after chemical treatment was sieved with a 250 µm and a 15 µm mesh. Because of the large amount of mineral residue which still occurred in the samples

after sieving, ZnCl2 was applied to separate the lighter organic material from the heavier mineral particles like pyrite. This lighter material was transferred from the test-tube and sieving was repeated with a 15 µm mesh. The remaining organic material was mounted on two slides per sample with glycerine jelly. The slides are stored in the collection of the Section Palaeoecology, Laboratory of Palaeobotany and Palynology, Utrecht University, The Nether- lands. Pollen and spore identification was mainly based on Schulz (1967), Morbey (1975), Lund (1977) and Schuurman (1976, 1977, 1979). Around 300 terrestrial palynomorphs per sample were counted. Lycopodium spores were counted concomitantly, but they were excluded from the terrestrial palynomorph sum. The palynomorph concentrations (absolute number of palynomorphs per gram) were calculated based on the fossil palynomorphs counted, the Lycopodium spores counted, the dry weight of the sample, and the total number of Lycopodium

52 CHAPTER 2

spores added to the sample. Relative abundances were calculated and plotted using the Tilia and TgView computer programs (Grimm, 1991-2001).

3. Bio- and C-isotope stratigraphy of the Eiberg Basin

High-resolution correlation of several T-J boundary sections in the Eiberg Basin has been established by Bonis et al. (2009a), Von Hillebrandt et al. (2007) and Ruhl et al. (2009), based on biological and geochemical proxies, respectively. Here we present an integrated framework for the transition of the late Rhaetian Kössen Fm limestones to the marls of the Tiefengraben Mb in Kuhjoch, Hochalplgraben and Tiefengraben (Fig. 2). The onset of the initial and main negative C-isotope excursions of the Eiberg Basin are biostratigraphically confined by the last occurrence (LO) of the late Rhaetian ammonite Choristoceras marshi and the first occurrence (FO) of the Psiloceras spelae tirolicum ammonite. The latter is the marker for the base of the Jurassic (Von Hillebrandt et al., 2007; Von Hillebrandt and Krystyn, 2009). The established high resolution palynological zonation scheme for the Eiberg Basin (Kürschner et al., 2007; Bonis et al., 2009a) together with the FO of macro- and microfossil species and acmes of marine and terrestrial palynomorphs, support the correlation between different section in the Eiberg Basin based on C-isotope stratigraphy (Ruhl et al., 2009). The FO of Cerebropollenites thiergartii is approximately contemporaneous with the FO of Psiloceras spelae tirolicum in the Eiberg Basin and is suggested as a useful palynological marker for the T-J boundary (Bonis et al, 2009a). The very similar C-isotope record of the Tiefengraben and St. Audrie’s Bay (UK) sections enables correlation outside the Tethys Ocean (Kürschner et al., 2007). The FO of Cerebropollenites thiergartii pollen in Austria and the UK is time-equivalent relative to the C-isotope records. This further supports the recognition of the initial and main C-isotope excursions at the T-J transition interval in the Eiberg Basin and its correlation to the T-J boundary record of St. Audrie’s Bay (UK).

4. Results

Dinoflagellate cysts (mostlyRhaetogonyaulax rhaetica) and foraminiferal test linings are the main constituents of the aquatic palynomorphs at the onset of the initial CIE (Fig.3). The lower half of the initial CIE directly follows at the top of the limestones and coincides with black shale deposits. This interval is characterized by highly increased TOC values of up to 13 14% (Fig. 3). The negative shift in the δ Corg signature and increased TOC values in Kuhjoch

53 CHAPTER 2

cm Kuhjoch Hochalplgraben Tiefengraben

cm 500 2000

TPi cm 3000 5

4 TPi 0 1500 TPi cm 4

2500 1000

3 TH

1000 TH 2 TH TJB 2000 500 2 TPo

500 RPo RPo RPo

1500 1 0 1 Eiberg RL RL

-32 -30 -28 -26 -24 0 13 -32 -30 -28 -26 -24 δ Corg [‰] 13 δ Corg [‰]

5 1000 FO Parapsiloceras calliphyllum FO Ischyosporites variegatus 4 FO Psiloceras cf. pacificum Heliosporites reissingeri acme 3 FO Psiloceras ex gr. P. tilmanni Acanthotriletes varius acme RL 2 FO Psiloceras spelae tirolicum Conbaculatisporites spp. acme 1 LO Choristoceras marshi Pinuspollenites minimus acme

FO Cerebropollenites thiergartii Cymatiosphaera spp. acme 500

RL Rhaetipollis-Limbosporites zone RPo Rhaetipollis-Porcellispora zone TPo Trachysporites-Porcellispora zone TH Trachysporites-Heliosporites zone

TPi Trachysporites-Pinuspollenites zone 0 -34 -32 -30 -28 -26 -24 Fault 13 δ Corg [‰]

Figure 2: Correlation of Kuhjoch, Hochalplgraben and Tiefengraben based on C-isotope stratigraphy, palynomorph assemblage zones and bio-events. The palynomorph assemblage zones are based on Kürschner et al. (2007) and Bonis et al. (2009 a). Dotted lines are correlation lines based on C-isotope stratigraphy from Ruhl et al. (2009). The shaded area represents the high resolution studied transition interval from the Kössen Fm to the Kendlbach Fm. The lithostrati-graphy of these sections was given in Kürschner et al. (2007) and Von Hillebrandt et al. (2007).

54 CHAPTER 2

and Hochalplgraben co-occur with a peak abundance of up to 80% of Classopollis meyeriana relative to all palynomorphs. Terrestrial palynomorph concentrations from Kuhjoch show a dramatic increase to values of up to 30000 palynomorphs per gram (ppg) (Fig. 3). In Tiefen- graben, the lower half of the initial CIE, the highly increased TOC values and peak Classopollis meyeriana abundances are missing (Fig. 3). A spore dominated interval (Acanthotriletes varius, Baculatisporites spp., Calamospora tener, Carnisporites anteriscus, Conbaculatisporites spp., Convolutispora microrugulata, Deltoidospora spp. and Todisporites spp.) directly succeeds the Classopollis meyeriana peak abundance at Kuhjoch. Maximum spore abundance increases to 55% (~2650 ppg) relative to terrestrial palynomorphs. The spore interval coincides with 13 relatively constant and extreme negative δ Corg values of -31‰. This interval is succeeded by a mass abundance of green algae, the prasinophyte Cymatiosphaera spp., of up to 80% relative to all palynomorphs (Fig. 3 and Fig. 4). Aquatic palynomorph concentrations during peak Cymatiosphaera abundance are higher than terrestrial palynomorph concentrations in Kuh- joch. In the post initial CIE interval, Cymatiosphaera %, Classopollis meyeriana % and TOC % return to relatively low pre-excursion values and dinoflagellate cysts become abundant again. Cymatiosphaera abundance increases again (to ~23%) in the post initial CIE interval at 1570 cm in the more shallow and restricted Tiefengraben section. Observed changes in aquatic palynomorphs are similar in Hochalplgraben and Kuhjoch. Although sampling resolution is much lower in Tiefengraben, changes in relative aquatic palynomorph abundance are similar to the sections in the centre of the basin.

5. Discussion

The marine extinctions and (floral) turnovers during the end-Triassic and the transition to the Jurassic (Hallam, 2002; Tanner et al., 2004; Kiessling et al., 2007; McElwain et al., 2007; Lucas and Tanner, 2008) coincide with two marked negative C-isotope excursions of up to 7‰, in several records world-wide (Pálfy et al., 2001; Hesselbo et al., 2002; Guex et al., 2004; Ward et al., 2004; Galli et al., 2005; Galli et al., 2007; Kürschner et al., 2007; Ward et al., 2007; Williford et al., 2007; Van de Schootbrugge, 2008; Ruhl et al., 2009). Major extinction events in the Phanerozoic are often contemporaneous with anoxic events and black shale deposition. Some examples are: a late Early Cambrian extinction event (Hallam and Wignall, 1999 and references therein), the Frasnian-Famennian mass extinction (Pujol et al., 2006; Bond and Wignall, 2008), the Devonian-Carboniferous Hangenberg mass extinction event (Caplan and Bustin, 1999; Hallam and Wignall, 1999) and the end-Permian extinction event (Taka- hashi et al., 2009). The early Toarcian Ocean Anoxic Event coincides with major marine extinctions (Wignall and Bond, 2008 and references therein) and distinct negative excursions in the C-isotope signature (Hesselbo et al., 2000; Kemp et al., 2005). Also the base of the Jurassic in North West Europe (Hallam, 1995) and Canada (Wignall et al., 2007) is marked

55 CHAPTER 2 drier 0 0 0

wetter 0 3 0 0 Foraminiferal test linings Acritarchs Botryococcus sp. Prasinophytes Dinoflagellate cysts 0 0 2 0 0 0 Terrestrial ( grey ) Terrestrial & aquatic ( black ) 0 1 palynomorphs conc. [ppg] 0 0 0

0 0

sp. Botryococcus 1 1 0 0 0

8 8 8 Foraminiferal test linings test Foraminiferal 0 0 0

6 6 6 Prasinophytes 0 0 0

4 4 4 Aquatic Acritarchs 0 0 0

2 2 2

palynomorphs [%] Dinoflagellate cysts Dinoflagellate 0 0 0 4 8 0 6 0 0 3 6 [%] 0 4 0 0 2 4 0 2 0 0 1 2 polypartita Cymatiosphaera 0 0 0 0 6 0 0 8 8 0 5 0 0 0 [%] 6 6 4 0 3 0 0 4 4 0 2 0 0 Classopollis 2 2 meyeriana 0 1 0 0 0 4 1 8 6 2 CIE 1 5 CIE 0 CIE 6 1 4 8 4 3 initial 6 initial initial 2 TOC [%] TOC 4 2 1 2 0 0 0 4 5 5 2 2 - - 2 - ] 6 7 ‰ 2 2 - [ -

7 g

2 r - 8

9 o 2 2 - - 9 2 0 1 - 3 3 - - δ 13C 2 3 1 3 3 3 - - - Tiefengraben Kuhjoch Hochalplgraben 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 2 0 0 6 8 4 2 6 0 4 0 9 8 7 6 5 4 3 7 6 5 4 3 2 1 - - 5 4 5 5 6 6 7 7 3 2 2 2 2 2 2 2

(cm)

(cm) 1 1 1 1 1 1 1 1 Depth

(cm)

Depth Depth E. Mb E. E. Mb E. Tiefengraben Mb Tiefengraben Tiefengraben Mb Tiefengraben Tiefengraben Mb Tiefengraben Eiberg Mb Eiberg

Kendlbach Fm Kendlbach Fm K. Kendlbach Fm Kendlbach Fm Kössen Kendlbach Fm Kendlbach Fm K.

Rhaetian uppermost Rhaetian uppermost Rhaetian uppermost Dark grey marl Black shale interval interval interval interval interval interval interval ? Spore interval cursion interval Light grey marl Red marl (Schattwald beds) Limestone Classopollis Classopollis pre-excursion interval pre-e x Classopollis post-excursion interval post-excursion interval post-excursion interval Cymatiosphaera Cymatiosphaera Cymatiosphaera Cymatiosphaera Lithological key

56 CHAPTER 2

13 Figure 3 (left): Lithology, δ Corg curve and total organic carbon (TOC) of bulk sedimentary organic matter across the Kössen Fm-Kendlbach Fm transition in Tiefengraben, Kuhjoch and Hochalplgra- ben. Also indicated are the abundance of Classopollis meyeriana and Cymatiosphaera spp. relative to all palynomorphs, the composition of the aquatic palynomorph fraction and the absolute palynomorph concentrations. The dataset is available on request. The schematic humidity curve is based on PCA analysis (Chapter 3). by oxygen restricted facies and black shale deposition. Several new T-J boundary sections in the Eiberg Basin demonstrate high TOC values of up to 14% in the uppermost Rhaetian coinciding with the onset of the initial CIE (Ruhl et al, 2009). Finely laminated sediments in this interval suggest anoxic bottom water conditions. We propose a model to explain reduced mixing and anoxic conditions related to abrupt environmental changes in the end-Triassic Eiberg Basin.

Figure 4: a) Cymatiosphaera sp. with a fine wall structure (Hochalplgraben), b) Cymatiosphaera sp. with a coarse wall structure (Kuhjoch), c) Both Cymatiosphaera types from the peak abundance interval (Hochalplgraben)

5.1 Fresh-water influx and regional climate change in the Eiberg Basin

Our high resolution geochemical and palynological study across the initial CIE at the T-J transition interval in the Eiberg Basin shows dramatic changes in the relative abundance of certain marine and terrestrial palynomorphs. Abundant Classopollis meyeriana pollen at the base of the studied section may suggest warm and dry regional climatic conditions (Vakhrameev, 1981; 1987; 1991). A sudden and dramatic twentyfold increase in terrestrial pollen and spore concentrations of up to 30000 ppg marks the onset of the initial CIE. Classopollis meyeriana represent up to 90% of the palynomorphs in this interval. Increased pollen influx into the

57 CHAPTER 2

basin can be explained by an abrupt increase in seasonality in a semi-arid region. Classopollis meyeriana and amorphous organic matter bearing soils are washed into the Eiberg Basin, causing strongly increased TOC values. Additionally, the breakdown of carbonate production caused a condensed interval and therefore an increase in concentration of terrestrial organic material. A spore-dominated zone indicating wetter conditions directly follows the peak TOC values in Kuhjoch. Subsequent dramatically increased Cymatiosphaera abundance of up to 80% (relative to all palynomorphs) and increased absolute marine palynomorph concentrations suggest a basin-wide acme of this particular prasinophyte species. These green algae are known to prefer brackish to fresh water conditions as well as increased nutrient availability (Prauss and Riegel, 1989; Mudie, 1992; Guy-Ohlson, 1996 and references therein; Prebble et al., 2006). The prasinophyte acme in the Eiberg Basin coincides with the absence of foraminifera (Kürschner et al. 2007). Prasinophyte abundance appears sensitive to grazing, which may explain an additionally increased abundance when foraminifera are absent (Metaxas and Scheibling, 1996). The (almost) absence of foraminiferal linings and a decreasing number of foraminifera species (Von Hillebrandt et al., 2007) directly following the Kössen limestones may thus be an alternative explanation for part of the prasinophyte mass occurrence (Kürsch- ner et al. 2007). However, prasinophytes may be considered as disaster species that are highly successful in adapting to changing environments (Guy-Ohlson, 1996) as they survive widespread extinctions in the middle Palaeozoic (Tappan, 1980). The decrease in foraminifera may be caused by the calcification crisis resulting from acidification of the atmosphere and oceans by volcanic activity and methane release. In the Eiberg Basin, the initial CIE concurs with the prasinophyte acme and the absence of foraminifera which underlines the acidification hypoth- esis and consequent calcification crisis. A prasinophyte peak at the T-J boundary interval corresponding to the δ13C negative excursion is also reported from the Csővár Basin in Hun- gary (Götz et al., 2009) supporting the possibility of an interregional signal. Recently, Van de Schootbrugge et al. (2007a) suggested that green algal phytoplankton blooms may be symptomatic of elevated carbon dioxide levels in the atmosphere and oceans. A relatively wetter climate concurrent with the initial CIE is inferred from multivariate statistical analysis of palynological data from Hochalplgraben and Kuhjoch (Chapter 3). This is reflected by an increased spore abundance at the transition from the limestones of the Kössen Fm to the marls of the Kendlbach Fm in palynomorph records from the Eiberg Basin (Karle, 1984; Bonis et al., 2009a). Increased humidity led to increased run-off, enhanced erosion and increased supply of nutrients and terrestrial organic matter into the basin. Although the Kössen Fm represents a full marine basinal facies (Kuss, 1983; Golebiowski and Braunstein, 1988; Golebiowski, 1990), increased run-off and decreased salinity of the surface waters likely led to the formation of a fresh-water lens in the semi-restricted Eiberg Basin. This is reflected in the aquatic palynomorph abundance by an increase in prasinophytes and a decrease in dinoflagellate cysts (Fig. 3). The reduced salinity of the surface waters possibly led to stratification of the water column and anoxic bottom water conditions concurring with the initial CIE. This is supported by millimetre scale organic rich finely laminated sediments that

58 CHAPTER 2

are present throughout the Eiberg Basin at the base of the Tiefengraben Mb (Fig. 5). The sea-level lowstand coinciding with the initial CIE possibly caused a semi- but not completely closure of the Eiberg Basin that further enhanced a strongly reduced circulation. Further- more, the onset of the transgression during the T-J boundary interval could have caused a spread of oxygen-deficient bottom waters (Hallam, 1990b; Hallam, 1995) which may have contributed to increased preservation in shallower settings. Increased prasinophyte abundances through geological time are often related to temperate and cool waters, e.g., higher palaeolatitudes or in concurrence with major global glacia- tion episodes (Prauss, 2007). However, the lack of major glaciations in the late Triassic (Frakes et al., 1992; Satterley, 1996; Hallam and Wignall, 1999) and the relatively low palaeolatitude of the Eiberg Basin favour a decreased salinity and increased nutrient availability rather than lower sea surface temperature. Several short-term end-Triassic climate changes with CO2 values of up to 3000 ppm based on pedogenic carbonate nodules have been reported for the Rhaetian (Cleveland et al., 2008). A fourfold increase of atmospheric CO2 and global warming across the T-J boundary is also suggested from stomatal frequency values (McElwain et al., 1999). Whether these climate changes are caused by the onset of CAMP volcanism (Nomade et al., 2007) or mainly linked to the release of methane hydrates is still subject to debate. Possible end-Triassic global warming may have caused a pole ward expansion of the Hadley cells and a subsequent shift in climate belts. Modelling studies support a monsoonal climate in the Upper Triassic western margin of the Tethys Ocean (Sellwood and Valdes, 2007). The relatively low palaeolatitude of the Eiberg Basin during this period (Kent and Tauxe, 2005) could have caused an increased sensitivity for enhanced monsoonal activity in the Western Tethys realm. Alternatively, global warming could have led to an enhanced hydrological cycle which resulted into increased precipitation in the western Tethys realm.

5.2 Mesozoic black shale deposition and climate change

During geological history prasinophyte mass occurrences are often linked to black shale deposition (Riegel, 2008) which suggests a climatological and/or oceanographic control (Prauss and Riegel, 1989). Generally, deposition of black shales is explained either by the stagnation/preservation model or by the upwelling/enhanced marine productivity model (e.g., Heimhofer et al., 2006; Meyers, 2006; Negri et al., 2009). A combination of both mechanisms is also possible, with an increased freshwater and nutrient input via river runoff leading to stratification and increased productivity (Negri et al., 2009). The onset and main phase of black shale deposition in the Eiberg Basin resulted from increased terrestrial organic matter influx from the hinterland with enhanced preservation caused by a stagnant water column. The end of black shale deposition is marked by increased primary productivity of green algae as a result of less saline surface waters with increased nutrient conditions. Black shale events in the Mesozoic, particularly in the Cretaceous often resulted from changes in climate and the hydrological

59 CHAPTER 2

Figure 5: Thin section of finely laminated organic rich sediments at the base of the Tiefengraben Mb in the Eiberg Basin (sample from Schlossgraben, for description and lithostratigraphy see Ruhl et al., 2009).

regime, with a similar mechanism as we propose for the T-J transition period in the Eiberg Basin. A causal relationship between enhanced volcanism, gas-hydrate dissociation and climate change possibly led to Oceanic Anoxic Events (Jenkyns, 2003). The early Toarcian OAE is

marked by a warming climate trend, which was probably related to increased CO2 values (Hesselbo et al., 2000) and that was coupled to an enhanced hydrological cycle and riverine input, causing less saline surface waters, stratification and anoxic bottom water conditions in most NW European epicontinental basins (Mailliot et al., 2009). During the early Cretaceous Valanginian Weissert OAE, volcanism of a large igneous province was presumably responsible

for an increase of CO2, triggering a climate change and accelerated hydrological cycling (Erba et al., 2004).

60 CHAPTER 2

Semi-arid conditions at the Jurassic-Cretaceous boundary in Dorset (UK), partly inferred from a Classopollis dominated vegetation, are succeeded by more humid conditions in the middle Berriasian (Schnyder et al., 2009). An increasingly wetter climate led to enhanced freshwater influx from the hinterland and organic-matter enrichment (with TOC values of up to 8.5%) of littoral sediments in an evaporitic and semi-closed environment. An increased spore content was documented during this transition from a semi-arid to a semi-humid phase that was furthermore marked by acmes of various algal types (Schnyder et al., 2009). Enhanced terrestrial organic matter influx and preservation is thus linked to a climate-driven changing hydrological regime in the coastal area. Also mid-Cretaceous episodes of black shale deposition coincide with times of wetter climate (e.g., Meyers, 2006). A greenhouse climate was present due to increased rates of seafloor spreading and the evolution of associated volcanism-derived CO2. This warmer climate caused increased sea surface evaporation and subsequently a wetter climate and enhanced hydrological cycle (Meyers, 2006). Increased continental runoff would have delivered more soil derived nutrients to the coastal ocean and increased thermohaline stratification. This is a strikingly similar model as we propose for the end-Triassic. During the Cretaceous Barremian-Early Aptian, finely laminated black shales (e.g., Blätterton horizons, Fischschiefer) with TOC contents of up to 7% have been deposited in the Lower Saxony Basin (Mutterlose et al., 2009). Comparable with our palynomorph record from the Eiberg Basin, the laminated sediments show a strong influx of terrestrial organic matter, a low abundance of dinoflagellate cysts and an enrichment of prasinophytes. Although the main prasinophyte constituents are different (Pterospermella, Leiosphaera), this group is interpreted as an indicator for reduced salinity of surface waters (Below and Kirsch, 1997). Strong terrigenous influx with a subsequent weaker marine signal caused reduced salinity of the surface waters. This set the right conditions for laminated sediment deposition in the late Barremian (Mutter- lose et al., 2009). Salinity driven stratification of the water column during more humid periods was caused by significant freshwater runoff. Seasonally increased runoff from the hinterland caused also an increased surface water primary productivity (Mutterlose et al., 2009). A final example of black shale deposition related to enhanced continental runoff occurred in the late Aptian Vocontian Basin (western Tethys Ocean) (Heimhofer et al., 2006). This so-called ‘Niveau Jacob’ is characterized by enriched TOC values, a laminated sedimentary texture and strongly increased terrestrial palynomorph abundances. The increased fern spore abundance in these sediments probably indicates warmer and more humid conditions in the hinterland. Increased runoff caused stratification of the water column, anoxic bottom water conditions and increased preservation of the (largely terrestrial) sedimentary organic matter. In addition, the strongly enhanced input of land plant-derived organic matter also may have led to increasing oxygen-deficient conditions in the bottom waters (Heimhofer et al., 2006). Oxidation of the highly increased (sedimentary) organic matter content in the Eiberg Basin at the T-J transition could have further enhanced dysoxic conditions in the bottom waters during black shale formation.

61 CHAPTER 2

6. Conclusions

High-resolution correlation of Triassic-Jurassic boundary sections (Kuhjoch, Hochalplgraben and Tiefengraben) in the Eiberg Basin, based on biological and geochemical proxies, shows several remarkable changes in the palaeoenvironment. One particular event is the transition from the limestones of the Kössen Fm to the marls of the Kendlbach Fm. This transition is accompanied by the deposition of black shales that are present throughout the Eiberg Basin at the base of the Tiefengraben Mb. The lower half of the initial negative carbon isotope excursion coincides with increased TOC values, highly increased terrestrial palynomorph concentrations and peak Classopollis meyeriana abundance. This is interpreted as organic soils being washed into the basin caused by increased seasonality (see also Chapter 7). We propose a model in which increased humidity (supported by the higher abundance of spores) may have led to increased run-off, enhanced erosion and increased supply of nutrients and terrestrial organic matter into the basin. Increased run-off and decreased salinity of the surface waters led to the formation of a fresh-water lens in the semi-restricted Eiberg Basin. This is reflected by the highly increased Cymatiosphaera abundance. Increased absolute marine palynomorph concentrations suggest a basin-wide acme of this particular prasinophyte species. The reduced salinity of the surface waters probably led to stratification of the water column and anoxic bottom water conditions coinciding with the initial CIE. Also in other Mesozoic time periods,

changes in climate (e.g., CO2 increase) and an enhanced hydrological cycle were main drivers of enhanced organic matter preservation.

Acknowledgements

The authors acknowledge funding from the “High Potential” stimulation program of Utrecht University. We thank J. van Tongeren and N. Welters for assistance in the laboratory and A. van Dijk (Int. Geoch. Lab., UU-NITG) for performing the C-isotope analyses. We thank S. Hesselbo and an anonymous reviewer for their constructive reviews of the manuscript.

CHAPTER 3

Abrupt climate change during the end-Triassic mass extinction

The end-Triassic is a period characterized by mass extinction, large-scale volcanism, major perturbations in the global carbon cycle and climate change. High resolution palynological data and multivariate statistical analysis have been used to infer relative changes in temperature and humidity and are correlated with the initial negative carbon isotope excursion (CIE). The palynomorph record indicates that a conifer dominated sclerophyllous hardwood vegetation is replaced by a mixed spore plant (e.g., ferns, fern allies, mosses and liverworts) and gymnosperm vegetation. Abrupt warming coincides with the negative carbon isotope shift. The CIE is also simultaneous with a trend from a relatively dry to a more humid climate. We propose that the abrupt global warming at

the initial CIE is the result of CO2 outgassing from CAMP volcanism and additional marine methane release which caused a northward shift of the tropical summerwet biome on land adjacent to the western Tethys realm.

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1. Introduction

The latest Triassic (~200 Ma) is marked by a period of extinctions in both the marine and terrestrial realm, although the severity remains a point of discussion (Lucas and Tanner, 2008). Explanations for the biotic turnover have included both gradual (e.g., sea level change) and catastrophic mechanisms like volcanism or an impact (Olsen et al., 2002a; Tanner et al., 2004; Hesselbo et al., 2007). One of the most frequently proposed catastrophic mechanisms is large-scale basalt volcanism, known as the Central Atlantic Magmatic Province (CAMP) (e.g., Wignall, 2001; Hesselbo et al., 2002; Marzoli et al., 2004; Schaltegger et al. 2008). There is still controversy on climate change across the Triassic-Jurassic (T-J) boundary. General circulation modelling studies show that the Mesozoic earth was several degrees centigrade warmer than now, generating higher atmospheric humidity and a greatly enhanced hydrological cycle (Sellwood and Valdes, 2007). Climate model experiments indicate that the western margin of the Tethys is influenced by monsoonal climate with rains that started to peak in December through April, and evaporation exceeding precipitation for the rest of the year (Sellwood and Valdes, 2007). Proxy data indicating climate changes during the T-J boundary interval are sparse and often of low temporal resolution. A fourfold increase (from 600 to

2100-2400 ppm) of atmospheric CO2 across the T-J boundary was inferred from stomatal frequency analysis (McElwain et al., 1999), suggesting global warming. Carbon isotope compositions of pedogenic calcite from palaeosol formations indicate a relative stability of

CO2 (a rise of 250 ppm) across the boundary (Tanner et al., 2001). These contrasting results could be caused by differences in temporal resolution (Beerling, 2002a). Furthermore, the suitability of the pedogenic-isotopic palaeobarometer is questioned because of the compromising effect of massive dissociation events from methane hydrate reservoirs (Retallack, 2002a). Distinct changes in high resolution δ18O oyster records from the UK suggest a warming bottom water temperature trend from the Triassic to the Jurassic (Korte et al., 2009). By contrast, terrestrial palynomorph records of the Late Rhaetian to earliest Hettangian show a pronounced cooling (Hubbard and Boulter, 2000). Two significant negative carbon isotope excursions (CIE) (known as ‘the initial and the main negative CIEs’) have been recognized in many T-J boundary successions (e.g., Pálfy et al., 2001; Hesselbo et al., 2002; Ruhl et al., 2009). The short-lived initial excursion coincides with the latest Triassic biotic crisis (Lucas and Tanner, 2008). This excursion, of up to 7‰, has often been linked to CAMP volcanism, and/or methane hydrate release (Beerling and Berner, 2002; Hesselbo et al., 2002; Hesselbo et al., 2007). However, a direct causal link between this important perturbation in the carbon cycle and climate change has not yet been shown because T-J palaeoclimate proxy records do not have the same time resolution as the C-isotope records. Palynological data proved to be a useful tool to infer climate change based on the botanical affinities of the parent plants (Barrón et al., 2006; Galfetti et al., 2007). Hubbard and Boulter (1997, 2000) were pioneers in using a multivariate statistics approach on pre-quaternary palynological data to unravel

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changes in climate (thermophilly and drought tolerance). To summarize: a warmer, a cooler and stable climate across the T-J boundary have previously been put forward. In order to evaluate high resolution climate changes coincident with the initial negative carbon isotope shift, we applied a palynological, and multivariate statistical approach on two T-J boundary sections from the Eiberg Basin, Austria.

2. Materials and methods 2.1 Studied sections

The Eiberg Basin, which was located in the western Tethys realm, contains marine key sections to study the T-J boundary interval (Morbey, 1975; Kürschner et al., 2007; Von Hillebrandt et al., 2007; Ruhl et al., 2009; Bonis et al., 2009a, b). We focus on the Kuhjoch (recently approved by the ICS as the GSSP section for the base of the Jurassic) and Hochalpl- graben sections (Fig. 1). Present-day latitude of the sections is 47°, which corresponds to a palaeolatitude of ~27° during the end-Triassic (Kent and Tauxe, 2005). The Upper Triassic sediments in the Eiberg Basin are characterized by a lithological change from limestones of the Kössen Formation to marls and clays of the Kendlbach Formation (Fig. 3). The base of the Kendlbach Formation is characterized by black shale deposition, the initial negative CIE, an acme of prasinophyte cysts, and major changes in palynology (Ruhl et al. 2009; Bonis et

B Tethys ocean 8˚ 10˚ 12˚ 14˚ 16˚ 18˚ 50˚ 50˚ km N CAMP 0 50 100 Czech Republic 49˚ 49˚ Germany

48˚ 48˚ Hochalplgraben Austria Kuhjoch 47˚ 47˚

Italy 46˚ 46˚ 8˚ 10˚ 12˚ 14˚ 16˚ 18˚

Figure 1: a) Map (modified from Quan et al., 2008) showing the approximate position of the Central Atlantic Magmatic Province (CAMP) and the Eiberg Basin during the end-Triassic b) Present day location of the Kuhjoch (47°29’02’’N/11°31’50’’E) and Hochalplgraben (47°28’20’’N/11°24’42’’E) sections 66 CHAPTER 3

al., 2009b). Both sections were sampled on high resolution for stable organic carbon isotope- and palynological analysis and samples were processed according to standard procedures (Ruhl et al., 2009; Bonis et al., 2009b).

2.2 Multivariate statistical analysis

A summary of the relative pollen and spore abundances dataset from Kuhjoch and Hochalpl- graben was made using a linear ordination method, Principal Components Analysis (PCA), as the gradient lengths of the datasets did not exceed 3 SD (standard deviations) (Lepš and Šmilauer, 2003). All analyses were done with a square-root transformation of the species data and the data was centred by variables (taxa). The results from Kuhjoch and Hochalplgraben are displayed as the species scores on the first and second axis in PCA ordination diagrams (Fig. 2). The two main ordination axes are the dimensions through the dataset which explain the largest variance in species composition and can be translated in terms of the environmental and/or climatic gradient that controls the dataset.

3. Results

The first PCA axis from Kuhjoch explains 44.8% of the total variance within the dataset (Fig. 2). Most spores (produced by moisture-loving plants as ferns, liverworts and horsetails) score negative on the first axis.Vitreisporites bjuvensis is a pollen taxon produced by a seedfern with a hygrophilous affinity (Barrón et al., 2006) and has a high negative score on the first axis as well. Classopollis meyeriana has a high positive score on the first axis. Previous research by Vakhrameev (1981; 1991) shows a relationship between Classopollis pollen produced by Cheirolepidiaceous conifers and aspects of climate, suggesting that these conifers were thermophillous. During increasing aridification that limits the areal of the moisture-loving plants, this conifer species becomes dominant or even mono-dominant (Vakhrameev, 1987). Albian palynological assemblages from central Asia and southern Kazakhstan have a very low percentage of Classopollis corroborated by the appearance of fern remains which is interpreted as a wetter climate (Vakhrameev, 1987). However, as the Cheirolepidiaceae represent a large and well diversified group in the Mesozoic, they may have occurred under a wider range of environmental conditions. Not all Classopollis species have to be indicative of a (semi)arid climate, and a coastal habitat, or mangrove, has been suggested (e.g. Batten, 1974; Watson; 1988; Abbink, 1998). C. torosus belongs to the Hirmeriellae which probably represent coastal plants (e.g. Abbink, 1998). In each PCA ordination diagram C. torosus points to a different direction than C. meyeriana (Fig. 2 this chapter, and Fig. 4 in chapter 7). This means that their pollen producing parent plants had a different habitat and ecological preference.

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Conbaculatisporites spp. (F) A

Baculatisporites spp. (F)

2.4

+ Calamospora spp. (H) Acanthotriletes varius (CM) Carnisporites 1.2 anteriscus (CM) Classopollis meyeriana (Ch) Todisporites spp. (F)

Polypodiisporites polymicroforatus (F) 0.0 Ricciisporites † tuberculatus (L) * Araucariacites temperature Polypodiisporites australis (C) ipsviciensis (F) Vitreisporites pallidus (S) Trachysporites -1.2 Deltoidospora spp. (F) fuscus (F) Vitreisporites bjuvensis (S) PCA axis 2, 22.7% explained PCA

Classopollis - torosus (Ch) -2.4 Tsugaepollenites pseudomassulae (C) * Spores indet † Ovalipollis Rhaetipollis pseudoalatus (G) germanicus (G) -3.6 -4.5 -3.0 -1.5 0.0 1.5 3.0 4.5 + PCA axis 1, 44.8% explained - humidity

3.0 Classopollis torosus (Ch) B -

1.5 Classopollis sp. (Ch) Tsugaepollenites pseudomassulae (C) Classopollis meyeriana (Ch)

0.0

Baculatisporites spp. (F) humidity

-1.5 Concavisporites spp. (F) Ricciisporites Deltoidospora spp. (F) tuberculatus (L) PCA axis 2, 24.9% explained PCA +

-3.0

Vitreisporites spp. (S) Polypodiisporites polymicroforatus (F)

-4.5 -5.0 -4.0 -3.0 -2.0 -1.0 0.0 1.0 2.0 3.0 4.0 5.0 + - PCA axis 1, 38.7% explained temperature

Figure 2: Principal components analysis (PCA) ordination diagram of pollen and spore taxa in Kuhjoch (a) and Hochalplgraben (b). Abbreviations are of the parent plant group. C: conifer, Ch: Cheirolepidiaceous conifer, CM: club mosses, H: horsetails, F: ferns, L: liverworts, G: gymnosperms, S: seed ferns.

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While C. meyeriana is clearly produced by drought resistant plants, C. torosus might be adapted to a wetter environment (or were more salt tolerant). The first axis from Kuhjoch may then be interpreted as a ratio between palynomorph types indicative of relatively wet and relatively dry conditions at this location (Fig. 2). Classopollis meyeriana has high positive values on the second axis (which explains 22.7% of the total variance) while Classopollis torosus has high negative values. Studies from Greenland (Raunsgaard Pedersen and Lund, 1980; Koppelhus, 1997) and Siberia (Rovnina, 1972) report that Classopollis is rare and if it occurs in the record, it is mostly Classopollis torosus. The occurrence of mainly Classopollis torosus in high latitude records suggests that this pollen type was produced by a Cheirolepidiaceae species better adapted to colder conditions than the parent plant of Classopollis meyeriana. Therefore, the second axis from Kuhjoch is interpreted to represent a ratio between palynomorph types indicative of relatively cold versus relatively warm conditions. The high negative score on the second axis of the Vitreisporites bjuvensis confirms this interpretation, as this taxon is also known from Greenland (Raunsgaard Pedersen and Lund, 1980). The first PCA axis from Hochalplgraben, which explains 38.7% of the total variance, possibly represents a ratio between palynomorph types indicative of relatively cold and relatively warm conditions while the second axis (24.9% of total variance) likely represents a ratio between palynomorph types indicative of relatively wet and relatively dry conditions. This is mainly based on the relative position of Classopollis meyeriana and Classopollis torosus on axis 1 and the position of conifer- produced pollen on the positive side of axis 2, and spores together with Vitreisporites spp. on the negative side of axis 2.

4. Discussion

The palynological record from the Eiberg Basin reflects major changes in the palaeoenvironment at the base of the Kendlbach Formation, co-occurring with the initial CIE (Fig. 3). Plant groups were derived from the botanical affinities of the palynomorphs (Table 1). The palynomorph record shows a change from a conifer dominated hardwood vegetation to an increased spore plant abundance (e.g., ferns, fern allies, mosses and liverworts). An acme in pollen of thermophillous conifers, mainly Cheirolepidiaceae, characterizes the lowermost part of the Tiefengraben Member which is followed by an increase in seed fern and spore plant abundance. The sample scores on the first and the second axes from the PCA plotted through time results in relative humidity and temperature records (Fig. 3). During deposition of the lowermost part of the sections, the climate was characterized by a warm semi-arid subtropical climate. Abrupt warming coincides with the negative carbon isotope shift, where the most negative C-isotope values correspond with the maximum in temperature. There is a change to a more humid and relatively cooler climate with the return to more positive C-isotope values. One of the mechanisms to explain the initial CIE is the onset of CAMP volcanism

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Depth Sample score Sample score A 13 (cm) δ Corg [‰] axis 1: humidity axis 2: T Spores Plant groups 70

Kendlbach Fm 20 Tiefengraben Mb Tiefengraben 10

uppermost Rhaetian -10

-20

Eiberg Mb -32 -30 -28 -26 -24 -2 -1 0 1 2 -1 -0.5 0 0.5 1 0 20 40 60 0 20 40 60 80 100

Kössen Fm [%] [%] wetter drier colder warmer

Depth Sample score Sample score 13 B (cm) δ Corg [‰] axis 1: T axis 2: humidity Spores Plant groups 300

290

280

270

260

Kendlbach Fm 250 Tiefengraben Mb Tiefengraben

240

uppermost Rhaetian 230

-31 -29 -27 -25 -1 0 1 2 -0.8 0 0.81.2 0 20 40 60 0 20 40 60 80 100 E. Mb K. Fm [%] [%] colder warmer wetter drier

Dark grey Light grey Red marls Limestone Black shale Lithology marls marls (Schattwald beds)

Cheirolepidiacean Seedferns Mosses & Liverworts Ferns conifers Plant groups Other spore Other conifers Other Gymnosperms Fern allies producing plants

13 Figure 3: (color version on p. 205): Chronostratigraphy, lithostratigraphy, δ Corg data (Ruhl et al., 2009), PCA sample scores for axis 1 and 2, relative abundance of spores and the relative abundance of plant groups of Kuhjoch (a) and Hochalplgraben (b). The grey shaded band emphasizes the most striking changes.

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(Hesselbo et al., 2002). This large-scale flood basalt volcanic activity could account for both the input of 12C enriched carbon into the global carbon cycle as well as increasing temperatures due to increased CO2 values in the atmosphere. In addition, global temperature rise can cause dissociation of gas hydrates and injection of oxidized methane into the ocean/ atmospheric system (e.g., Jenkyns, 2003). Modelling studies suggest that volcanic CO2 outgassing fails to fully account for the light C-isotope signal (Beerling and Berner, 2002).

The ‘best-fit’ scenario is that warming, due to a build-up of volcanically derived CO2, triggers destabilization of seafloor methane hydrates and the catastrophic release of CH4 (Beerling and Berner, 2002, Chapter 4). Additionally, the heating of organic-rich shales and petroleum bearing evaporites around sill intrusions could have led to greenhouse gas and halocarbon generation in sufficient volumes to cause global warming, ozone depletion, and a negative carbon isotope excursion comparable to the end-Permian environmental crisis (Svensen et al., 2009). Global climate zones, or biomes, comparable with the ones we know from the present- day situation were also present during the end-Triassic and Jurassic periods (Kent and Olsen, 2000; Rees et al., 2000; Sellwood and Valdes, 2007). The Late Triassic tropical humid belt was of similar or perhaps slightly narrower width as the modern tropical humid belt (Kent and Olsen, 2000). The low palaeolatitude of the studied sections indicates that the Eiberg Basin was positioned near the boundary between a summerwet tropical and a semi-arid subtropical biome. A northward shift of these biomes, possibly triggered by volcanism induced global warming, may explain the transition from the Cheirolepidiaceae dominated vegetation, adapted to semi-arid conditions, to a vegetation dominated by moisture loving spore producing plants. Additionally, global warming may have led to an enhanced hydrological cycle which resulted into increased precipitation in the western Tethys realm. A comparison of our results with previous end-Triassic palaeoclimate proxy records is difficult because of their lower temporal resolution. Our results show for the first time direct evidence for abrupt climate change associated with the initial marine CIE and late Triassic mass extinction. A marked increase in greenhouse gasses as a result of CAMP volcanism and/ or destabilization of seafloor methane hydrates and the catastrophic release of CH4 gases during the initial CIE caused abrupt global warming. Although an increase in atmospheric

CO2 associated with a negative isotope excursion has been inferred from stomatal frequency analysis (McElwain et al., 1999) the exact stratigraphic correlation of the Greenland continental record with the marine realm is still unclear (Hesselbo et al., 2002). Also the precise stratigraphic correlation of the CAMP basalt flows with the initial CIE is still controversially discussed (Hesselbo et al., 2002, Hounslow et al., 2004, Whiteside et al. 2007). Nevertheless, the present study supports the idea that abrupt global warming was an important factor for the end-Triassic biotic crisis. Our results contradict with the previous palynological study by Hubbard and Boulter (1997, 2000), which showed a pronounced cooling event near the base of the Hettangian based on the disappearance of fernlike palynomorphs (temperate climatic conditions) and increase in a conifer dominated ‘boreal’ association. These contrasting results may be caused by the difference in temporal resolution. Furthermore, interpretations with

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regard to ecological preferences of the palynomorph producing plants differ. For example, Classopollis meyeriana, pollen produced by the thermophillous Cheirolepidiaceous conifers were mistakenly assigned to a cold-temperate climate. Geochemical evidence suggested a sudden intensification in continental weathering following global warming associated with the emplacement of the CAMP, perhaps amplified by methane hydrate destabilisation (Cohen and Coe, 2007). Our wetter climate trend inferred from palynology fits very well with these results. Additionally, a short humid event at the beginning of the Hettangian was reported from northern Spain (Barrón et al., 2006). A major open question remains about the pace and length of the climate change at the T-J transition. In the model from Beerling and Berner (2002) the initial excursion is assumed to took place over about 70 kyr. The CIE covers 40 cm in Kuhjoch and Hochalplgraben. As the warming occurs within 10 cm, this would imply about 15000 yrs based on constant sedimentation rates. This is a first order approximation and a better time frame is needed (e.g., cyclostratigraphy) for records which contain the initial CIE. The best documented example of similar rapid climate change is the Palaeocene-Eocene thermal maximum (PETM), ca 55.5 Ma. The PETM is also characterized by global warming, wetter climate conditions (i.e. an accelerated hydrological cycle), vegetation changes and a prominent negative C-isotope excursion (~5‰ in TOC records), with an estimated duration of the main body between 71.000 and 120.000 yrs (Wing et al., 2005; Nicolo et al., 2007; Sluijs et al., 2007). The CIE reflects the geologically rapid injection of13 C depleted carbon (by methane hydrate release and/or volcanism) into the global exogenic carbon pool (Sluijs et al., 2007), which is similar to the interpretation of the end-Triassic CIE. Palaeoclimate reconstructions from leaf margin analysis (temperature) and leaf area analysis (precipitation) based on PETM fossil floras from the Bighorn Basin (USA) show an abrupt warming (within ~ 10 ky), with a short dry period during the onset, and wetter conditions during the upper part of the event (Wing et al., 2005). This is similar to our end-Triassic record where also a brief drier period is evident from a small increase in xeromorph elements in the palynomorph assemblages (Cheirolepidiaceous conifers) in the lower part of the CIE which is shortly followed by an increasingly wetter climate as indicated by an increase in hygrophillous taxa (e.g. ferns, Caytoniales) (Fig. 3). A recent analogue to flood basalt volcanism is the 1783-1784 Laki eruption (e.g., Thordarson and Self, 2003). Despite the fact that this is small-scale it gives us some indications about the effects on climate. One of the mechanisms we haven’t dealt with thus far is the

emission of sulphur dioxide (SO2), and consecutive atmospheric cooling which lasted for 2-3 years after the Laki eruption. To date, there is little geological evidence for cooling associated

with continental flood basalt eruptions suggesting little long-term impact of SO2 emissions (Wignall, 2001). As this cooling operates over a short time scale, from months to decades (e.g., Wignall, 2001), it is very difficult to trace it back in deep-time records. Regional terrestrial acidification from sulphuric acid deposition during CAMP eruptions was suggested from German and Swedish palynological studies (Van de Schootbrugge et al., 2007b, 2009). At the T-J boundary, gymnosperm forests were abruptly replaced by an acidophile, low light adapted

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vegetation consisting of herbaceous ferns and fern allies (Van de Schootbrugge et al., 2007b, 2009). The increased spore abundance is interpreted as a pioneer assemblage commonly found in disturbed ecosystems, in this case caused by global warming and the release of pollutants during flood basalt volcanism. However, the existence of this spore spike is controversial. First, the spore dominated interval is restricted to the Triletes beds which means it could be caused by the local depositional environment. Second, the Jurassic ammonite fauna lacks representatives typical of the lowermost Jurassic Psiloceras planorbis subzone which indicates that a hiatus may be present at the top of the Triletes beds. Therefore it is possible that it is not a spore ‘spike’ but rather a prolonged interval of an increased spore abundance induced by climate changes. And as mentioned by Van de Schootbrugge et al. (2008) ‘the duration of the Triletes beds is difficult to constrain’. Third, the correlation of the Triletes beds to the Austrian Schattwald beds based on the C-isotope records is ambiguous (Van de Schootbrugge et al., 2008; Bonis et al., 2009a). Our results suggest a vegetation pattern across the T-J transition that is more complex on a regional scale rather than a global floral turnover (McElwain et al., 1999) or a decline in arborescent vegetation caused by the release of pollutants (e.g., sulphur dioxide) and toxic compounds (Van de Schootbrugge et al., 2007b, 2009). The present study shows that integrated high resolution C-isotope and quantitative palynomorph records are valuable in unravelling detailed patterns in end-Triassic climate change. To better understand the global significance of these patterns, more sections with a firm stratigraphic framework should be investigated in detail.

5. Conclusions

The terrestrial palynomorph records from two key T-J boundary sections, located in the western Tethys realm, were used to infer climate changes. Principal components analysis suggests that temperature and humidity were the main drivers for vegetation changes, based on the position of the pollen and spore taxa in the ordination diagrams. The end- Triassic initial CIE is simultaneous with a trend from a relatively dry to a more humid climate. Rapid warming (~15000 yrs) coincides with the negative shift. Peak warming coincides with the most negative C-isotope values. The low palaeolatitude of the Eiberg Basin implies that the sections were located close to the transition between the semi-arid subtropical biome and the summerwet tropical biome. We suppose that abrupt climate change caused by atmospheric greenhouse gas injection from large scale volcanism may have resulted into a northward shift of these biomes.

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Table 1

Botanical affinities of the palynomorphs found in the Kuhjoch and Hochalplgraben sections

Plant groups (Figure 3) Botanical affinity POLLEN Cheirolepidiaceous Conifers Coniferophyta - Cheirolepidiaceae Classopollis meyeriana Classopollis murphyi Classopollis sp. Classopollis torosus

Other conifers Other Coniferophyta Araucariacites australis Araucariaceae Pinuspollenites minimus Pinaceae Quadraeculina anellaeformis Podocarpaceae Perinopollenites elatoides Taxodiaceae Tsugaepollenites pseudomassulae Taxodiaceae Lunatisporites rhaeticus Voltziaceae Platysaccus spp. Voltziaceae Triadispora sp. Voltziaceae

Seed ferns Pteridospermophyta Vitreisporites bjuvensis Caytoniales Vitreisporites pallidus Caytoniales Vitreisporites spp. Caytoniales Alisporites diaphanus Corystospermales Alisporites radialis Corystospermales Alisporites robustus Corystospermales Striatoabieites aytugii Vesicaspora fuscus

Other gymnosperms Chasmatosporites apertus Cycadophyta - Cycadales Cycadopites spp. Cycadophyta, Ginkgoales, Peltaspermales Ephedripites spp. Gnetales Lagenella martini Ovalipollis pseudoalatus Rhaetipollis germanicus

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SPORES Mosses & Liverworts Bryophyta & Hepatophyta Mosses Bryophyta Stereisporites australis Stereisporites infrapunctus Stereisporites sp. div. Plant groups (Figure 3) Botanical affinity POLLEN Liverworts Hepatophyta Cheirolepidiaceous Conifers Coniferophyta - Cheirolepidiaceae Porcellispora longdonensis Classopollis meyeriana Ricciisporites tuberculatus Classopollis murphyi Classopollis sp. Fern allies (Horsetails & Club mosses) Sphenophyta & Lycophyta Classopollis torosus Horsetails Sphenophyta Calamospora tener Equisetales Other conifers Other Coniferophyta Araucariacites australis Araucariaceae Clubmosses Lycophyta Pinuspollenites minimus Pinaceae Aratrisporites parvispinosus Isoetales Quadraeculina anellaeformis Podocarpaceae Aratrisporites spp. Isoetales Perinopollenites elatoides Taxodiaceae Lycopodiacidites rugulatus Lycopodiales Tsugaepollenites pseudomassulae Taxodiaceae Densosporites fissus Selaginellales Lunatisporites rhaeticus Voltziaceae Foveosporites spp. Selaginellales Platysaccus spp. Voltziaceae Heliosporites reissingeri Selaginellales Triadispora sp. Voltziaceae Limbosporites lundbladii Selaginellales Densoisporites nejburgii Pleuromeia Seed ferns Pteridospermophyta Acanthotriletes varius Vitreisporites bjuvensis Caytoniales Camarozonosporites laevigatus Vitreisporites pallidus Caytoniales Camarozonosporites rudis Vitreisporites spp. Caytoniales Carnisporites anteriscus Alisporites diaphanus Corystospermales Carnisporites lecythus Alisporites radialis Corystospermales Carnisporites leviornatus Alisporites robustus Corystospermales Carnisporites spiniger Striatoabieites aytugii Carnisporites sp. div. Vesicaspora fuscus Cingulizonates rhaeticus Leptolepidites spp. Other gymnosperms Retitriletes gracilis Chasmatosporites apertus Cycadophyta - Cycadales Retitriletes sp. div. Cycadopites spp. Cycadophyta, Ginkgoales, Peltaspermales Tigrisporites microrugulatus Ephedripites spp. Gnetales Triancoraesporites ancorae Lagenella martini Ovalipollis pseudoalatus Rhaetipollis germanicus

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Ferns Pteridophyta Baculatisporites spp. Osmundales - Osmundaceae Conbaculatisporites spp. Osmundales - Osmundaceae Todisporites sp. div. Osmundales - Osmundaceae Perinosporites thuringiacus Filicales - Cyatheaceae Zebrasporites interscriptus Filicales - Cyatheaceae Zebrasporites laevigatus Filicales - Cyatheaceae Converrucosisporites luebbenensis Filicales - Dipteridaceae Deltoidospora spp. Filicales - Dicksoniaceae, Cyatheaceae, Dipteridaceae, Matoniaceae Concavisporites spp. Filicales - Matoniaceae Polypodiisporites ipsviciensis Filicales - Schizaeaceae Polypodiisporites polymicroforatus Filicales - Schizaeaceae Kyrtomisporis laevigatus Kyrtomisporis speciosus Lophotriletes verrucosus Trachysporites fuscus

Other spore producing plants Asseretospora gyrata Cosmosporites elegans cf. Cyclotriletes sp. Foveolatitriletes potoniei Platyptera trilingua Polycingulatisporites bicollateralis Semiretisporis gothae

Atmospheric methane injection caused end-Triassic mass extinction

The end-Triassic mass extinction (~201.5 Ma), marked by major terrestrial ecosystem changes and 50% marine biodiversity loss, has been attributed to a major volcanic episode during the break-up of Pangaea. Here, we present C-isotope data of long-chain n-alkanes derived from epicuticular waxes of land plants. Our data show a 5-6‰ negative excursion in compound specific C-isotope records, coincident with the extinction interval. This suggests strong 13C depletion of the end-Triassic atmosphere, within 5-10 kyr. The magnitude and rate of C-cycle disruption suggest the injection of ~7-8x103 Gt of isotopically light carbon from the methane-hydrate reservoir. Concurrent vegetation changes reflect a strong warming event and an enhanced hydrological cycle. Hence, our data, for the first time, link the end-Triassic events to massive carbon release and associated climate change.

78 CHAPTER 4

The end-Triassic mass extinction (~201.5 Ma; Schaltegger et al., 2008), one of the five major extinction events of the Phanerozoic (Raup and Sepkoski, 1982), is marked by up to 50% of marine biodiversity loss and terrestrial ecosystem changes (Raup and Sepkoski, 1982; Sepkoski, 1996; Olsen et al., 2002; McElwain et al., 2009). This event closely matches a 13 distinct negative excursion in δ CTOC (CIE) records (Hesselbo et al., 2002; Ruhl et al., 2009) and a potential fourfold increase in atmospheric CO2 concentrations (McElwain et al., 1999). The end-Triassic C-cycle perturbation has been attributed to large-scale carbon release, caused by a major volcanic episode with deposition of the Central Atlantic Magmatic Province (CAMP) during the break-up of Pangaea. However, deposition of this large igneous province continued for ~610 kyr (Whiteside et al., 2007), much longer than the ~20 to 40 kyr duration of the end-Triassic extinction event. The magnitude of the negative CIE also varies significantly between different geological basins, possibly due to changes in the source of the sedimentary organic matter. These observations question the reality of an end-Triassic global carbon cycle turnover and causal relationships to the mass extinction event (Van de Schoot- brugge et al., 2008; Van de Schootbrugge et al., 2009). Here we present compound specific C-isotope records from the western Tethys Ocean (including the Global Stratotype Section and Point (GSSP) for the base of the Jurassic), which span the end-Triassic mass extinction interval. Changes in the carbon isotopic composition of long chain n-alkanes, derived from epicuticular plant waxes, directly reflect changes in atmospheric13 C values since the carbon in these compounds is incorporated from atmospheric CO2. Furthermore, preservation of these organic molecules is unaffected by diagenetic alteration of the organic matter (Pancost and Boot, 2004). Hence, this allows for the first accurate reconstruction of the end-Triassic C-cycle perturbation, in a biostratigraphically well-constrained framework. Sediments for this study come from a 60 cm upper Rhaetian (latest Triassic) interval in the Kuhjoch and Hochalplgraben outcrops (Ruhl et al., 2009). These sediments were deposited in the intra-platform Eiberg Basin at the continental margin of the western Tethys Ocean (Von Hillebrandt et al., 2007). The studied interval in both sections directly succeeds the transition from limestones of the Kössen Formation (Fm) to marls of the Kendlbach Fm and coincides with marine extinctions and assemblage changes (Fig. 1). Organic compounds derived from plant waxes, are preserved in these over 200 Million year old, but immature sediments. Analyses of the chemical structure and carbon isotopic composition of the isolated n-alkane fractions, show a ~6‰ negative excursion from Rhaetian base values of ~-29‰, which is 2-3‰ larger than previously assumed (Beerling and Berner, 2002) (Fig. 2). This negative excursion in compound specific C-isotope records represents the first convincing evidence for end-Triassic atmospheric 13C depletion. The onset of the negative CIE probably transpired within ~5-10 kyr, based on the astronomically constrained ~20-40 kyr duration of the complete event. This suggests that large amounts of isotopically light carbon were rapidly released to the end-Triassic atmosphere. Major vegetation changes at the studied time interval potentially modified the magnitude of the negative CIE in the terrestrial higher plant C-isotope records, due to differential

79 CHAPTER 4

Kuhjoch, Austria

1200 Organic rich shales Organic poor shales 1100 Limestones LO Triassic organisms FO Jurassic organisms 1000 Assemblage changes

900 UK Bay, Audrie’s St.

800 Newark supergroup basins, USA Queen Charlotte Islands, Canada

700 Ammonites Jurassic Hettangian

600

500 Triassic Rhaetian

400 Pollen/spores Molluscs Echinoidea Foraminifera Ostracodes

Stratigraphic position (cm) 300

200 CAMP volcanism CAMP 100 Kendlbach Fm Tiefengraben Mb Tiefengraben 0 Methane event Triassic pollen/spores Triassic -100 Eiberg Mb Kössen Fm Palynology

-200 Foraminifera Dinoflagellates Ammonites -32 -31 -30 -29 -28 -27 -26 -25 -24 Ostracodes LO Conodonts

13 Radiolarian turnover bulk δ CTOC Bivalve extinction Theropod dinosaur turnover

13 Figure 1: End-Triassic δ CTOC negative excursion in concurrence with mass extinction. 13 The δ CTOC record of Kuhjoch (Ruhl et al., 2009), the GSSP for the base of the Jurassic, with the end-Triassic negative CIE closely matching continental and marine assemblage changes (Ward et al., 2001; Hesselbo et al., 2002; Olsen et al., 2002; Von Hillebrandt et al., 2007).

80 CHAPTER 4

carbon isotopic fractionation among plant groups. Although this is not suggested by changes in the Average Chain Length (ACL) of n-alkanes (Supplementary Information Figs. 2e and f), Cheirolepidiaceous conifer abundance strongly increases in concurrence with the negative CIE (Bonis et al., 2009). Modern conifer derived n-alkanes are relatively enriched in 13C, which is caused by lower stomatal conductance relative to other plant groups (Pedentchouk et al., 2008). A rapid transition from a mixed angiosperm-conifer flora to a purely angiosperm flora at the Paleocene-Eocene Thermal Maximum amplified the observed negative CIE by 1 to 2‰ (Schouten et al., 2007; Smith et al., 2007). Assuming similar physiological mechanisms in Mesozoic conifers as today, the observed end-Triassic higher plant n-alkane δ13C excursion is even dampened relative to atmospheric values. The actual magnitude of the end-Triassic C-cycle disruption may thus be even larger than observed. 13 The end-Triassic negative CIE in δ CTOC records was, based on model calculations, previously ascribed to the release of ~8000-9000 Gt carbon from volcanogenic gaseous CO2 in the CAMP, which also destabilized ~5000 Gt carbon from the methane hydrate reservoir (Beerling and Berner, 2002). However, the modeled release of carbon from these two reservoirs would result in a ~3‰ depletion of the exogenic carbon pool only. The ~6‰ magnitude and the short (~20-40 kyr) duration of the observed negative CIE do therefore not match with

CAMP related CO2 release. A simple mass balance calculation, using end-Triassic boundary conditions (Beerling and Berner, 2002), shows that 5-6‰ atmospheric-13C depletion is be best explained by the release of ~6900-8200 Gt carbon as methane (with δ13C values of -60‰). Combustion of subsurface organic rich strata, associated with flood basalt lavas, has also been proposed to have contributed to the magnitude of the end-Triassic C-cycle perturbation (van de Schootbrugge et al., 2009). Carbon release purely from this source (with δ13C values of ~-25‰), would involve an input at least two times larger than for methane, possibly as much as 23000 Gt. The injection of 6900-23000 Gt of carbon during the end-Triassic likely had a profound impact on global climate. Statistical analyses of palynological data spanning this time interval (Supplementary Information) show a strong warming event and enhanced hydrological cycle, directly coinciding with the onset of the negative CIE (Fig. 2). This suggests a strong causal relationship between massive carbon release, associated climate change and terrestrial ecosystem turnover. Massive carbon release to the atmosphere and subsequently the ocean, results in strong ocean acidification (Caldeira and Wickett, 2003; Zachos et al., 2005). In the end-Triassic, this likely caused major carbonate dissolution and stressed marine ecosystems and marine extinctions (Hautmann, 2004; Hautmann et al., 2008). Similar events of rapid carbon release to the atmosphere (e.g. at the Paleocene-Eocene Thermal Maximum (Zachos et al., 2005) and the in early Toarcian (Kemp et al., 2005), suggest a ~100 kyr duration for δ13C recovery of the exchangeable carbon reservoirs. This is in line with the residence time of carbon in the exogenic carbon pool (Dickens, 2003), but longer than the suggested ~20-40 kyr duration of the observed end-Triassic C-cycle perturbation. Dilution of the exchangeable carbon reservoirs with 13C enriched carbon from dissolving

81 CHAPTER 4

δ13C n-alkanes -38 -37 -36 -35 -34 -33 -32 -31 -30 -29 -28 40

0 Onset of bio- calcification crisis

-10 Stratigraphic position (cm)

-20

Kuhjoch ~ 5‰ -30 -34 -33 -32 -31 -30 -29 -28 -27 -26 -25 -24 - Temperature + - Humidity + 13 δ CTOC

δ13C n-alkanes -38 -37 -36 -35 -34 -33 -32 -31 -30 -29 -28 -27 -26 280 ~ 6‰ ~ 5‰ 270

260

250

240 Onset of bio- calcification crisis

230 Stratigraphic position (cm)

220 Hochalplgraben ~ 5‰ 210 -36 -35 -34 -33 -32 -31 -30 -29 -28 -27 -26 -25 -24 - Temperature + - Humidity + 13 δ CTOC 13 δ CTOC 13 δ C-[C25-C27-C29-C31-C33-C35] 13 δ C-[C17-C18-C19-C20-C21-C22-C23]

Figure 2 (color version on p.206) n-alkane biomarker C-isotope and climate proxy-records.

13 The combined n-C25 to n-C35 odd-chain-length (green) and n-C17 to n-C23 (blue) δ Cn-alkane signature from the end-Triassic mass extinction interval in Kuhjoch and Hochalplgraben. The 5-6‰ negative CIE coincides with strong relative warming and enhanced hydrological cycling based on statistical analysis of palynological data.

end-Triassic carbonate ramps under high CO2 conditions (Hautmann, 2004; Hautmann et al., 2008) can explain the rapid return of the atmospheric-δ13C signal to pre-excursion values. Enhanced surface ocean productivity and increased preservation of carbon in organic rich black shales coinciding with the negative CIE (Bonis et al., 2009b), likely contributed to the

82 CHAPTER 4

rapid recovery of the δ13C signal. This may be further enhanced by diminished background release of isotopically light carbon caused by changes in the gas-hydrate capacitor (Dickens, 2003). A delay in the recovery of the mixed terrestrial and aquatic (possibly algae; Han and

Calvin, 1969) short-chain n-alkane (n-C17 to n-C23) C-isotope records, relative to terrestrial long-chain C-isotope records (Fig. 2), suggests a more sluggish exchange between ocean and atmosphere. Such a partitioning column, with enhanced storage of 13C depleted carbon in the deeper parts of the ocean (Schouten et al., 2000). The end-Triassic mass extinction interval with rapid and large-scale carbon release, may now be regarded as a natural deep-time analogue similar to today’s anthropogenic carbon emissions. The present cumulative anthropogenic carbon release of over 5000 Gt (Caldeira and Wickett, 2003) will likely enhance greenhouse warming by several degrees (Allen et al., 2009) and substantially lower oceanic pH values (Caldeira and Wickett, 2003). Earths biosphere is also projected to experience major disruption of ecosystems, with associated loss of biodiversity (Sala et al., 2000). The direct link between massive carbon release and the end-Triassic mass extinction, strongly suggests that modern day ecosystems will experience a further loss of biodiversity, not only by habitat reduction but also by carbon release driven rapid climate changes.

Supplementary Information Materials

Two T-J boundary sections in the Northern Calcareous Alps (Kuhjoch, the Global boundary Stratotype Section and Point for the base of the Jurassic, and Hochalplgraben), where studied 13 for the δ C composition of long-chain n-alkanes. The study focuses on a 60 cm sedimentary interval at the very base of the Tiefengraben Mb (Kendlbach Fm) that succeeds the Eiberg Mb (Kössen Fm) in the uppermost Rhaetian. Sediments were deposited in the Eiberg Basin. This intra-platform shelf basin was situated on the continental margin of the western Tethys Ocean. This basin extended over several hundred kilometers from east to west and was located in between upper Triassic Oberrhät reef systems (Ruhl et al., 2009; Von Hillebrandt et al., 2007). Eight samples from Kuhjoch and nine samples from the Hochalplgraben section were 13 processed for this study. The same samples were previously studied for its δ CTOC composition (Ruhl et al., 2009). Eleven samples from Kuhjoch and ten samples from Hochalplgraben were previously studied for palynomorph abundances (Bonis et al., 2009a,b; Chapter 3).

83 CHAPTER 4

C25

C27 A

C29

C31 Relative abundace

C33

Time (min)

C25

C27

C29 Relative abundace

C31

C33

Time (min)

Supplementary Figure 1: Total ion chromatogram from the a-polar adductable fraction from sediments from (a) the first half of the negative carbon isotope excursion (CIE) at Kuhjoch (sample S-3; 0.5 cm) and (b) the second half of the negative CIE at Hochalplgraben (sample Hin-6; 252 cm). The labeled long-chain odd-carbon numbered n-alkanes are used to calculate the Average Chain Length (Fig.2e and f, supplementary information).

84 CHAPTER 4

13 Methods: δ Cn-alkane measurements

Between 10-20 g of fresh sediment of each sample was freeze dried and subsequently powdered. Organic compounds were extracted from the sediment, with a Dionex 200 Accelerated Solvent Extraction (ASE) equipment and a dichloromethane-methanol (9:1) solution. The

total lipid extracts were rinsed over a Na2SO4 column and separated in polar and a-polar compound classes with a hexane/dichloromethane (1/1) and hexane/dichloromethane (9/1) mixture, respectively. Elemental sulfur was removed from the a-polar fraction with activated copper. Copper fragments were activated with 2M HCl and subsequently rinsed with demi- water, methanol and dichloromethane. Straight chain n-alkanes were isolated from the a-polar fraction using the urea-adduction method. For this, the dry residue was dissolved in a 200 μl

methanol/urea (~10%, H2NCONH2, Merck) solution. Subsequently, 200 μl acetone and 200 μl hexane were added to the solution and than frozen and dried under nitrogen flow. The n-alkane compounds were captured in the formed urea crystals, which were subsequently washed with hexane to remove non-adductable branched and cyclic alkanes. Urea crystals, containing the adductable normal alkanes were then dissolved in 500 μl methanol and 500 μl distilled water. These compounds were subsequently extracted from the solution with hexane. This process was repeated 2-3 times to eliminate non-adductable alkanes from the residue.

Supplementary Table 1: The δ13C n-alkane composition of individual n-alkanes per sample. Standard deviations are given for samples that are measured in duplicate. The calculated Average Chain Length (ACL) is based on the n-C25 to n-C33 n-alkanes.

Kuhjoch Sample Stratigraphic δ13C n -alkane ACL position (cm) C17 C18 C19 C20 C21 C22 C23 C24 C25 C26 C27 C28 C29 C30 C31 C32 C33 C34 C35 C25 - C33 S-9 20 -28.1 -28.4 -28.1 -28.9 -28.1 -28.2 -27.7 -28.4 -28.2 27.3 S-8 15 -29.0 -30.6 -29.8 -28.5 -28.8 -28.3 -28.8 -28.3 -27.1 -28.0 -27.6 -30.6 -26.6 -28.6 -28.6 27.8 S-7 8 -33.7 -31.2 -35.2 -31.7 -32.6 -31.3 -32.0 -30.4 -30.3 -30.0 -30.2 -30.9 -28.9 -30.6 -30.6 27.6 S-6 (average) 7 -36.0 -36.1 -35.7 -36.9 -34.5 -34.0 -32.9 -33.2 -32.7 -31.7 -32.3 -32.6 -30.7 26.9 S-6 S.D. 1.6 0.6 0.3 0.0 0.3 0.0 0.7 0.6 1.6 0.4 1.9 1.7 S-5 6 -34.1 -35.7 -35.4 -36.3 -36.3 -35.8 -37.0 -34.7 -34.1 -33.0 -33.4 -33.3 -33.2 -32.8 -32.4 -32.5 -31.7 -32.9 -31.8 27.5 S-3 (average) 0.5 -37.2 -36.1 -36.0 -35.9 -37.0 -35.4 -35.6 -34.7 -34.5 -33.9 -33.7 -34.1 -33.6 -34.5 -35.8 27.8 S-3 S.D. 4.1 2.3 0.9 0.3 0.4 0.2 0.6 0.5 0.5 0.7 0.6 0.7 1.1 1.6 3.7 S-2 -1 -33.8 -33.5 -33.6 -34.2 -34.7 -34.7 -35.7 -34.7 -34.6 -34.2 -34.4 -33.6 -33.5 -33.4 -33.5 -33.4 -33.6 -33.6 -33.4 27.8 S-1 -4 -31.7 -31.4 -31.9 -32.7 -32.0 -32.3 -32.3 -32.8 -32.8 -33.8 -34.0 -34.1 -33.1 -31.8 -32.0 -30.8 28.3

Hochalplgraben Sample Stratigraphic δ13C n -alkane ACL position (cm) C20 C21 C22 C23 C24 C25 C26 C27 C28 C29 C30 C31 C32 C33 C34 C35 C25 - C33 Hin-9 273 -31.6 -30.1 -30.1 -29.0 -29.1 -28.7 -28.8 -27.3 -27.4 -26.9 -28.0 -25.8 -27.5 -26.0 27.6 Hin-8 266 -29.1 -29.4 -28.5 -29.0 -28.4 -28.2 -26.9 -27.1 -26.4 -27.6 -25.7 -26.3 -25.1 27.8 HinA-7 259 -31.6 -31.8 -32.1 -32.4 -31.8 -31.8 -31.1 -31.6 -31.1 -32.1 -30.6 -30.8 -31.9 -31.0 -29.4 28.2 Hin-6 252 -36.6 -37.1 -35.5 -38.4 -35.1 -34.0 -32.9 -33.5 -32.7 -31.2 -31.6 -29.8 -29.5 -27.9 -30.4 -29.8 27.2 Hin-5 245 -35.5 -36.9 -36.8 -37.8 -34.9 -33.9 -33.8 -35.1 -36.0 -33.0 -34.6 -30.8 -30.4 -28.4 -31.0 -30.9 27.2 Hin-4 242 -35.4 -36.9 -37.0 -39.6 -37.5 -36.4 -35.6 -35.7 -33.9 -34.2 -34.8 -33.9 -34.7 -33.5 27.0 Hin-2 235 -33.0 -30.0 -32.1 -30.7 -30.9 -30.7 -30.2 -32.3 -32.4 -31.3 -29.1 -28.5 28.7 Hin-1-A 230 -28.6 -29.8 -29.1 -29.3 -28.5 -30.4 -29.2 -30.7 -26.9 -27.2 29.1 HinB-4 (average) 229 -32.8 -31.3 -30.9 -31.3 -31.3 -30.4 -31.5 -30.9 -31.4 -30.0 -31.8 -30.4 -31.0 -27.1 -25.8 -25.0 29.0 HinB-4 S.D. 0.3 0.8 0.1 1.1 0.6 1.5 0.9 0.1 0.3 1.3 0.4 1.5

85 CHAPTER 4 13 13 δ Cn-alkanes δ Cn-alkanes -40-39 -38-37-36 -35 -34-33-32 -31-30-29-28-27 -26-25-24 -38 -37-36 -35 -34 -33-32 -31 -30-29 -28 -27-26 -25 -24 280 40 n-C25 n-C25 n-C 270 27 30 n-C27 n-C29 n-C29 n-C 260 31 20 n-C31 n-C33 n-C33 n-C35 n-C35 250 13 10 13 δ CTOC δ CTOC

240 0

230 -10 Stratigraphic position (cm)

220 -20 A B 210 -30 -40-39 -38-37-36 -35-34-33 -32-31-30 -29-28-27 -26-25-24 -38 -37-36 -35 -34-33 -32 -31-30 -29 -28-27 -26-25 -24 13 13 δ CTOC δ CTOC 13 13 δ Cn-alkanes δ Cn-alkanes -40-39 -38-37-36 -35 -34-33-32 -31-30-29-28-27 -26-25-24 -38 -37-36 -35-34 -33-32 -31 -30-29 -28 -27-26 -25 -24 280 40 13 n-C19 δ CTOC n-C17 n-C21 270 n-C20 30 n-C18 n-C22 n-C21 n-C19 n-C23 13 260 n-C22 20 n-C20 δ CTOC n-C23 250 10

240 0

230 -10 Stratigraphic position (cm) 220 -20 C D 210 -30 -40-39 -38-37-36 -35-34-33 -32-31-30 -29-28-27 -26-25-24 -38 -37-36 -35 -34-33 -32 -31-30 -29 -28-27 -26-25 -24 13 13 δ CTOC δ CTOC CPI CPI 0 1 2 3 4 5 0 1 2 3 280 40

270 30

260 20

250 10

240 0

230 -10 Stratigraphic position (cm) 220 -20 E F 210 -30 26 27 28 29 30 26 27 28 29

ACLC25-C33 ACLC25-C33

-25 -26 -27 -28 -29 -30 -31 -32

-33 -34

C n -alkanes per sample -36 -35 13 δ -38 -37 G H -40 -39 17 19 21 23 25 27 29 31 33 35 17 19 21 23 25 27 29 31 33 35 n-alkane c-chainlength n-alkane c-chainlength

Hin-9, 273 cm Hin-5, 245 cm S-9, 20 cm S-6, 7 cm Hin-8, 266 cm hin-4, 242 cm S-8, 15 cm S-5, 6 cm Hin-6, 252 cm Hin-2, 235 cm S-1, -4 cm S-2, -1 cm HinA-7, 259 cm Hin-1A, 230 cm S-7, 8 cm S-3, 0.5 cm 86 HinB-4, 229 cm CHAPTER 4

Supplementary Figure 2 (left, color version on p. 207): Individual high molecular weight

13 13 odd-carbon numbered δ C n-alkane signatures and δ CTOC signatures (in grey) from (a) Hochalplgraben and (b) Kuhjoch. Individual low to middle molecular weight odd-carbon numbered

13 13 δ C n-alkane signatures and δ CTOC signatures (in grey) from (c) Hochalplgraben and (d) Kuhjoch. The calculated Average Chain Length (ACL) compared to Carbon Preference Index (CPI) values from (e) Hochalplgraben and (f) Kuhjoch. The C-isotope composition of individual n-alkanes per sample from (g) Hochalplgraben and (h) Kuhjoch.

Conbaculatisporites spp. (F) A Baculatisporites spp. (F)

2.4

+ Calamospora spp. (H) Acanthotriletes varius (CM)

d Carnisporites e 1.2 anteriscus (CM) Classopollis n

i meyeriana (Ch) a

l Todisporites spp. (F) p x e r e u

t Polypodiisporites a % polymicroforatus (F) r 7 e

. 0.0 p

2 Ricciisporites m 2 tuberculatus (L) †

e , * Araucariacites t

2 Polypodiisporites australis (C)

i ipsviciensis (F) Vitreisporites x Trachysporites a pallidus (S) -1.2 fuscus (F) A Vitreisporites Deltoidospora spp. (F) C bjuvensis (S) P

- Classopollis torosus (Ch) -2.4 Tsugaepollenites pseudomassulae (C) * Sporesindet † Rhaetipollis Ovalipollis germanicus (G) pseudoalatus (G) -3.6 -4.5 -3.0 -1.5 0.0 1.5 3.0 4.5 + PCA axis 1, 44.8% explained - humidity

3.0 Classopollis Supplementary Figure 3: torosus (Ch) B Principal components

- Tsugaepollenites pseudomassulae (C) analysis (PCA) ordination 1.5 Classopollissp. (Ch)

d Classopollis diagram of pollen and spore e meyeriana (Ch) n i a l

taxa in (a) Kuhjoch and (b) p x e

0.0

Hochalplgraben. Abbrevia- y % t i 9 d . i 4 tions are of the parent m Baculatisporitesspp. (F) 2 u

, h 2 plant group. C: conifer, Ch: s i Concavisporitesspp. (F)

x -1.5 Ricciisporites a Deltoidospora spp. (F) Cheirolepidiaceous conifer, tuberculatus (L) A C

CM: club mosses, H: P + horsetails, F: ferns, L: -3.0 liverworts, G: gymnosperms, Vitreisporites spp. (S) Polypodiisporites polymicroforatus (F) S: seed ferns.

-4.5 -5.0 -4.0 -3.0 -2.0 -1.0 0.0 1.0 2.0 3.0 4.0 5.0 + - PCA axis 1, 38.7% explained temperature 87 CHAPTER 4

The n-alkanes were identified through mass spectra, molecular ion mass and retention time using a Thermo-Finnigan Trace Gas Chromatograph (GC) Mass Spectrometer (Thermo- Finnigan Trace DSQ). The δ13C composition of individual n-alkanes (Fig. 2a, b, Table 1, supplementary information) was measured by injection of the a-polar adductable fraction on a HP 6890N Gas Chromatograph (GC) coupled to a Thermo-Finnigan Combustion III and Thermo-Finnigan Delta Plus XP Isotope Ratio Mass Spectrometer (GC-IRMS) at the Molecular Biogeochemistry laboratory at the department of Earth Sciences, Utrecht Univer- sity. The GC temperature was programmed to increase 20°C/minute from 70-130°C and 5°C/ minute from 130-320°C. Based on standard analyses the accuracy was better than 0.2 and 1.6‰ (depending on standard, with the latter based on the squalane standard with δ13C= -19.54‰). Only relatively small amounts of the n-alkane fraction were recovered from most samples in the studied time interval, leading to only few duplicate sample measurements. Standard deviations on duplicate sample measurements are reported in Table 1 (supplementary information).

Methods: Multivariate statistical analysis on palynological data

The relative pollen and spore abundance of Kuhjoch and Hochalplgraben was summarised using a linear ordination method, Principal Components Analysis (PCA), as the gradient lengths of the datasets did not exceed 3 standard deviations (SD) (Lepš and Šmilauer, 2003). All analyses were performed with a square-root transformation of the species data. The two main ordination axes explain the largest variance in a multi-dimensional species composition dataset. The first axis of the Kuhjoch dataset represents 44.8% of the variance (38.7% in the Hochalplgraben dataset), and 22.7% of the variance is represented along the second axis (24.9% in the Hochalplgraben dataset) (Fig. 3, supplementary information). The two main axes are interpreted as climatic gradients that control the dataset. Based on ecological preferences of the palynomorph producing plants, the two main ordination axes are interpreted to reflect temperature and humidity changes through time (Fig. 2, main text).

Results

The a-polar residue of each sample mainly comprises middle (C17-C23) to long (C25-C35) n-alkanes, with the latter demonstrating a distinctly odd over even predominance in chain length distribution (Fig. 1, supplementary information). The individual n-alkane C-isotope

records of both sections, demonstrate distinct negative excursions of ~ 6‰ (n-C25 to n-C35) 13 and ~ 5‰ (n-C17 to n-C23) (Fig.2a, b, c and d, supplementary information). The δ C composition of individual n-alkane compounds per sample is relatively constant with increasing chain length in the Kuhjoch record (Fig. 2h, supplementary information). However, the C-isotope

88 CHAPTER 4

composition slightly increases with increasing chain length in the Hochalplgraben record (Fig. 2g, supplementary information). The influence of potential vegetation changes on the n-alkane chain length distribution was studied by reconstruction of the Average Chain Length (ACL) per sample. The maturity of the long-carbon-chain n-alkane fraction and the possible influence of migrated hydrocarbons on then -alkane biomarker C-isotope records are studied by reconstruction of the relative abundance of odd versus even numbered n-alkanes (carbon preference index (CPI)). The ACL and CPI are calculated using the following formulae (Smith et al., 2007):

(1) ACL = (25A25+27A 27+29A29+31A 31+33A33)/ (A 25+A27+A29+A31+A33)

(2) CPI = (((A25+A27+A29+A31+A33)/ (A 24+A26+A28+A30+A32)) + ((A 25+A27+A29+A31+A33)/

(A26+A28+A30+A32+A34))) * 0.5 A is the area under the chromatogram peak for every individual n-alkane biomarker (Fig. 1, supplementary information). The Average Chain Length decreases from 28 to 27 in the Kuhjoch dataset, but independent of changes in the C-isotope composition of the longer-chain n-alkanes (Fig. 2f, supplementary information). The Hochalplgraben dataset is marked by a similar decrease from 29 to 28, but with lower values down to 27, in between (Fig. 2e, supplementary information). Although lower ACL values in the Hochalplgraben record coincide with the negative excursion in the combined long-chain n-alkane C-isotope record (Fig. 2e, supplementary information), there is no statistically significant relation. CPI values for the studied interval at Kuhjoch and Hochalplgraben are relatively constant at ~1.5 and ~2 respectively. The CPI value of mainly one sample (at 242 cm) from the latter section increases in concurrence with the negative shift in n-alkane C-isotope records. CPI values of subsequent samples are however significantly lower even though long-carbon-chainn -alkanes remain relatively depleted in 13C. There is no significant relationship between then -alkane C-isotope record and CPI values.

Acknowledgements

We thank laboratory technician G. Nobbe, who assisted with most of the stable-isotope measurements presented in this paper and H. Visscher and A. Sluijs for suggestions and corrections during the drafting of this paper. M. Deenen is gratefully acknowledged for thorough discussion. MR, NB and WK acknowledge funding from the High Potential program of Utrecht University.

89 CHAPTER 5

Milankovitch-scale palynological turnover across the Triassic – Jurassic transition at St. Audrie’s Bay, SW UK

A high-resolution palynological study of the Triassic–Jurassic boundary in the St. Audrie’s Bay section revealed a palynofloral transition interval with four pronounced spore peaks in the Lilstock Formation. Regular cyclic increases in palynomorph concentrations can be linked with periods of increased runoff, and correspond to the orbital eccentricity cycle. The spore peaks can be related to precession-induced variations in monsoon strength. An implication is that the initial carbon isotope excursion lasted about 20 kyr. Emergence during deposition of the Cotham Member had influence on one of the peaks, dominated by spore-producing pioneer plants (e.g., horsetails and liverworts). There is no compelling evidence of a global end-Triassic spore spike which, by analogy with the K-T boundary fern spike, could be related to a catastrophic mass extinction event. Climate change is a more plausible mechanism for explaining the increased amount of spores.

90 CHAPTER 5

1. Introduction

The Triassic–Jurassic (T–J) boundary interval, spanning one of the ‘big five’ episodes of mass extinction, is characterized by large-scale volcanism (e.g., Marzoli et al., 2004; Schaltegger et al., 2008), major carbon cycle perturbations (e.g., Hesselbo et al., 2002; Ruhl et al., 2009), climate change (e.g., McElwain et al., 1999; Korte et al., 2009) and pronounced vegetation changes (e.g., McElwain et al., 2007; Bonis et al., 2009a). Explanations for the end-Triassic biotic crisis have included both gradual (e.g., sea-level change) and catastrophic mechanisms such as volcanism or a bolide impact (Olsen et al., 2002a,b; Tanner et al., 2004; Hesselbo et al., 2007). The climate was significantly different from today: there was no ice present in the high palaeolatitudes (e.g., Frakes et al., 1992; Satterley, 1996; Hallam and Wignall, 1999) and the thermal contrast between the large low-latitude continental interior and the sea drove a strongly developed monsoon circulation (Parrish, 1993; Buratti and Cirilli, 2007; Sellwood and Valdes, 2007). The monsoonal activity influenced precipitation patterns, and consequently the floral distribution and vegetation development. In palynological records, major biotic crises are sometimes characterized by a highly increased spore abundance. The best known example is at the Cretaceous-Tertiary (K-T) boundary, 65 million years ago, where a spore spike was described at the boundary clay level (e.g., Nichols and Johnson, 2008). For the T–J boundary, the evidence for a global spore spike is equivocal. A spore peak in the T–J transition interval was recognized in the Newark Basin (USA) and subsequently linked to an impact-induced mass extinction (e.g., Fowell and Olsen, 1993; Olsen et al., 2002a,b). By contrast, gymnosperm forests on land adjacent to the Eiberg Basin (Austria) were gradually replaced by ferns and fern-associated vegetation (Kürschner et al., 2007; Bonis et al., 2009a). Van de Schootbrugge et al. (2009) interpreted T–J fern proliferation as a pioneer assemblage after the disturbance of the terrestrial ecosystems by the release of pollutants during flood basalt volcanism. However, that interpretation is rather ambiguous because it does not account for possible changes in the pollen/spore ratio due to sedimentological facies changes and related taphonomical processes (Neves effect) or environmental changes that may have resulted in an increased dominance of spore-producing plants (e.g., Traverse, 2007). In St. Audrie’s Bay (Somerset, UK), a key T–J boundary section, an interval with increased spore abundance is present, 8-6 m below the first occurrence of Jurassic ammonites (Hounslow et al., 2004; Warrington et al., 2008). In this paper, we present new high resolution and quantitative palynological data from St. Audrie’s Bay in order to shed new light on the ongoing discussion of the dynamics of end-Triassic vegetation changes. We focus on the spore record, and consider whether the spore increase is caused by a sudden proliferation of pioneer vegetation in the aftermath of the end-Triassic crisis, or if other factors such as changes in climate and sea level are more important. In addition, we compare the spore record in St. Audrie’s Bay with records from other regions to assess any similarities, and whether they can be used for long-range correlation.

91 CHAPTER 5

2. Study area and lithology

St. Audrie’s Bay is a classic T–J marine boundary section located at the west Somerset coast in south-western England (Fig. 1). The lithology and depositional environment have been reported in detail by Hesselbo et al., (2004), Hounslow et al., (2004), and Warrington et al. (2008) (and references therein). A short summary is given below: The earliest Rhaetian Williton Member (upper part of the Blue Anchor Formation) was deposited in a shallow marine environment (Warrington et al., 2008) and represents the initial Late Triassic marine transgression in SW Britain (Hesselbo et al., 2004). The distribution of facies in the succeeding Westbury Formation was controlled by fluctuations in relative sea level (Hesselbo et al., 2004). Three sedimentary cycles, representing alternating deposition in transgressive, littoral, high-energy environments and lower energy, stagnant or weakly oxygenated water bodies, may be present (Warrington et al., 2008). The transition from the Westbury Formation to the lower Cotham Member (Lilstock Formation) represents a shallow- ing of the depositional environment from shelf to peritidal water depths (Wignall and Bond, 2008). A 0.5 m-thick unit of deformed beds in the middle of the member is followed by an erosional surface penetrated by deep cracks that are considered to reflect temporary emergence (Warrington et al., 2008). The exposure was extremely brief as suggested by the unstratified, single generation fill of these desiccation cracks (Hesselbo et al., 2004). The deformed beds are interpreted as a seismite (Simms, 2003; 2007). An alternative explanation could be a more prolonged period of seismic activity associated with the onset of the Central Atlantic Magmatic Province (CAMP) (Wignall and Bond, 2008). The upper Cotham Member represents a coastal environment (Mander et al., 2008) and it contains the initial negative carbon isotope excursion (Hesselbo et al., 2002). The Cotham Member-Langport Member junction is interpreted by Hesselbo et al. (2004) to represent a flooding surface. The Langport Member (Lilstock Formation) was either deposited in a warm and shallow saline lagoonal environment (Warrington et al., 2008), in a broad shallow seaway, or during sea-level rise on a carbonate ramp (Hesselbo et al., 2004). The depositional environment of the top of Langport Member is also disputed. It may have formed either during relative sea-level fall, with regression culminating in sea-floor erosion and emergence at the top of the Langport Member (Wignall and Bond, 2008), or during sea-level rise during the final drowning of the carbonate ramp (Hesselbo et al., 2004). The Blue Lias Formation was deposited during a phase of rapid flooding, indicated by development of a laminated, organic- rich shale (Hallam, 1995, 1997; Warrington, 2008). The appearance of the ammonite Psiloceras planorbis at the base of bed 13 in the Blue Lias Formation was proposed as a boundary marker for the base of the Jurassic (Warrington et al., 1994, 2008).

92 CHAPTER 5

-10˚ -5˚ 0˚

58˚ A km 0 50 100

Tethys ocean 56˚ CAMP

54˚

52˚

St. Audrie’s Bay

50˚

Cratonic landmasses Marginal marine - Fluviolacustrine Deep ocean

Figure 1: a) present-day location of the St. Audrie’s Bay section (51°11’N, 3°17’W) b) approximate position of the St. Audrie’s Bay section during the Triassic–Jurassic boundary interval (modified from Quan et al., 2008). CAMP = Central Atlantic Magmatic Province 3. Methods

Sixty-three rock samples from the St. Audrie’s Bay section were selected for palynological analysis. The samples range from the upper part of the Williton Member to the Blue Lias Formation. Average sample spacing was ~1 m, but 10 cm or less throughout the Lilstock Formation (Fig. 2). Between 5 and 20 grams of sediment was crushed into small fragments and dried for 24 hours at 60°C. A Lycopodium spore tablet was added to each sample. Subse- quently, the samples were treated twice alternately with cold HCl (30%) and cold HF (40%) to remove the carbonates and silicates. The residues were sieved using a 250 µm and a 15 µm mesh. ZnCl2 was applied to separate the lighter organic material from the heavier mineral particles. The lighter fraction was transferred from the test-tube and sieved once more using a 15 µm mesh. The remaining organic material was mounted on two slides per sample with glycerine jelly. The slides are stored in the collection of the Section Palaeoecology, Laboratory of Palaeobotany and Palynology, Utrecht University, The Netherlands. Pollen and spore identification was mainly based on Schulz (1967), Morbey (1975), Lund (1977) and Schuurman (1976, 1977, 1979). The identified morphotaxa of spores, pollen and aquatic palynomorphs are listed in Appendix 1. About 300 terrestrial palynomorphs were counted per sample (see palynomorph sums in Figs. 3 and 4). Lycopodium spores were counted concomitantly, but excluded from the terrestrial palynomorph sum. The palynomorph concentrations (absolute number of palyno-

93 CHAPTER 5

morphs per gram in samples) were calculated based on the fossil palynomorphs counted, the Lycopodium spores counted, the dry weight of the samples, and the total number of Lycopodium spores added to the sample. Relative abundances were calculated and plotted using the Tilia and TgView computer programs (Grimm, 1991-2001). Terrestrial palynomorph assemblages were established by constrained cluster analysis using CONISS (Grimm, 1987) within Tilia. A subsequent qualitative analysis, scanning two complete slides per sample, was carried out to check if rare palynomorph taxa were present which could be of biostratigraphic value. The complete presence/absence dataset is available on request. A linear ordination method, principal components analysis (PCA), was carried out on the relative pollen and spore abundance data. By relating the palynomorph taxa with their botanical affinity (see Chapter 7) we interpreted the axes and revealed temperature and humidity gradients in the data. Frequency analysis was performed on the terrestrial palynomorph concentration record (0.3-28.5 m) and the relative spore abundance record (12-14.5 m) using AnalySeries 1.1.1 (Paillard et al., 1996). Data were linearly detrended before the Blackman-Tuckey method was applied (compromise predefined level, Barlett window). Power spectra for each proxy record are reported in cycles/cm with a 90% confidence interval (Fig. 7a and 7b). A Gaussian band-pass filter from the main peaks of each proxy record is reported in Figure 7c.

4. Results 4.1 Terrestrial versus aquatic palynomorphs

In the lower part of the Westbury Formation (below 1010 cm) the palynomorph assemblages are generally dominated by terrestrial palynomorphs. Higher aquatic palynomorph abundance has been recorded in the Williton Member (up to 75%) and in the lower Westbury Formation (up to 68%) (Fig. 2). Between 1010 cm and 1580 cm most samples are dominated by aquatic palynomorphs. Between 1580 cm and the top of the section sampled the assemblages consist mainly of terrestrial palynomorphs (55%-97%). A remarkable cyclic pattern is visible in the palynomorph concentration. The Lilstock Formation is characterized by a very low total palynomorph concentration and a higher aquatic than terrestrial palynomorph content (Fig. 2). Most samples from the Williton Member and Westbury Formation have a higher terrestrial

Figure 2 (right, color version on p. 208): Chronostratigraphy, lithology, sea-level changes and

13 δ Corg after Hesselbo et al. (2004), palynological sample positions, terrestrial:aquatic palynomorph ratio, pollen:spore ratio and concentrations of palynomorphs through the Triassic– Jurassic transition in the St. Audrie’s Bay section. Spore peaks are indicated with a shaded band and a number (1-5).

CHAPTER 5

Peak

p S 0 5 3 2 1 4b 4a 0

s 0 h

p 0 r

o 4

l 0

a 0 p

l 2

o T 0 0 0 5

/g dry sediment]

3 palynomorphs

Aquatic 2

estrial palynomorphs estrial

e T 0 5 1 0

s 5 e

p S 0 0 0 3 Concentration of palynomorphs [grains x 10 0 0 2

0 n

l 1

o P

e r o 0

0 p S 1 0 8 0

p r 0

o 4

n Mudcracks

o concentrations Shell e

l n Percentage [%]

P 2 a

c i

a 0

u 0

q A 1

h 0

p r 8

o 0

n 6

l 'Beef' calcite Dolostone

a 4

i r t s

Percentage [%]

e r

r 2 e T 5 2 - Sand Limestone 7 [‰] 2 - org

C Carbon isotopes Carbon 9 13 2 - δ δ 1 3 - fall

SB TST & sea-level change sea-level & HST? RSLC

FSST-LST Sequence stratigraphy Sequence TST Dark grey mudstone Dark grey shale laminated Dark grey

rise

Lithology Height (m) Height

9 8 7 6 5 4 3 2 1 0

18 13 12 11 10 28 27 26 25 24 23 22 21 20 19 17 16 15 14 Palynological sample sample Palynological 2

189 143 135 134 133 132 131 130 129 127 125 123 122 121 120 119 118 116 115 114 113 112 111 110 109 108 107 106 105 98 87 50 43 34 23 14 80 70

290 273 263 254 239 229 213 206 201 197 194 192 168 165 161 159 155 151 149 191 141 102 179 170 Bed Number Bed

– 8 7 6 5 4 3 2 3 6 7 9 1 2 4 5 8 16

18 22 23 24 25

42 29 25 12 11 12 13 15 16 17 10

Williton Mr Williton Member Member

Blue Lias Formation Lias Blue Westbury Formation Westbury Langport Cotham

Formation Lithostratigraphy

Blue Anchor Anchor Blue Lilstock Fm Lilstock marl grey Pale mudstone Medium grey

Rhaetian Hettangian Stage

95 CHAPTER 5

than aquatic palynomorph content. Exceptions are the samples from 130, 400, 1010, and 1120 cm. The terrestrial palynomorph content is also higher in the Blue Lias Formation, except in some of the lowest samples (1460–1580 cm). The trends in terrestrial and aquatic palynomorph concentrations are simultaneous (Fig. 2).

4.2 Terrestrial palynomorphs

Significant changes occur in the pollen:spore ratio across the T–J boundary interval (Fig. 2). In the Westbury Formation the samples are characterized by a high amount of pollen. Only in the sample at 610 cm does spore abundance exceed 50% (spore peak 1). Samples from the Lilstock Formation have varying pollen and spore peak abundances. The total palynomorph concentration is very low in this interval (Fig. 2). Two spore peaks (2: 53% and 3: 68%) are present in the Cotham Member and two even higher spore peaks (4: 96% and 5: 76%) in the Langport Member. Pollen represents the most abundant palynomorph category throughout the Blue Lias Formation although there are some small increases in spore abundance. Within the terrestrial palynomorph fraction, four different assemblages (SAB1–SAB4) are distinguished, based on cluster analysis (Fig. 3): Assemblage SAB1. This assemblage occurs between the base of the section studied and the top of the Westbury Formation. It is characterized by high amounts of pollen, mainly Classopollis meyeriana, Classopollis torosus, Ovalipollis pseudoalatus and Rhaetipollis germanicus. In the lower part of the assemblage, C. torosus and O. pseudoalatus are more abundant than in the upper part. A prominent spore type is Ricciisporites tuberculatus with a major peak (71%) at 610 cm and minor peaks at 290 and 880 cm. Granuloperculatipollis rudis has its last common occurrence at the top of this assemblage. Assemblage SAB2 occurs between the base of the Cotham Member and 1275 cm, in the upper Cotham Member. The base of the assemblage is marked by an increasing abundance of Vitreispo- rites bjuvensis and Tsugaepollenites pseudomassulae. Another feature is an increase in the amount of spores, including Polypodiisporites polymicroforatus, Deltoidospora spp., Carnisporites anteriscus, Baculatisporites spp., Conbaculatisporites spp., Todisporites spp. and Concavisporites spp. Also of note is an acme of Porcellispora longdonensis at 1250 cm. Classopollis meyeriana, C. torosus and Ricciispo- rites tuberculatus are present throughout SAB2. Ovalipollis pseudoalatus, Rhaetipollis germanicus, V. bjuvensis and Lunatisporites rhaeticus have their last occurrence at the top of this assemblage. Two spore peaks are present (Fig. 4). Major components of the first one (2: 53%) at 1210 cm, are Polypodiisporites polymicroforatus, Porcellispora longdonensis, Heliosporites reissingeri, Concavisporites spp., C. anteriscus and Todisporites spp. The second peak (3: 68%) at 1250 cm consists mainly of P. polymicroforatus, C. tener, P. longdonensis, Deltoidospora spp. and Todisporites spp.. Assemblage SAB3 occurs between 1275 cm in the Cotham Member and the top of the Langport Member (1435 cm). A main feature is the absence of V. bjuvensis and T. pseudomassulae. Samples within this assemblage show a large variation in species composition. Classopollis meyeriana is the most abundant pollen in the lower samples and an acme of C. torosus (74%)

96 CHAPTER 5

occurs at 1390 cm. The most abundant spores are H. reissingeri, Deltoidospora spp., Acanthotriletes varius, Baculatisporites spp., Conbaculatisporites spp., Todisporites spp., Concavisporites spp. and Trachysporites fuscus. Two spore peaks are present (Fig. 4). The first peak (4: up to 96% spores) is based on four samples of which the lower two (1340 and 1350 cm: peak 4a) differ in species composition from the upper two (1360 and 1370 cm: peak 4b). The two lower samples consist mainly of Deltoidospora spp., A. varius, H. reissingeri, Concavisporites spp., Conbaculatisporites spp. and Trachysporites fuscus. The two upper samples have a high abundance of H. reissingeri, accompanied by Deltoidospora spp.. The major components of the second peak (5: 75% spores) at 1430 cm are Deltoidospora spp., A. varius and Conbaculatisporites spp.. Assemblage SAB4 occurs between 1435 cm till the top of the studied section. This assemblage is characterized by the dominance of C. meyeriana (75–100%) and common Pinuspollenites minimus. The amount of spores increases slightly at some levels, with H. reissingeri as a major component. The largest spore “peak” is only 22%. Of biostratigraphic importance are the first occurrences of Cerebropollenites thiergartii and Ischyosporites variegatus at 1850 cm.

4.3 Aquatic palynomorphs

The three most prominent aquatic palynomorph types in the record are dinoflagellate cysts (Heibergella sp. A, Rhaetogonyaulax rhaetica, Dapcodinium priscum), Micrhystridium spp., and cf. Leiosphaeridia (Fig. 5). Notably, Heibergella sp. A is only present in one sample in the Williton Member where it has an abundance of 88%. At the transition from the Cotham Member to the Langport Member there is a shift from an assemblage dominated by dinoflagellate cysts (mainly R. rhaetica) and prasinophytes to one dominated by prasinophytes (cf. Leiosphaeridia) and acritarchs (Micrhystridium spp.). Botryococcus disappears in the Langport Member at 1360 cm. Leiofusa jurassica is an acritarch of biostratigraphic importance which has its first occurrence at 1850 cm and a minor peak abundance (5%) in the Blue Lias Formation at 1890 cm.

4.4 Principal Components Analysis

The two main ordination axes are the dimensions through the dataset with the largest variance in species composition (Fig. 6). These axes are explained in terms of the environmental or climatic gradient that controls the dataset. The first axis represents a gradient from relatively warm (e.g., C. meyeriana) to ‘cold’ (e.g., C. torosus) palynomorph taxa (Fig. 6). The second axis represents a gradient from relatively dry (e.g., C. meyeriana and other conifers) to wetter taxa (spore producing plants). The sample scores on the second axis are used to derive a trend in relative humidity (Fig. 7). In general, climate was relatively dry throughout the time represented by the section studied. The Lilstock Formation is characterised by an interval of wetter phases, corresponding with the spore peaks.

CHAPTER 5

Orbell (1973) Orbell zone Rhaetipollis zone Heliosporites

zones Palynomorph assemblage Palynomorph SAB3 SAB4 SAB1 SAB2 s 4 2 e 2 r 2 a 0 2 u S 8 q 1 6 s

f 4

u 1 o

s 2 N 1 0 m 1 O u 8

l s h p r o a 4

o 2

m o

n T

s u t

5 0 y l a p

l a i r

a g e i r a t

v s u 3 0

s e

s u

s e r r e

r o j

a 1 0

. p p

t i r o p s e

s o

Other spores Other +

s e

t i r o p

y h

. p p

r o n i

y h t i r o p

m s e

s i

c a r

v a s e

t i r o p

. p p

s s i c n o

t i r o p

s e

t a l u s u

s i d o

r e g i n i p

T t i r o p

s i s i r e

c a b n o

s e

t n a t a l u

t i r o p

e e

c a r B

o s i n r a t i r o p

p C

s u i r a

s i n r

S a v

s e

t e l i r 4 0

. p p

s t o h

t n a 2 0

a r o p

c A

s o d i o

2 0

t l e D

8 0

i r e g n i

s u 6 0

s i e r

t a r o

s e

f o

r 4 0

c i

r e n e t i r o p t m

2 0

y l o p s o i l e

s i

a r o p

s e s o

2 0

t i r o p m a l a

s n e n o d g n o l

s i i d o p

2 0

a r o p y l o P

s i l l e

2 0

c r o P s u

t a l u

i i

6 0

c r e b u

s e

t r a g r e i

h 4 0

s i s u

t i r o p

s e

s i i

m r o

m i n i

c 2 0

c i

R s u

s e

f e a l l e n a

Other pollen Other

t r e p a

t i n e l l o p o r b e r e

e a l u

s e

t i n e l l o p

s a n i l u

s i l a r C

s a

s u n t i r o p

s u a

i P

s o

t a s e

t i

m o d u e

c e a r d a u Q

s u d i l l a p

s p

s a h

c a i r a s e

s e C

t i r o p s e d i o

c u a r A

s i e r

t a l e s u

s u n a h p a i d

c i

t i V

s e

t i n e l l o p e a g u

t e a h r T

t i r o p

s i d u r

s e

s i l A . p p

t i r o p

s i l l o p

n i

s i

s e

t i n e l l o p o n i r e P

t a l u s

l u

l c Cerebropollenites

t a n u L

s n e g i

s u t i p o d o a

s P e

s i a r o p

t i r o p

s a

s n e c r e p o l u n a r G

c i

v u j b

s u

s e V

s a l a n o

s e

c i n a

z n E

t i r o p

m r e g

s i e r 2 0

t i V

s i l l o p i

s u

2 0

t e a h

t a l a o d u e

s p

4 0

s i l l o p i l a 2 0

v O

8 0

s u

6 0

s o r o t

4 0

s i l l o p o s

2 0

s a l C 1 0

8 0

a n a i r e

6 0

y e

4 0

s i l l o p o Percentage [%] s

2 0

s a l C

0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Height (cm) Height 0 0

0 0 found after a qualitative analysis is indicated by a black dot. black a by indicated is analysis qualitative aftera found Relative pollen sporeand abundances through theTriassic–Jurassic transition theinSt. 70 60 50 4 30 20 10 00 9 9 80 70 60 50 4 30 20 10 00 800 40

2 2 2 2 2 2 2 2 1 1 1 1 1 1 1 1 1 1

80 70 60 50

30 20 10

Williton Mr Williton

Member Member

Lithostratigraphy

Westbury Formation Westbury Blue Lias Formation Lias Blue Langport Cotham Formation

Lilstock Fm Lilstock Blue Anchor Anchor Blue

Stage Rhaetian Hettangian

Psiloceras planorbis Psiloceras FO FO shown;importanttaxabiostratigraphicallyare and abundant mostthe section. OnlyBay Audrie’s the complete dataset is available on request. The shadedfirst Theoccurrence portionblack. plotted of in abundances the exaggeration of of the curves is a 5 times thiergartii Figure 3:

CHAPTER 5

zones

Palynomorph assemblage Palynomorph

n s

. i SAB2 SAB3

Spore peak number peak Spore

p s

u c s

p s

u 5 3 2

estrial palynomorph sum [n] sum palynomorph estrial

a 4b 4a r

s e

s i

e m

t e

2 1 2 8 1 6 1 1 8 4 5 0 7 2 6 6 6 9 4 2 6

7 6 9

i e

r 5 8 4 3 7 6 1 7 9 3 9 1 1 0 9 6 9 2 0 4 7 T 3 2 5 t

i 3 2 3 3 2 2 3 1 3 3 3 2 3 3 2 2 2 3 4 2 3

r o

p s o

p s u c s Other Spores Other

n n

i t

r r

p s y

a e

n s

h c

u s c

p s

e . s

r b

r T

p s e

o o t r

p s

o r C

p i

e p s

o i

S t

p s

r T

r i v o

p s a c C

t n

a o

u c

i r

g 0

i s s i

e 0 r

i 0 s r

u o 4 i

p s

a v

o 0 i

s 2

s t

p t l

i 0 a r

l t p s

4 o

u c

r t

i s

o n 0 n

e 2 p s

u a c

n t

d s

d i

t g 0

t i

n 2

o o

r p s a

e 0

i c

c 2

e i

p s

e c 0 o

s 2 o

p s u

P o

s 0 i

i a

l 2

a m a

s t s

t u

i 0 t

a r

n 2

e u

Polypodiisporites polymicroforatus Polypodiisporites

s t

s o a

i c

u c

p s

Other Pollen Other

s e

t u t

r e n

r o

a c

e P

u h

i r

p s

. s

A e

p l

a t

p s e

i r

p s

n s o

m s

s u

n p s

a t

t s

u s s

i C p

h u

o n n

p i

u l

u c

e o

i a

l d

l a

d l p

a c y

r s

d s

e r o C

e s s

e p

u i

P i t l

o i l

p s

i i

p s

t i

n s

l e

r l p

l u c t

i i

u c

n V

i t p

a r

e e

r p

i s

l n

u s

s T e v

a u l u

t j r

l G p

a t

o e

e 0 t

i d a

r 2 u

e s

p s

p 0

t 2 l i

V o

l 0

a v 2 O

u s

r 0

t 6

i l

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o 0

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o P record and the lithology from the spore P org

3 0 C

Aquatic palynomorphs Aquatic 0 13

errestrial palynomorphs errestrial δ

0 T 1

0 /g dry sediment]

ollen 3 P 8 6

4 Spores 2 0 Concentration [grains x 10 4 2 - Enlargement oftherelative terrestrial palynomorph 5 2 - [%] 6 2 - org 7

C - Carbon isotopes Carbon 8 2 Initial shift 13 - δ 9 2 - 0 3 -

14 13 12 Palynological sample Palynological 98

130 129 127 125 123 122 121 120 119 118 116 115 114 113 112 111 110 109 108 107 106 105 102

Bed number Bed 2 3 4

Cotham Mbr Cotham Langport Mbr Langport

Lithostratigraphy Lilstock Fm Lilstock peak interval in the Lilstock Formation. The shaded portion of the curves is a 5 times exaggerationtimes 5 a curvesportion is the ofshaded TheLilstockFormation. the intervalin peak 2).(seeFig. numbered areSpore peaks black. plotted in abundances the of Figure 4 (color version on p. 209): p. on version (color 4 Figure abundances and concentrations in relation to the to relationconcentrations in and abundances

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m [n] m u

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f o n i Percentage [%] D

0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Height (cm) Height Relative aquatic palynomorph abundances through the Triassic–Jurassic transition

0 0

0 0 70 60 50 4 30 20 10 00 9 9 30 20 10 00 80 70 60 50 4 800 40

2 2 2 2 2 2 2 2 1

80 70 60 50 30 20 10 1 1 1 1 1 1 1 1 1

Williton Mr Williton

Member Member

Lithostratigraphy

Westbury Formation Westbury Blue Lias Formation Lias Blue Langport Cotham Formation

Lilstock Fm Lilstock Blue Anchor Anchor Blue

Stage Rhaetian Hettangian igure 5: F section.Bay St.Audrie’s the in

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4.8 Heliosporites reissingeri +

3.6

Deltoidospora spp. 2.4 Acanthotriletes varius

d (humidity) Conbaculatisporites spp. e Todisporites spp. n

i Trachysporites fuscus a l 1.2

x p Polypodiisporites polymicroforatus e Calamospora tener 8 %

7 . 0.0 1

, Vitreisporites bjuvensis 2

s Granuloperticulatipollis rudis x i

a -1.2 Ricciisporites tuberculatus A

C Rhaetipollis germanicus P Classopollis torosus -2.4

- Ovalipollis pseudoalatus Classopollis meyeriana -3.6 -6 -4 -2 0 2 4 6 + PCA axis 1 , 46.5% explained (temperature) -

Figure 6: Principal components analysis (PCA) biplot of the pollen and spore percentage data. The plot shows a gradient from relatively warm to cooler taxa along axis 1, while axis 2 represents a drier to wetter vegetation gradient.

y s h e p p A C ra to g ) ti m so a c i r ( n st t o e o h b Sample score axis 2 g h ig r Terrestrial palynomorph ta it e a Spore abundance S L H C concentration (17.8 % explained) 2.0E+12 Power spectrum: Terrestrial palynomorph conc [ppg] 2800 1.8E+12 100kyr Eccentricity peak, 2700

with 470cm periodicity n

1.6E+12 a 2600 90% confidence interval i g

n 2500

1.4E+12 a t t 2400 e

1.2E+12 H 2300

1.E+12 n 2200 o i t 8.E+11 a 2100 m Relative power r

o 2000 F

6.E+11 s 1900 a i L 4.E+11

e 1800 u l

2.E+11 B 1700 1600 0.E+00 0.000 0.004 0.008 0.012 0.016 0.020 1500 t r r e o b p 1400 m g m F

Cycles/cm e n

a M k L c B 80000 r e o 1300 m t b a s h l m t i e Power spectrum: Spore abundance [%] o L C M 1200 Precession peak, with 100cm periodicity n a

70000 i

90% confidence interval t 1100 e a

60000 h 1000 R n

o 900 i t 50000 a 800 m r

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Relative power 30000 e

W 400 20000 300 r 200 r

10000 M

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u 0 l F 0.00 0.04 0.08 0.12 0.16 0.20 B -31 -29 -27 -25 0 100 200 0 20 40 60 80 100 -1.0 0.0 1.0 2.0 3 Cycles/cm δ 13Corg [‰] [grains x 10 /g dry sediment] Percentage [%] - Humidity +

Figure 7 (color version on p. 210): Power spectra of (a) the terrestrial palynomorph concentration with a main periodicity of ~470 cm and (b) the relative spore abundance with main periodicity of ~100cm. Gaussian band-pass filters reflect periodic changes in both proxy records (c) and are tentatively assigned to the astronomical ~100 kyr eccentricity and ~20kyr precession cycles.

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4.5 Frequency analysis

Time-series analyses of terrestrial palynomorph proxy data show strong oscillations throughout the T–J transition (Fig. 7c). Six periodic increases in terrestrial palynomorph concentrations, of up to 200x103 palynomorphs/gram sediment, coincide with low relative spore abundances. A power-spectrum (Fig. 7a) reflecting the main periodic oscillation in the terrestrial palynomorph concentration record, is marked by a >90% significant peak with a ~470 cm periodicity. Large fluctuations in relative spore abundance are confined to the Lilstock Formation. At least four periodic fluctuations with a ~100 cm periodicity (Fig. 7b) coincide with one ~470 cm periodic cycle in the terrestrial palynomorph concentration record. The fundamental precession frequencies of the Earth’s orbital parameters decreased to ~20 kyr in the early Jurassic (Berger et al., 1992). The duration of eccentricity cycles however, remained constant. The distinct wavelengths of the oscillations in our proxy-records can be tentatively linked to the orbitally controlled precession and ~100 kyr eccentricity frequencies that are corrected for the late Triassic.

5. Discussion 5.1 Palynology

The St. Audrie’s Bay palynological record is characterized by a marked palynofloral change within the Cotham Member at 1275 cm. Conifer-dominated hardwood vegetation was replaced by a mono-dominant Cheirolepidiaceous forest (Fig. 3). Orbell (1973) proposed older (Rhaetipollis) and younger (Heliosporites) palynomorph zones from the British T–J transition and placed the boundary between these at a rapid decline in the abundance of Ovalipollis ovalis, Rhaetipollis germanicus and Ricciisporites tuberculatus. This boundary corresponds to the boundary between SAB2 and SAB3 (Fig. 3) but its exact position is questionable as the decline in the numbers of these taxa is not usually synchronous (Warrington, 2005). Our results are in good agreement with the previous palynomorph studies of St. Audrie’s Bay (e.g., Hounslow et al., 2004; Warrington et al., 2008). However, the present higher resolution reveals new palynological findings. The first is that there are two transitional zones within each of the palynofloras recognized by Orbell (1973). The upper part of theRhaetipollis zone, SAB2, shows an acme of Vitreisporites bjuvensis and Tsugaepollenites pseudomassulae and an increase in relative spore abundance (e.g., Porcellispora longdonensis, Polypodiisporites polymicroforatus). The lower part of the Heliosporites zone, SAB3, is characterized by the absence of O. pseudoalatus and Rhaetipollis germanicus. This zone also has a high abundance of spores, dominated by Heliosporites reissingeri, Deltoidospora spp., and Acanthotriletes varius. The second is that the first occurrence (FO) ofCerebropollenites thiergartii appears to be a useful biostrati-

102 CHAPTER 5 spp. Calamospora Osmundaceae spores Dipteridaceae- Matoniaceae spores Polypodiisporites Polypodiisporites polymicroforatus Most abundant taxa Northwest Germanic Basin (~80%)

Anapiculatisporites Kyrtomisporis Porcellispora Reticulatisporites Todisporites Verrucosisporites Dictyophyllidites Deltoidospora Granulatisporites* Converrucosisporites* Most abundant taxa Newark Basin (89%) 6 28 19 14 9 7 9 7 18 14 8 6 5 8 7 7 5 5 67 81 6 15 34 max % spp. spp. (>5%)

spp. spp. spp. spp. spp. spp. spp. Deltoidospora Heliosporites reissingeri Heliosporites Acanthotriletes varius Deltoidospora Conbaculatisporites Trachysporites fuscus Calamospora tener Polypodiisporites polymicroforatus Todisporites spp. Deltoidospora reissingeri Heliosporites Concavisporites anteriscus Carnisporites Porcellispora longdonensis Todisporites Concavisporites Porcellispora longdonensis Polypodiisporites polymicroforatus tuberculatus Ricciisporites Heliosporites reissingeri Heliosporites Conbaculatisporites Deltoidospora Acanthotriletes varius Most abundant taxa 78 68 53 71 96 75 Total % Total 4a 3 2 1 4b 5 Peak nr. ~1345 1250 1210 610 ~1365 1430 Depth (cm) St. Audrie’s Bay(96%)

and The most abundant Table 1. sporesofthespore peaks from St. Audrie’sBay,theNewark Basin (FowellOlsen,and 1993; Fowell et northwestthe 1994)and al., GermanicBasin(Van de Schootbruggeet2009).al., The bracketsfollowing percentage in the location name is the maximum totalrelative spore abundance. infirmus *Granulatisporites Converrucosisporitescameronii aredominating thefern-spike assemblagess(Fowell Olsen,and 1995).

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graphic marker enabling the correlation between terrestrial and marine realms (Kürschner et al., 2007; Bonis et al., 2009a). In the St. Audrie’s Bay section the FO of this pollen taxon is close to that of Psiloceras planorbis (Fig. 3). Ischyosporites variegatus and the acritarch Leiofusa jurassica have their FO in the Blue Lias Formation at the same level as C. thiergartii (Figs. 3 and 5). Also in an earlier study from St. Audrie’s Bay, Leiofusa jurassica was found in the T–J transition interval (Hounslow et al., 2004). A Jurassic acme of this species was recorded from Danish sections and from the Northwest Germanic Basin (Lund, 1977; Dybkjær, 1991). According to Lund (1977) the basal part of the Pinuspollenites-Trachysporites zone is characterized by this acme. This is in concordance with results from the Eiberg Basin (Austria, western Tethys), where the base of the Pinuspollenites-Trachysporites zone correlates with the Jurassic TH zone, containing the FO of C. thiergartii (Bonis et al., 2009a). Van de Schootbrugge et al. (2007a) reported an FO of Leiofusa jurassica slightly lower in the section (~1820 cm). Finally, the increased spore abundance in the Lilstock Formation as described by Hounslow et al. (2004) and Warrington et al. (2008) now appears to comprise multiple spore peaks (Figs 2-4), the main constituents of which are summarized in Table 1; the nature and cause of these peaks are discussed below.

5.2 Climate change

Orbitally induced variations in solar radiation (Milankovitch cycles) have exerted a strong influence on the Earth’s climate throughout geological time (Berger et al., 1992; Olsen and Kent, 1996; Van der Zwan, 2002; Popescu et al., 2006; Ruddiman, 2006). The amount of incoming solar radiation in present-day low-latitude systems depends mainly on precession. Maximum precession leads to maximum insolation and corresponds to times of maximum monsoon intensity (Vollmer et al., 2008). Changes in monsoonal activity have immediate consequences for atmospheric circulation (Crowley et al., 1992), the magnitude of precipitation rates (Vollmer et al., 2008), runoff, and weathering patterns which potentially translates terrestrial changes to the marine realm (Crowley et al., 1992). The large Pangaean landmass may have intensified the monsoon system because the larger land area could retain more heat (e.g., Crowley et al., 1992). A modelling study by Kutzbach (1994) suggests that rainfall and runoff would undergo cyclic changes with periods of 23,000 yrs over a substantial part of the Pangaean (sub)tropics. The low palaeolatitude (~30°) of the St. Audrie’s Bay section (Kent and Tauxe, 2005) suggests an influence of monsoonal activity. Cyclic fluctuations in palynomorph concentrations (with a ~470 cm periodicity) are present in the St. Audrie’s Bay record (Figs. 2 and 7). Additionally, four distinct peaks of relative spore abundance (with ~100 cm periodicity) occur within one of the longer cycles. Periods of increased spore abundance probably reflect wet phases, which may be related to intensified monsoon activity on the precession scale. The different periodic oscillations in terrestrial palynomorph records may be tentatively assigned to astronomical climate forcing with a ~100 kyr eccentricity and

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~20 kyr precession wavelength (Fig. 7). A qualitative and quantitative analysis of the palynofacies revealed orbital cycles in the lower Jurassic in the southern UK (Waterhouse, 1999a, b). In agreement with the present study, Waterhouse (1999b) reports an average of five precession cycles in one eccentricity cycle. The precession cycle acted mainly on the terrestrial environment, probably via climate- controlled variations in runoff that affected terrestrial organic debris (Waterhouse, 1999a), and the 100 kyr eccentricity cycle controlled relative sea-level (Waterhouse, 1999b). However, in the absence of ice-sheets (e.g., Frakes et al., 1992; Satterley, 1996; Hallam and Wignall, 1999) during the T–J period a 100 kyr sea-level cycle is unlikely. Late Triassic (Norian) playa cycles in the Mid-German Basin were associated with varying monsoon activity (Vollmer et al., 2008). The palaeoclimate model for the Mid-German Basin implies highest rainfall when summer solstice passed through perihelion (earth closest to the sun) in the northern hemisphere (Vollmer et al., 2008, p.12, fig. 9). Insolation was highest, Tethyan seawaters evapo- rated, and moisture was transported to the north, where it precipitated in the playa system (monsoonal maximum). Insolation was lowest when the northern hemisphere summer occurred during passage of the solstice through aphelion. Decreasing evaporation of Tethyan seawater supplied less moisture and a drier climate resulted (monsoonal minimum). The wet and dry period in the Mid-German Basin occur within a single precession cycle. A similar palaeoclimate model may be applied to the St. Audrie’s Bay record, with the spore peaks representing monsoonal maxima. Also Late Triassic palynological data from Northern Spain revealed peak abundance of hygrophytic plants that may reflect the strong monsoon precipitation regime (Gómez et al., 2007). Olsen and Kent (1996) considered that precession-related cycles in precipitation (including the powerful effect of the eccentricity cycles on precession) were a consistent feature of tropical climate during most times in Earth history. A modelling study by Crowley et al. (1992) suggests that climate responses to 100 kyr eccentricity forcing can occur over low-latitude land areas involved in monsoon fluctuations. Kemp and Coe (2007) recognized 100 kyr eccentricity cycles in the Late Triassic at St. Audrie’s Bay on the basis of the rock colour. The Newark Basin succession shows that lake levels were controlled by Milankovitch modulation of the monsoon systems of Pangaea (Olsen and Kent, 1996; Olsen and Kent, 1999); unfortunately, pollen and spores were too sporadically preserved to produce time-series data for that succession (Olsen and Kent, 1996). The most intense rainy seasons occurred when the precession cycle results in perihelion in northern hemisphere summer during times of maximum eccentricity. We suggest that the high terrestrial palynomorph concentrations in St. Audrie’s Bay are linked to an abrupt increase in seasonality in a semi-arid region (enhanced monsoonal activity) and more runoff during eccentricity maxima (Fig. 7c). Coincidence of occurrences of high total organic carbon (TOC) values and black shales with the high terrestrial palynomorph concentrations support this interpretation (Ruhl et al., 2010.). An intensified monsoon system may induce extension of the climate belts. Because rain penetrated further into the hinterland, vegetation such as Cheirolepidiaceae and other gymnosperms could cover a larger area, and enhanced seasonal

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runoff would transport a relatively large amount of pollen, as is reflected in the low spore abundance during eccentricity maxima (Fig. 7c). A similar pattern, with enhanced seasonal runoff causing high terrestrial palynomorph concentrations, has been suggested for the Eiberg Basin in Austria (Bonis et al., 2009b).

5.3 Sea-level change

A marine extinction scenario driven by sea-level fall and the loss of shallow-marine habitat space has been put forward (Hallam and Wignall, 1999), while according to Hesselbo et al., (2004) ‘it is unlikely that sea-level fall played a significant role in the T–J boundary extinctions in either a local or a global context’. A main question is to what extent sea-level changes influenced the observed palynomorph distribution patterns, as the marginal marine facies of the St. Audrie’s Bay section may have been highly sensitive to such changes (Hesselbo et al., 2004). A transgression represented by the upper part of the Blue Anchor Formation (Hesselbo et al., 2004; Warrington et al., 2008) is confirmed by a peak abundance of the dinoflagellate cyst Heibergella sp. A which was reported by Palliani and Buratti (2006), but any details about this cyst are currently unknown. The transgressive systems tract represented by the Westbury Formation (Hesselbo et al., 2004) is reflected in the high abundance of dinoflagellate cysts in the aquatic palynomorph association; the sample fom 1010 cm in particular has a very high concentration (and relative abundance) of the dinoflagellate cystRhaetogonyaulax rhaetica (Figs. 2 and 5), a feature described by Orbell (1973) as a Rhaetogonyaulax population bulge possibly related to a decrease in the salinity. Only two dinoflagellate genera (Dapcodinium and Beaumontella) are found in the basal Jurassic deposits. Prasinophytes as well as acritarchs dominate the Blue Lias Formation which is in agreement with Van de Schootbrugge et al. (2007a), who suggested that seawater was warmer, had a lower salinity and that conditions prone to stratification and the development of anoxia were created. Sea-level changes could influence the relative amount of spores in the record. Spores are relatively heavy and more difficult to transport than pollen (e.g., Dybkjær, 1991). During low sea level larger amounts of spores might be expected, while at high sea level pollen with a better buoyancy (e.g., bisaccates) would predominate. At St. Audrie’s Bay the spore spike interval that occurs in the Lilstock Formation is characterized by a sea-level lowstand (Fig. 2; Hesselbo et al., 2004). The Cotham Member was deposited in an environment that was subject to sub-aerial exposure (Hesselbo et al., 2004) and spore peak 3 coincides with a level with desiccation cracks (Fig. 4). One of the most abundant spores in this peak is Porcellispora longdonensis, a large, and heavy spore which may imply that sea-level change influenced the presence of this peak. Furthermore, P. longdonensis and Calamospora tener are both produced by pioneer plants (bryophytes and horsetails, respectively) which could have invaded the newly emergent coastal areas. Another feature of the St. Audrie’s Bay record is spore peak 4b, which consists almost completely of Heliosporites reissingeri. In the NW German Basin, the relatively

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common occurrence of Heliosporites in the brackish Lower Rhaetian and marine Hettangian suggests that the parent plant inhabited a coastal environment, possibly in a marsh or mangrove/swamp (Lund, 2003). Heliosporites was produced by lycophytes, possibly Selaginellaceae (Schulz, 1967), and may represent a pioneer plant (Van Konijnenburg-Van Cittert, pers. comm.). Heliosporites is also reported from a Rhaeto-Liassic flora from a lacustrine environment in Airel, northern France (Muir and Van Konijnenburg-Van Cittert, 1970). Almost the entire palynomorph assemblage from this locality consists of Classopollis harrisii sp. nov. (>99%). This assemblage is comparable to the pattern in St. Audrie’s Bay. One could argue that this H. reissingeri peak was caused by a sea-level fall resulting in an extension of the coastal area to be invaded by these lycophytes. However, this peak occurs within a transgressive systems tract, suggesting that the increased H. reissingeri abundance reflects a vegetation change induced by climate and not by sea level. Although the lowest total palynomorph concentrations are present within the Lilstock Formation, most samples are characterized by a higher concentration of aquatic palynomorphs which strengthens the idea that the Lilstock Formation was at least partially marine most of the time (Figs. 2 and 4). Furthermore, the spore peaks do not correspond with a lower abundance (both % and concentration) of aquatic palynomorphs. In all samples with spore peaks dinoflagellates are also present. Peak 2 even coincides with a very high abundance of Rhaetogonyaulax rhaetica, which is thought to have been more adapted to open marine conditions (Courtinat and Piriou, 2002; Kürschner et al., 2007). Apart from spore peak 3, there is no direct evidence that the increased amount of spores coincided with a lower sea level. Terrestrial palynomorph associations from the Blue Lias Formation are dominated by C. meyeriana (75–100%). End-Triassic-basal Jurassic Classopollis meyeriana-dominated palynofloras are documented from the continental Newark Basin, USA (Fowell et al., 1994; Olsen et al., 2002a), the Argana Basin, Morocco (Whiteside et al., 2007), and shallow marine/ coastal successions in northern and eastern Spain (Barrón et al., 2006; Gómez et al., 2007). These occurrences imply a change to warmer and/or more arid climate. However, a monotonous Classopollis assemblage is absent from the earliest Jurassic of the Danish Subbasin (Dybkjær, 1991) which suggests that the relatively drier climate was restricted to the interior of Pangaea. Cheirolepidiaceae are most probably wind pollinators (Ziaja, 2006) which produced a large amount of pollen produced per plant and the abundance of Classopollis might be amplified. Such small pollen are easily transported so that sea-level change (transgression) could well have affected Classopollis dominance. Although part of the Cheirolepidiaceous group may have had a coastal habitat (Batten, 1974; Watson, 1988; Abbink, 1998; Abbink et al., 2004), the interval of increased Classopollis meyeriana abundance in the Blue Lias Forma- tion of St Audrie’s Bay is unlikely to be related to an extension of the coastal plain area as this part of the succession represents a transgression. The sea-level curve constructed by Hesselbo et al. (2004) does not follow the changes in the spore:pollen or terrestrial:aquatic palynomorph ratios (Fig. 2). Therefore, it is more likely that climate, rather than sea level, was the main influence on the palynomorph record.

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Another reason for regarding the spore peaks as climate-related is that it is difficult to explain a sea-level change every 20,000 years, especially because the Late Triassic was a non-glacial interval and glacio-eustasy could not be expected (Satterley, 1996). The larger cycles in palynomorph concentrations are also hard to explain by sea-level change occurring every 100,000 years. The changes in terrestrial and aquatic palynomorph concentrations are simultaneous (Fig. 2) which may imply that marine productivity was controlled by enhanced nutrient supply via river runoff.

5.4 The end-Triassic spore spike

It has been suggested that the palynological records from the Newark Basin and St. Audrie’s Bay, with an upward increase in spore diversity and abundance, followed by low diversity assemblages dominated by Classopollis meyeriana, represent comparable microfloral turnovers (e.g., Hesselbo et al., 2002; Hounslow et al., 2004; Whiteside et al., 2007). However, correlation of these records is equivocal. The species composition of the spore spike (or ‘fern spike’) in the Newark Basin is different from that from St. Audrie’s Bay (Table 1) and it is the only one documented in a continental succession (Olsen et al., 2002b), which could indicate local climate changes. The Newark Basin spore peak coincides with a 60% regional extinction of palynoflora and occurs in a coal/smectite clay (Fowell and Olsen, 1993; Olsen et al., 2002b), deposited in a swamp environment where higher spore abundance would be expected. The fern spike in the Newark Basin has been linked to an impact, based on a small iridium anomaly and the occurrence of this spike ~20 kyr before the extrusion of the Orange Mountain basalt (Olsen et al., 2002a). Whiteside et al. (2007) suggested that there is no causal relationship between the T–J extinction and the onset of CAMP volcanism. However, there is ongoing discussion about the exact position of the T–J boundary in the eastern North America Basins (Hounslow et al., 2004; Kozur and Weems, 2005; Lucas and Tanner, 2007b). According to Kozur and Weems (2005) there are no age-diagnostic palynomorphs or other fossils to prove that the extinction and replacement of the diverse P. densus microflora, of taxa with Norian to Rhaetian, or even longer, ranges, such as Enzonalasporites vigens, Carnisporites spiniger, Patinasporite densus, Vallasporites ignacii, Granuloperculatipollis rudis, Classopollis meyeriana and Classopollis torosus) by the low diversity C. meyeriana palynoflora occurred at the T–J boundary. Palynological assemblages from sedimentary rocks just above the North Mountain basalt in the Fundy Basin are dominated by bisaccates such as Lunatisporites rhaeticus and Alisporites parvus and appear to be Triassic in age, indicating that CAMP volcanism may have triggered the T–J environmental crisis (Cirilli et al., 2009). Although spore peaks from the Newark Basin and St. Audrie’s Bay may be contemporaneous, they might not have a causal relationship. Fowell et al. (1994) suggested that the T–J boundary spore spike is analogous to that at the Cretaceous/Tertiary (K-T) boundary. However, in contrast with the T–J boundary, the spore spike just above the level of the extinction of Cretaceous pollen (65 Myr) has a global

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occurrence (Nichols and Johnson, 2008). It occurs at 40 different localities in four basins in the USA, two basins in Canada, one basin in Japan and two basins in New Zealand and in most places consists of a single taxon such as Cyathidites or Laevigatosporites (Nichols and Johnson, 2008 and references therein). The K-T boundary spore spike was linked to an extraterrestrial impact scenario. An impact site is identified, and anomalous iridium concentrations and impact-sourced shock-metamorphosed mineral grains occur in numerous K-T boundary sections. After the impact, ferns took temporary advantage of the absence of seed plants and dominated the landscape as pioneer communities (e.g., Tschudy et al., 1984). Fleming and Nichols (1990) formally defined the spike as a palynological assemblage composed of 70-100% fern spores of a single species occurring within an interval 0-15 cm above the K-T boundary. The reported end-Triassic spore spikes from the Newark Basin (Fowell and Olsen, 1993; Fowell et al., 1994), St. Audrie’s Bay, and the Northwest Germanic Basin (Van de Schootbrugge et al., 2009) all consist of various species (Table 1). The transitional palynomorph assemblages in St. Audrie’s Bay (SAB2 and SAB3) even include four distinct peaks (Fig. 4), with very low spore concentrations. These characteristics are inconsistent with a scenario of pioneer plants invading bare lands after a mass extinction. Therefore, the end- Triassic spore spike is not comparable with that at the K-T boundary and was probably not caused by an impact. This is further indicated by the lack of definite evidence of an impact scenario, such as shocked quartz or an impact crater (Lucas and Tanner, 2007b; Simms, 2007). Gymnosperm forests in northwest Europe were transiently replaced by ferns and fern-associated vegetation (Van de Schootbrugge et al., 2009). This has been interpreted as a pioneer assemblage commonly found in disturbed ecosystems, in this case suggested to be caused by global warming and the release of pollutants during CAMP flood basalt volcanism. It is important to note that these assemblages in Germany are derived from the so-called Triletes Beds, named after the regular occurrence of megaspores which have not been transported over long distances. Hence the increase of fern spores could also be facies-induced. All the locations mentioned by Van de Schootbrugge et al. (2009) that are marked by fern proliferation were at about the same palaeolatitude, suggesting that they may have been within the same humid climate belt. For example, a spore increase in the Tatra Mountains, Slovakia, is interpreted as reflecting a sudden increase in humidity (Ruckwied and Götz, 2009). Volcanic activity associated with changes in oceanic and atmospheric circulation patterns could regionally result in increasing precipitation and/or humidity. Ruckwied et al. (2008) also related an increase of trilete spores within the T–J boundary interval in a terrestrial coal-bearing series in Hungary to increasing humidity, rather than catastrophic events. A spore spike (mainly Concavisporites and Deltoidospora) in the marine Csővár section (Hungary) has been linked to CAMP volcanism (Götz et al., 2009) but the maximum spore abundance is only ~35% at its (short) peak abundance. Van de Schootbrugge et al. (2009) also mention a volcanism-induced fern proliferation in northern Spain (Gómez et al., 2007), but this interpretation is debatable as the very high percentages of Classopollis were represented

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separately from the other palynomorphs by Bárron et al. (2006) who interpreted the fern increase as due to a short humid event in an arid desert-like palaeoenvironment at the beginning of the Jurassic. After this humid event, ‘plant communities were reduced to Cheirolepid- iaceous formations with undergrowth containing scarce lycophytes and ferns’ (Bárron et al., 2006). It was even suggested that peaks of relatively greater abundance of hygrophytic plants might reflect the strong monsoon precipitation regime that dominated in Pangaea during the Late Triassic (Gómez et al., 2007). According to Van de Schootbrugge et al. (2009) Polypodiis- porites polymicroforatus dominates T–J boundary assemblages in Austria. However, though it occurs in the palynoflora of the end-Triassic Schattwald beds it forms only ~10-20% of the terrestrial palynomorph assemblage and occurs with a wide variety of other taxa (Bonis et al., 2009a, p.13-16, fig. 4); a gradual proliferation of ferns in the Austrian sections is here interpreted as reflecting a change to a more humid climate. A fern spike has, according to Van de Schootbrugge et al. (2009), also been documented in continental successions in Greenland. However, the high relative proportion of fern taxa there comprises macrofossils (McElwain et al., 2007) and at present there is no published high-resolution palynological study from Greenland. A taphonomic artefact cannot be excluded in this instance, as this material was deposited in a coal swamp and ‘it is not unusual for peat-forming vegetation to contain a high proportion of fern taxa’ (McElwain et al., 2007). Limited quantitative information from North China suggests that a change from gymnosperm dominance to fern dominance can be recognized as far east as the Junggar Basin. There Late Triassic assemblages containing a wide variety of gymnosperm pollen, but typically without Classopollis, are succeeded by one with nearly 60% fern spores (Lu and Deng, 2005). We consider that linking increased spore abundance to volcanism and pollutants is pre-mature and that the reality of a supra-regional end-Triassic ‘spore spike’ has to await confirmation by more high-resolution palynological records. Possibly, local depositional environment and climate changes influenced fern abundances to a greater degree than suggested by Van de Schootbrugge et al. (2009).

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6. Concluding remarks

High-resolution palynological study of the St. Audrie’s Bay section revealed that there are two transitional zones within each of the palynofloras as recognized by Orbell (1973), and that these contain four pronounced end-Triassic spore peaks. The present study shows there is no single unambiguous global end-Triassic spore spike, but that there is a more complex pattern of spore distribution. Cyclic patterns are observed in the palynomorph records. It is unlikely that the spore peaks from the St. Audrie’s Bay section would show a cyclic pattern if they were linked to fern proliferation after major mass extinction caused by an impact or volcanic activity. We suggest that the spore peaks in the St. Audrie’s Bay section are related to precession-induced changes in monsoon strength and precipitation and/or humidity. This implies that the total duration of the spore peak interval was about 80-100 kyr and that the initial carbon isotope excursion lasted about 20 kyr.

Acknowledgements

We acknowledge funding from the “High Potential” stimulation program of Utrecht Univer- sity. We are grateful to S. Hesselbo for providing the samples. J. van Tongeren and N. Welters are acknowledged for their assistance in the laboratory. M. Hounslow is thanked for his guidance in the field. We thank M. Deenen and W. Krijgsman for the useful discussions. We gratefully acknowledge the thoughtful comments by H. Visscher on an earlier version of the manuscript.

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Appendix 1 Alphabetical list of palynomorphs identified in the St. Audrie’s Bay section * encountered only after qualitative analysis, † encountered only in the quantitative analysis Pollen Alisporites diaphanus (Pautsch) Lund 1977 Alisporites radialis (Leschik) Lund 1977 Alisporites robustus Nilsson 1958 Araucariacites australis Cookson 1947 Callialasporites turbatus (Balme) Schulz 1967* Cerebropollenites thiergartii Schulz 1967 Chasmatosporites apertus (Rogalska) Nilsson 1958 Chasmatosporites major Nilsson 1958* Classopollis meyeriana (Klaus) Venkatachala & Góczán 1964 Classopollis murphyi Cornet & Traverse 1975 Classopollis torosus (Reissinger) Klaus 1960, emend. Cornet & Traverse 1975 Cycadopites Wodehouse 1933 Enzonalasporites vigens Leschik 1955 Ephedripites Bolchovitina 1953 ex. Potonié 1958 Eucommiidites granulosus Schulz 1967* Eucommiidites major Schulz 1967* Eucommiidites troedssonii Erdtman 1948 Granuloperculatipollis rudis Venkatachala & Góczán 1964 Lunatisporites rhaeticus (Schulz) Warrington 1974 Ovalipollis pseudoalatus (Thiergart) Schuurman 1976 Perinopollenites elatoides Couper 1958 Pinuspollenites minimus (Couper) Kemp 1970 Platysaccus Naumova 1937 Quadraeculina anellaeformis Maljavkina 1949 Rhaetipollis germanicus Schulz 1967 Triadispora Klaus 1964* Tsugaepollenites pseudomassulae (Maedler) Morbey 1975 Vesicaspora fuscus (Pautsch) Morbey 1975 Vitreisporites bjuvensis Nilsson 1958 Vitreisporites pallidus (Reissinger) Nilsson 1958 Bisaccate indet A* Bisaccate indet B*

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Spores Acanthotriletes varius Nilsson 1958 Annulispora folliculosa (Rogalska) de Jersey 1959* Aratrisporites centratus Leschik 1956* Aratrisporites minimus Schulz 1967 Asseretospora gyrata (Playford & Dettman) Schuurman 1977 Baculatisporites Thomson & Pflug 1953 Calamospora tener (Leschik) Maedler 1964 Camarozonosporites laevigatus Schulz 1967 Carnisporites anteriscus Morbey 1975 Carnisporites lecythus Morbey 1975 Carnisporites leviornatus (Levet-Carette) Morbey 1975 Carnisporites megaspiniger Morbey 1975 Carnisporites spiniger (Leschik) Morbey 1975 Carnisporites Maedler 1964† Cingulizonates rhaeticus (Reinhardt) Schulz 1967* Conbaculatisporites Klaus 1960 Concavisporites Pflug 1953 Converrucosisporites luebbenensis Schulz 1967 Cosmosporites elegans Nilsson 1958 cf. Cyclotriletes Deltoidospora Miner 1935 Densoisporites nejburgii (Schulz) Balme 1970 Densosporites fissus (Reinhardt) Schulz 1967 Foveolatitriletes potoniei Maedler 1964* Foveosporites Balme 1957† Foveosporites multifoveolatus Doering 1965* cf. Guthoerlisporites magnificusBharadwaj 1954† Heliosporites reissingeri (Harris) Muir & Van Konijnenburg - Van Cittert 1970 Ischyosporites variegatus (Couper) Schulz 1967 Kyrtomisporites laevigatus Maedler 1964 Kyrtomisporites speciosus Maedler 1964b Leptolepidites Couper 1953 emend. Schulz 1967 Limbosporites lundbladii Nilsson 1958 Lophotriletes verrucosus Schulz 1967 Lycopodiacidites rhaeticus Schulz 1967 Lycopodiacidites rugulatus (Couper) Schulz 1967 Neochomotriletes triangularis (Bolchovitina) Reinhardt 1962 Nevesisporites bigranulatus (Levet-Carette) Morbey 1975 Paraklukisporites foraminis Maedler 1964b†

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Perinosporites thuringiacus Schulz 1962 Polycingulatisporites bicollateralis (Rogalska) Morbey 1975* Polycingulatisporites triangularis (Bolchovitina) Playford & Dettmann 1965* Polypodiisporites ipsviciensis (de Jersey) Playford & Dettman 1965 Polypodiisporites polymicroforatus (Orlowska-Zwolinska) Lund 1977 Polypodiisporites sp.† Porcellispora longdonensis (Clarke) Scheuring 1970 emend. Morbey 1975 Retitriletes clavatoides (Couper) Doering, Krutzsch, Mai & Schulz 1963 Retitriletes gracilis (Nilsson) Doering, Krutzsch, Mai & Schulz 1963 Ricciisporites tuberculatus Lundblad 1954 Stereisporites australis (Cookson) Schulz 1970 Stereisporites punctatus Schulz 1970 Stereisporites radiatus Schulz 1953 Stereisporites seebergensis Schulz 1966* Stereisporites sp.* Taurocusporites verrucatus Schulz 1967* Tigrisporites microrugulatus Schulz 1967* Todisporites Couper 1958 Trachysporites fuscus Nilsson 1958 Triancoraesporites ancorae (Reinhardt) Schulz 1967* Uvaesporites Döring 1965 Uvaesporites microverrucatus Schulz 1967 Vallasporites ignacii Leschik 1956* Verrucosisporites cheneyi Cornet & Traverse 1975† Verrucosisporites† Zebrasporites interscriptus (Thiergart) Klaus 1960 Zebrasporites laevigatus (Schulz) Schulz 1967 Spore indet A*

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Dinoflagellate cysts Beaumontella langii (Wall) Below 1987 Cleistosphaeridium mojsisovicsii Morbey 1975 Dapcodinium priscum Evitt 1961 emend. Below 1987 Dapcodinium sp.† Hebecysta brevicornuta Bujak & Fisher 1976 Heibergella aculeata Bujak & Fisher 1976 Heibergella asymmetrica Bujak & Fisher 1976 Heibergella sp. A Rhaetogonyaulax rhaetica (Sarjeant) Loeblich & Loeblich 1968 Rhaetogonyaulax wigginsii (Stover & Helby) Lentin & Williams 1989 Suessia swabiana Morbey 1975 emend. Below 1987

Acritarchs Leiofusa jurassica Cookson & Eisenack 1958 Micrhystridium Deflandre 1937 emend. Sarjeant 1967 Veryhachium Deunff 1958

Prasinophytes Cymatiosphaera polypartita Morbey 1975 Cymatiosphaera ‘small’† cf. Leiosphaeridia Eisenack 1958 Pleurozonaria sp. W Lund 2003 Pterospermella Eisenack 1972 Tasmanites Newton 1875 Tytthodiscus cf. faveolus Morbey 1975

Chlorococcales Botryococcus Kuetzing 1849

Pediastrum

Foraminiferal test linings

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Changing CO2 conditions during the end-Triassic inferred from stomatal frequency analysis on Lepidopteris ottonis (Goeppert) Schimper and Ginkgoites taeniatus (Braun) Harris

End-Triassic fluctuations in atmospheric carbon dioxide (CO2) concentration are reconstructed by the use of stomatal frequency analysis on a single plant species: the seedfern Lepidopteris ottonis (Goeppert) Schimper. The stomatal index shows no significant intra- and interpinnule variation which makes it a suitable proxy for past relative

CO2 changes. Records of decreasing stomatal index and density from the bottom to the top of the Wüstenwelsberg section (Bavaria, Germany)

indicate rising CO2 levels during the Triassic-Jurassic transition. Addition- ally, stomatal frequency data of fossil ginkgoalean leaves (Ginkgoites

taeniatus (Braun) Harris) suggest a palaeoatmospheric CO2 concentration of 2750 ppmv for the latest Triassic.

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1. Introduction

The transition from the Triassic to the Jurassic period is characterized by large perturbations in organic carbon-isotope records (e.g., Pálfy et al., 2001; Hesselbo et al., 2002; Ruhl et al., 2009). These perturbations are related to changes in the global carbon cycle, and have often been linked to massive volcanic CO2 input associated with the Central Atlantic Magmatic Province (CAMP), and/or methane hydrate release (Beerling and Berner, 2002; Hesselbo et al., 2002, 2007; Jenkyns, 2003). Additionally, degassing of organic-rich shales and petroleum bearing evaporites as a result of widespread sill intrusion could have led to greenhouse gas and halocarbon generation in sufficient volumes to cause a negative carbon-isotope excursion (Svensen et al., 2009). The stomatal density (SD) and stomatal index (SI) of leaf cuticles are both inversely related to atmospheric CO2 concentration during growth (e.g., Woodward, 1987; Kürschner et al. 1998) and can be used to reconstruct palaeoatmospheric CO2 levels on short times scales (Beerling, 1993; McElwain et al., 1995; Wagner et al., 1999) and for deep time (e.g., Van der Burgh et al., 1993; McElwain et al., 1999, Royer et al., 2001a; Royer, 2001; Beerling, 2002b; Beerling et al.,

2002; Retallack, 2001, 2009; Kürschner et al., 2008). A fourfold increase (from 600 to 2100-

2400 ppmv) of atmospheric CO2 across the Triassic-Jurassic boundary was suggested based on stomatal frequency analysis (McElwain et al., 1999). On the other hand, carbon isotope compositions of pedogenic calcite from palaeosol formations indicate a relative stability of atmospheric

CO2 (a rise of 250 ppmv) across the boundary (Tanner et al., 2001). These contrasting results could have been caused by differences in temporal resolution (Beerling, 2002a). Furthermore, the suitability of the pedogenic-isotopic palaeobarometer may be questioned because of the compromising effect of massive dissociation events from methane hydrate reservoirs (Retallack,

2002a). The advantage of the stomatal frequency records over other CO2 proxies is that the SI of leaves responds immediately to CO2 change (e.g., Wagner et al., 1996), which makes this parameter ideal for detecting rapid fluctuations (e.g., Wagner et al., 1999; Royer et al., 2001b; Wagner et al., 2005). The study by McElwain et al. (1999) was based on a mixed selection of ginkgoalean and cycadalean species. Because the SI is species specific (e.g., Kürschner et al., 1996; Royer, 2001; Roth-Nebelsick, 2005; Retallack, 2009) it would be advantageous to examine a section with the same species occurring in several stratigraphic layers.

In order to detect changes in the end-Triassic atmospheric CO2 regime, in this study we present a stomatal frequency analysis of a temporal leaf record of a single plant species, Lepidopteris ottonis (Goeppert) Schimper. Leaves of this seed fern were collected from a series of successional organic-rich layers exposed in the Wüstenwelsberg quarry (Bavaria, southern Germany). The intra- and interpinnule variability of the stomatal frequency parameters will be checked before using this species as a proxy for relative CO2 changes. The uppermost macro-fossil layer contains both L. ottonis and Ginkgoites taeniatus (Braun) Harris. By using modern Ginkgo biloba as ‘nearest living equivalent’, stomatal frequency data of the latter species enable a quantitative CO2 estimate for the latest Triassic.

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2. Material and methods 2.1 The Wüstenwelsberg quarry

The studied section is located in a quarry near the village of Wüstenwelsberg (Fig. 1). The sediments were deposited in the Germanic Basin and are characterized by an alternation of clay and sandstone layers (Fig. 2). The stratigraphic sequence was sampled at three different locations within a few meters distance from 6˚ 8˚ 10˚ 12˚ 14˚ each other. The base of the each location is km indicated with a 0 in the lithological column 0 50 100 54˚ 54˚ (Fig. 2). The lowermost part of the section Hamburg consists of an organic rich clay layer (1) with very abundant Lepidopteris ottonis leaves. This Berlin Hannover layer is followed by a thick sandstone (the Utrecht 52˚ 52˚ ‘Hauptsandstein’), which probably belongs Germany to the Postera beds (Ruhl and Kürschner, in Koln prep.). On top of the sandstone lies a ~1m

Coburg Prague thick interval with alternating clay and 50˚ Wüstenwelsberg 50˚ quarry Luxembourg Nurnberg sandstone layers. Some L. ottonis specimens were found in this interval (1a). The consecutive sandstone (~1.5 m thick) is followed by a Munich 48˚ 48˚ Salzburg 10 m thick clay interval (’Hauptton’) which

Dijon Innsbruck represents the middle Rhaetian Contorta Bern beds (Ruhl and Kürschner, in prep.). The Bolzano

46˚ 46˚ clay interval contains several fossil layers 6˚ 8˚ 10˚ 12˚ 14˚ (k2c, wz58, wz53, k2a, k2o, wz57 and k2b, Figure 1: Present-day location of the studied section, the together indicated as level 2, Fig. 2b) with a coordinates of the Wüstenwelsberg quarry are 50°08’N/10°48’E. diverse macroflora, includingL. ottonis. Occasionally, coal seams and large pyrite

Figure 2 (right, color version on p. 211): a) Lithology of the Wüstenwelsberg section, position of the fossil levels and position of the palynology samples. For practical reason, three different locations within a few meters distance from each other were chosen for sampling. The base of each location is indicated with a 0 in the lithological column. The asterisk indicates the position where a hole was excavated to sample the lowermost clay layer with abundant Lepidopteris ottonis leaves. b) Detail of level 2 with the position of all the layers.

118 CHAPTER 6 B A * k2b wz57 k2o k2a wz53 wz58 k2c level 2 detail level

9 8 7 6 ‘Hauptton’ Fine sands Fine sands Medium grey shales Dark grey shales Light grey shales Sandstone

Lithology Fossil leaves levels leaves Fossil 3 2 1 Palynology sample 1a *

4 3 2 1 0 6 5 4 3 2 1 0

6 5 4 3 2 1 9 8 7 0 0 2 1 1 1 1 1 1

Depth (m)

‘Hauptton’ ‘Hauptsandstein’

Triletes beds Triletes Contorta beds Contorta Postera beds? Postera

Jurassic Triassic transition interval transition

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nodules are present. On top of the clays lies another sandstone (~1.5 m thick) followed by the uppermost organic rich clay layer (3) of the section. This interval belongs to the late Rhaetian Triletes beds (Ruhl and Kürschner, in prep.) and contains a more diverse macroflora than the Hauptton. Ginkgoites taeniatus is by far the most dominant species and rare L. ottonis leaves have been found.

2.2 Description of the fossil leaf material

2.2.1 Lepidopteris ottonis (Goeppert) Schimper 1869

Lepidopteris ottonis belongs to the Peltaspermales (Pteridospermopsida, Gymnospermae) and was morphologically and taxonomically described from Sweden (Antevs, 1914; Lundblad, 1950), East Greenland (Harris, 1926; 1932), Poland (Barbacka, 1991), and Germany (Schim- per, 1869; Kelber and Van Konijnenburg-Van Cittert, 1997). Figure 3a shows a L. ottonis specimen from the lowermost fossil level in the Wüstenwelsberg quarry. In the scope of this study we will only focus on the main characteristics of the cuticle (Fig. 3c and 3e). The upper leaf surface is thicker than the lower leaf surface. Cuticularisation is strong, but the degree of cuticularisation varies between different pinnule cuticles. The epidermal cells have a polygo- nal, at times somewhat oblong shape with thick straight walls and occasionally sinuous extensions. Cells are more rectangular over the veins. Papillae in the middle of the epidermal cells are usually present. Stomata are irregularly distributed on both cuticle surfaces (amphi- stomatic), but are far more numerous on the lower one. Occasionally, they are present over the veins. The stomata have 4-7 (mostly irregular) subsidiary cells with papillae on top covering the stomatal aperture and sunken guard cells. Two adjacent stomata are common and even three adjacent stomata have been found. It was already mentioned by Harris (1926) that ‘the great majority of the stomata are in fact somewhat irregular’.

2.2.2 Ginkgoites taeniatus (Braun) Harris 1935

Ginkgoites taeniatus belongs to the Ginkgoales (Gymnospermae) and was morphologically and taxonomically described from East Greenland (Harris, 1935). Figure 3b shows fragments of G. taeniatus leaves from the uppermost fossil level (3) in the Wüstenwelsberg quarry. The cuticle is thicker on the upper side and it shows slightly elongated or isodiametric cells with straight walls (Fig. 3f). Papillae in the middle of the epidermal cells are absent. The cells along the veins are slightly more elongated. The stomata are irregularly distributed over the surface. Subsidiary cells have large papillae overhanging the stomatal pore. The thinner lower cuticle (Fig. 3d) has isodiametric cells between the veins and elongated along them. The epidermal cell walls are nearly straight. Papillae on the epidermal cells are usually absent. Stomata are confined to strips between the veins. The guard cells are sunken and surrounded by about five subsidiary cells which sometimes have overhanging papillae.

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cm cm A B

C D

50 μm 50 μm

E F

50 μm 50 μm

Figure 3 (color version on p. 212): a) A Lepidopteris ottonis leaf from level 1. b) Fragments of Ginkgoites taeniatus leaves from level 3. c) abaxial cuticle surface of Lepidopteris ottonis from level 1, slide 1-A-A. d) abaxial cuticle surface of Ginkgoites taeniatus from level 3, slide Gi-3-A. e) adaxial cuticle surface of Lepidopteris ottonis from level 1, slide 1-A-A. f) adaxial cuticle surface of Ginkgoites taeniatus from level 3, slide Gi-3-A

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2.3 Maceration technique

Cuticles were picked directly from the rock surface. They were macerated according to the

standard procedure (e.g., Kerp, 1990) using Schulze’s reagent (30% HNO3 with a few KClO3 crystals) and subsequently treated with 5–10% potassium hydroxide (KOH). Macerated cuticles were rinsed with water. The upper and lower cuticle surface were separated and embedded in glycerine jelly on microscopic slides. The slides are stored in the collection of the Section Palaeoecology, Laboratory of Palaeobotany and Palynology, Utrecht University.

2.4 Stomatal frequency analysis

Stomatal frequency is conventionally expressed in terms of stomatal density (SD) and stomatal index (SI). The SD is the number of stomata per unit leaf area. The SI is the number of stomata expressed as a percentage of the total number of cells. The SI (%) was calculated after Salisbury (1927): SI = [SD/(SD+ED)]*100. ED is the epidermal cell density per unit leaf area. SD and ED were measured on the abaxial (lower) side of the leaf with a magnification of 330x and a field of view as large as possible of 0.1287 mm2. In contrast to the SD, SI expresses stomatal frequency independently of variation in epidermal cell size. Epidermal cell expansion can be caused by e.g. sunlight intensity, water availability, salinity, and soil nutrient deficiency. Therefore, SI is a more sensitive parameter for detecting stomatal frequency changes (e.g., Van der Burgh et al., 1993; Kürschner, 1996; Kürschner et al., 1996; Beerling, 1999; Royer, 2001; Beerling and Royer, 2002a, 2002b). The ‘stomatal ratio method’ was applied to recon-

struct CO2 values. In this method, a ‘nearest living equivalent’ (NLE) is used, defined as an extant species which is, as far as possible, of comparable ecological setting and/or structural similarity to its fossil counterpart (McElwain and Chaloner, 1995). The stomatal ratio (SR) was calculated by dividing the SI of the NLE by the SI of the fossil. McElwain and Chaloner

(1995) standardized the stomatal ratios of Late Carboniferous conifers against CO2 estimates based on a long term C-cycle model (McElwain et al., 1999). With the use of this Carbonifer-

ous standard, 1SR = 2RCO2 = 600 ppm CO2 (McElwain et al., 1999), CO2 values were

calculated. RCO2 is the ratio of atmospheric CO2 concentration estimated from stomatal ratios relative to a preindustrial value of 300 ppm (McElwain et al., 1999).

2.5 Palynology

Twenty-one palynological samples from the Wüstenwelsberg section were selected for palynological analysis. Between 5 and 20 grams of sediment was crushed into small fragments and dried for 24 hours at 60°C. Subsequently, the samples were treated twice alternately with cold HCl (30%) and cold HF (40%) to remove the carbonates and silicates. The residues were

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sieved using a 250 µm and a 15 µm mesh. ZnCl2 was applied to separate the lighter organic material from the heavier mineral particles. The lighter fraction was transferred from the test-tube and sieved once more using a 15 µm mesh. The remaining organic material was mounted on two slides per sample with glycerine jelly. The slides are stored in the collection of the Section Palaeoecology, Laboratory of Palaeobotany and Palynology, Utrecht University.

3. Results 3.1 The age of the Wüstenwelsberg section

The fossil layers in the Wüstenwelsberg quarry contain a diverse end-Triassic flora (Table 1). The upper level has the most diverse flora and is dominated byGinkgoites taeniatus (Table 1). Lepidopteris ottonis, present in all levels, is restricted to the Rhaetian stage (Antevs, 1914; Barbacka, 1991; Kelber and Van Konijnenburg-Van Cittert, 1997). In Greenland, sediments from the Rhaetian and the Lower Liassic are part of the Lepidopteris zone and the Thaumatop- teris zone, respectively (McElwain et al., 1999; McElwain et al., 2007). The family of L. ottonis (Peltaspermaceae) went extinct at the Triassic-Jurassic boundary (e.g., Ash, 1987). The conifer Stachyotaxus elegans (Table 1) is also characteristic of the Rhaetian stage (Kelber and Van Konijnenburg-Van Cittert, 1997). In East Greenland, Ginkgoites taeniatus is characteristic of the lower and middle part of the Thaumatopteris zone (Harris, 1935), lowermost Jurassic. Fifteen out of the twenty-one palynological samples were productive (Fig. 2, 4, see foldout in the back of the thesis). Ricciisporites tuberculatus dominates (>95%) most of the samples and is therefore shown separate from the other pollen and spores (Fig. 4). From the base up to 15 m, the most abundant pollen types are Vitreisporites bjuvensis, Ovalipollis pseudoalatus and Perinopollenites elatoides. Persistent spore types are Todisporites spp., Deltoidospora spp., Calamospora tener and Concavisporites spp. The amount of spores is highest in the darker organic rich layers. Based on the abundance of Vitreisporites bjuvensis, Ovalipollis pseudoalatus and the presence of Rhaetipollis germanicus, Lunatisporites rhaeticus and Rhaetogonyaulax rhaetica the samples till 15 m are assigned to the end-Triassic. In the uppermost palynological sample from Wüstenwelsberg Ricciisporites tuberculatus is almost absent. This sample consists for 50% of pollen with Classopollis meyeriana, Classopollis torosus, Araucariacites australis, Pinuspollenites minimus and unidentified deformed pollen as the most abundant elements. Most likely, the deformed pollen belongs to Classopollis but it is not possible to distinguish between Classopollis meyeriana and Classopollis torosus. Most abundant spore taxa are Deltoidospora spp., Calamospora tener, Trachysporites fuscus and Concavisporites spp. Remarkable is also the presence of 5% aquatic palynomorphs, mainly consisting of the dinoflagellate cystDapcodinium priscum. The higher abundance of Pinuspollenites minimus, Trachysporites fuscus, D. priscum and the absence of Vitreisporites bjuvensis, Ovalipollis pseudoalatus, Rhaetipollis germanicus, Luna-

123 CHAPTER 6 femalefructification Lepidopteris rare frequent at least 2 species broad leavesbroad Remarks bracts Seedferns Coniferales,Palissyaceae Coniferales Coniferales Conifers?Czekanowskiales?or Seedferns Seedferns Seedferns, Peltaspermales Horsetails,Equisetales Seedferns, Corystospermales Seedferns, Peltaspermales Ginkgophyta,Ginkgoales Ferns,Filicales, Diperidaceae Ferns,Filicales, Dipteridaceae Ferns,Filicales, Dipteridaceae Ferns,Filicales, Mationiaceae Ferns,Filicales, Mationiaceae Ferns,Osmundales, Osmundaceae Ginkgophyta,Ginkgoales Cycadophyta,Cycadeoidales/Bennettitales, Williamsoniaceae Cycadophyta,Cycadeoidales/Bennettitales, Williamsoniaceae Cycadophyta,Cycadeoidales/Bennettitales, Williamsoniaceae Cycads,Cycadales Affinity Cycads,Cycadales Cycads,Cycadales Cycads,Cycadales Cycads,Cycadales sp. nilssonii tietzei sp. cf. sp. sp. cf. sp. sp. Species list of macrofossils found in the Wüstenwelsberg section, based on the finds by Stefan Schmeissner, Günter Dütsch and the authors. the and Dütsch Günter Schmeissner, Stefan by finds the Wüstenwelsbergonsection, thebased in foundmacrofossils of Specieslist Rhaphidopteris Rhaphidopteris Pseudoctenis Pseudoctenis cf. cf. Stachyotaxus elegans Stachyotaxus Elatocladus sp. Cones Needles Ctenozamites wolfiana Ptilozamites heeri Pachypteris sp. ottonis Lepidopteris Peltaspermum Equisetites Clathropteris meniscoides Clathropteris brauniana Thaumatopteris angustiloba Phlebopteris muensteri Phlebopteris Cladophlebis spp. Ginkgoites taeniatus Schmeissneria Otozamites ? Dictyophyllum Pterophyllum Anomozamites spp. Level 3 Dorycycadolepis Species Ctenis sp. pterophylloides Nilssonia Nilssonia Table 1:

124 CHAPTER 6 femalefructification Lepidopteris malefructification Lepidopteris frequent leaf morphogenus leaf Seedferns, Peltaspermales Seedferns Coniferales Conifers Seedferns, Peltaspermales Ferns,Filicales, Diperidaceae Ferns,Filicales, Mationiaceae Ferns,Osmundales, Osmundaceae Seedferns, Peltaspermales probably Cycadophytaprobably Lycophytes,Selaginellales Cycads,Cycadeoidales/Bennettitales, Williamsoniaceae Ferns,Filicales, Diperidaceae Ferns,Filicales, Mationiaceae Ferns Horsetails,Equisetales Seedferns, Peltaspermales Seedferns Conifers Seedferns, Peltaspermales Conifers nilssonii nilssonii cf. cf. sp. sp. sp. Antevsia Seedfern Cones Needles Peltaspermum Phlebopteris muensteri Phlebopteris Todites sp. ottonis Lepidopteris Selaginella Taeniopteris sp. Level 2 Anomozamites Dictyophyllum muensteri Phlebopteris Spiropteris sp. Equisetites sp. ottonis Lepidopteris Seedfern Level 1a elegans Stachyotaxus ottonis Lepidopteris Level 1 elegans Stachyotaxus Dictyophyllum

125 CHAPTER 6

tisporites rhaeticus and Rhaetogonyaulax rhaetica suggests an early Jurassic age of this sample. The difference between the lower palynomorph assemblages and the one above fossil level 3 is similar to the difference between Triassic and Jurassic assemblages from the Eiberg Basin in Austria (Kürschner et al., 2007; Bonis et al., 2009a). Despite the presence of Ginkgoites taeniatus, which occurs in the lower Jurassic in East Greenland, it appears that level, 1, 1a, 2 and 3 are all end-Triassic in age based on the palynological record. This is confirmed by the carbon isotope stratigraphy (Ruhl and Kürschner, in prep.)

3.2 Number of counted fields of view

Relatively stable SI values for one Lepidopteris ottonis pinnule were derived after counting 7 fields of view (Fig. 5a) of 0.1287 mm2. An exception is pinnule WZ53-A-A, which still shows gradually decreasing values after 7 fields. Another pinnule from this level (WZ53-B-A) does show stable values after 6 fields. Unfortunately it was not always possible to count a minimum of 7 fields. Parts of the pinnule surface are unreliable for measurements due to the presence of veins or damage of the cuticle. Stable SI values for one Ginkgoites taeniatus leaf were obtained after counting 10 fields of view (Fig. 5b). For both species, the standard deviations of the SI values were low, e.g., a mean of 0.71 for L. ottonis and a mean of 0.57 for G. taeniatus.

3.5 A B 8

7 3.0

6 2.5 SI [%] SI [%]

2.0 4

3 1.5 0 2 4 6 8 10 12 14 16 18 0 2 4 6 8 10 12 14 16 Field of view Field of view 1-A-A 1-B-A 1a-A-A WZ53-A-A WZ53-B-A Gi-3-A Gi-3-B Gi-3-Ca Gi-3-D k2a-C-B WZ57-B-A k2b-D-A 3-A-A

Figure 5 (color version on p.213): a) Cumulative mean SI of Lepidopteris ottonis b) Cumulative mean SI of Ginkgoites taeniatus

126 CHAPTER 6

A 1-A-A SI ED SD Apex 6.32 252.25 17.00 Middle 6.59 241.88 17.13 4 Base 7.27 232.75 18.25 7.63 242 20 6.99 266 20 5.75 6.93 3 213 13 282 21 6.06 6.47 6.67 248 16 217 15 210 15 7.79 2 6.64 213 18 6.25 6.48 281 20 Pinnule width (mm) 255 17 231 16 7.39 238 19 1 6.07 7.02 232 15 225 17 SI 5.80 7.17 276 17 ED SD 246 19

0 1 2 3 4 5 6 7 Pinnule length (mm)

B 3-A-A SI ED SD Apex 4.57 162.83 7.83 6 Middle 4.15 161.33 7.00 4.42 Base 3.58 159.33 6.00 8 173 2.52 5 155 4

4 4.37 175 8 3.03 160 5 2.58 3 3.80 152 6 151 4 4.74 3.90 2.98 181 9 5.06 2 4.47 148 6 Pinnule width (mm) 163 5 5.13 8 4.86 8 3.87 169 9 176 171 148 4.49 9 4.85 149 6 SI 1 7 8 149 3.77 157 ED SD 6 5.00 153 171 9 0 1 2 3 4 5 6 7 8 9 10 11 12 Pinnule length (mm) Figure 6: a) intrapinnule variation of a pinnule from level 1 b) intrapinnule variation of a pinnule from level 3. ED and SD are indicated per field of view.

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3.3 Intrinsic leaf variability Lepidopteris ottonis

It is important to investigate the intra- and interpinnule variation of the ED, SD and SI for the reliability of measurements from smaller pinnule fragments. The ED, SD and SI were measured at different positions on a pinnule from level 1 (Fig. 6a) and level 3 (Fig. 6b). The SI and SD decrease and the ED increases from the base to the apex in pinnule 1-A-A while in pinnule 3-A-A all parameters increase from the base to the apex. Although there are differences in the mean SD, ED and SI for the apex, middle and base of both pinnules this seems to be arbitrary as there is a different pattern in both pinnules. Exceptionally higher or lower SI values occur at all positions of the pinnules. For example, the lowest SI values from pinnule 3-A-A are located at the base edge (2.52) as well as in the middle closer to the midrib (2.98). There is no pattern in fields of view close to the margin or close to the midrib. It seems that the measurements just capture the natural variability in a pinnule. However, to get the most reliable SI for one pinnule the whole pinnule should be counted, including positions on the apex, base and middle part of the pinnule. The interpinnule variability is checked by measuring four different pinnules from the same leaf (Fig. 7a). There is no significant variation in ED, SD or SI between A a b different pinnules from the same leaf (Fig. 7b, c, d). This implies that analysing one c d pinnule per leaf is enough to obtain consis-

tent ED, SD and SI values. B

240 220 200 [n/area]

D 180 E 160 a b c d Pinnule C 14

12 [n/area]

S D 8 a b c d Pinnule Figure 7: a) schematic representation of a measured leaf from 6 level 2 b) interpinnule variation of ED c) interpinnule variation D ] 5 [ %

of SD d) interpinnule variation of SI. Please note that the measurements are from a previous lower resolution study of S I 4

level 2 and therefore the ED, SD and SI values are not included 3 a b c d in Table 2. ED and SD values are indicated per field of view. Pinnule

128 CHAPTER 6

3.4 Temporal stomatal 0 0

frequency trends 2 2 0 0 0

There is some variability in the ED, 2 0

SD and SI between different leaves 0 8 within one level (Table 2, Fig. 8). This 1 ] 0 2 0 6

is probably a reflection of the natural m 1 m variability of a species. More important 0 0 [ n / 4 is that despite this variability, there are 1 ED 0 significant differences in the SI (and SD 0 2 and ED) throughout the record (Fig. 8). 1 0 0

From level 1 to level wz58 there is a 0 1

decreasing trend in the SI and SD. The 0 0 ED decreases from level 1 to level 1a but 8 0 4 remains stable from level 1a to level wz58. 1 The subsequent levels (wz53, k2a and 0 m ean 2 k2o) show higher SI values (Fig. 8). 1

Important to note is that these levels are 0 0 1 ] positioned the closest to the very high 2 m

organic-rich interval (Fig. 2). Level k2o, m 0 m edian 8 which is positioned within the darkest [ n /

shales, shows significantly higher SD and SD 0 6 ED values (Fig. 8). Above level k2o, SI u s

levels decrease towards the uppermost 0 4 level 3. In addition, in level 3 stomatal a e n i t t

frequency analysis was carried out on 0 e s 2

Ginkgoites taeniatus leaves. SD values are 7

lower than for Lepidopteris ottonis and ED i n k g o t values are positioned within the lower G 6 range of the values from L. ottonis. This results in a significantly lower SI (Fig. 8). 5 ] [ % o n i s t SI o t 4 i s e r 3

Figure 8: SI, SD and ED variations through the L e p i d o t

Wüstenwelsberg section. The dashed grey lines indicate 2 3 1 1a k2c that levels below and beneath are separated by a k2b k2o k2a Level wz57 wz53 wz58 sandstone layer.

129 CHAPTER 6

ottonis

Stomatal Lepidopteris Ginkgoitestaeniatus Table 2: frequency analysis data from and fossils found in the section Wüstenwelsberg 0.93 0.68 1.09 0.84 0.78 0.66 0.93 0.55 0.64 0.71 0.57 0.34 0.60 0.67 0.72 0.55 0.86 0.73 1.07 0.79 0.92 0.62 0.65 0.75 0.85 0.53 0.70 0.64 0.79 0.71 SI sd SI 0.81 0.62 4.24 5.21 4.56 4.95 3.92 5.59 5.10 4.94 5.47 5.26 5.35 3.78 6.20 5.70 5.37 6.01 4.65 5.83 6.36 5.48 5.04 6.70 5.02 4.36 6.13 4.44 6.21 4.40 5.30 5.11 SI 5.53 4.49 12.77 9.66 11.75 20.07 13.80 15.28 18.50 10.36 10.28 15.12 13.00 6.95 11.87 20.28 13.24 9.87 21.40 15.58 18.81 13.27 15.33 17.92 8.97 12.95 19.89 10.94 16.32 11.95 15.03 12.27 SD sd 15.95 12.44 54.03 66.06 58.85 77.72 49.22 89.93 69.95 62.18 93.27 67.17 89.87 66.84 117.88 86.60 71.06 63.04 68.01 106.69 90.44 72.86 70.66 135.04 62.18 57.86 123.32 67.03 120.73 62.45 70.26 67.95 SD 81.04 65.67 181.55 71.04 142.88 171.97 220.22 176.97 108.33 107.90 128.12 147.86 94.92 165.38 88.88 225.82 91.76 99.62 204.00 146.25 148.77 73.87 123.84 189.64 59.43 89.46 112.51 99.22 134.32 94.34 105.45 121.56 ED sd 142.19 142.94 1224.39 1200.79 1247.98 1478.65 1200.79 1518.34 1286.84 1193.02 1616.60 1199.13 1587.94 1706.76 1785.00 1420.08 1243.54 983.60 1369.84 1726.12 1331.15 1250.83 1329.74 1882.31 1174.70 1261.67 1884.48 1436.87 1819.48 1349.27 1245.08 1263.74 ED 1382.23 1417.10 19 12 7 4 12 16 7 10 9 14 16 5 12 7 14 9 4 11 11 16 11 16 7 9 15 8 54 17 51 65 n areas 43 16 K2a-E (average) 1a-E-A pinnules)(2 K2a-E-B 1a-D-A K2a-E-A 1a-C (average) WZ-53-D-A K2a-D-A 1a-C-B WZ-53-C-A K2a-C-B K2c-C-A 1-D-A 1a-C-A WZ-53-B-A K2a-B-A K2c-B-A 1-C-C 1a-A-A WZ-53-A-A K2a-A-A 1-A-A K2c-A-A WZ-58-A-A 1-B-A WZ-58-B-A Slide (level-leaf-pinnule) Slide 115 707 720 below SST 647 652 Depth (cm) Average 1 1a WZ-53 K2A 1 K2C WZ-58 Level Average 1a Average K2C Average WZ-58 Average WZ-53 Average K2A Lepidopteris ottonis Lepidopteris

130 CHAPTER 6 0.47 0.36 0.51 0.84 0.58 0.56 0.90 0.70 0.58 0.38 1.17 0.67 0.72 0.74 0.84 0.34 0.74 0.51 0.79 0.66 0.42 0.85 0.47 0.97 0.71 0.75 0.73 0.67 sd SI 0.57 4.31 4.10 6.15 4.75 4.51 5.70 5.60 4.87 2.43 5.71 4.69 5.23 2.45 5.70 5.03 4.81 4.50 2.61 5.10 5.12 4.51 4.10 2.40 5.02 5.49 5.07 4.87 4.43 SI 2.47 9.67 7.58 12.58 18.68 11.75 12.54 16.83 16.77 15.46 7.55 16.32 15.92 11.99 17.54 16.47 5.88 17.63 9.38 18.92 13.58 6.50 13.45 10.39 17.38 15.36 15.03 12.84 13.30 SD sd 11.80 66.06 58.85 119.91 73.40 73.28 87.08 84.79 90.24 49.96 89.38 55.38 102.76 39.42 84.79 69.24 59.96 77.72 46.08 79.66 72.78 55.96 53.97 43.86 80.31 91.74 71.01 65.90 70.09 SD 44.83 73.63 74.12 189.45 119.09 73.15 126.98 73.61 166.33 177.91 112.69 75.12 148.31 93.62 141.28 129.45 77.81 184.36 112.44 141.15 98.02 32.05 87.54 143.00 112.43 142.50 113.74 75.53 109.17 ED sd 131.74 1458.38 1372.33 1833.11 1456.84 1544.43 1436.18 1423.00 1749.64 1985.78 1479.03 1114.33 1858.40 1571.63 1393.33 1292.29 1186.91 1640.89 1719.86 1469.90 1339.63 1181.36 1252.61 1774.26 1519.02 1564.55 1315.96 1272.49 1486.88 ED 1762.88 14 7 7 9 7 21 11 17 14 10 8 9 14 11 11 7 8 14 8 11 5 18 14 40 9 45 22 49 56 n areas 3-D (average) 3-D-B K2o-D-A K2b-E-A 3-D-A K2o-C (average) K2b-D-A 3-B/C (sameleaf. average) Gi-3-D K2o-C-B K2b-C-A 3-C-A Gi-3-Ca K2o-C-A WZ-57-B-A K2b-B-A 3-B-A Gi-3-B K2o-B-A WZ-57-A-A K2b-A-A 3-A-A Gi-3-A

K2o-A-A Slide (level-leaf-pinnule) Slide 753 763 1480 1480 744 Depth (cm) WZ-57 K2B 3 3 K2O Average K2O Average WZ-57 Average K2B Average 3 Average 3 Level Ginkgoites taeniatus

131 CHAPTER 6

4. Discussion 4.1 Changes in the stomatal frequency analysis parameters

The intra- and interpinnule variation was checked before using the stomatal frequency of L.

ottonis as a proxy for palaeoatmospheric CO2. The intrinsic variability of L. ottonis pinnules is very low (Figs. 6 and 7). The SI of L. ottonis from Scania also showed almost identical values for different parts of the pinnules. Furthermore, the SI within a complete frond or frond

portions is fairly stable justifying the use of SI values from L. ottonis as a proxy for CO2 if this

species is sensitive to CO2 changes (Vörding, 2008). An overall decreasing trend in the SI and SD of Lepidopteris ottonis is present from the bottom to the top of the studied section (Fig. 8).

SI is inversely related to the ambient CO2 concentration during growth and can therefore be

used to reconstruct palaeoatmospheric CO2 (e.g., Woodward, 1987; Kürschner, 1996; Kürschner et al., 1996; McElwain et al., 1999; Royer, 2001; Retallack, 2001, 2009). To exclude the effect of changing ED as the main driver of SI (instead of a higher SD induced

by higher CO2 concentration) two SI simulations were carried out (Fig. 9a). The first one was calculating the SI with a constant (mean) ED value of 1410 n/mm2. The second one was calculating the SI with a constant (mean) SD value of 78 n/mm2. The simulated SI signal with constant ED values shows the same (amplified) pattern as the original SI record (Fig. 9a). This suggests that the changes in SD were indeed the main driver of changes in the SI. The uppermost SI value could be partly influenced by the strong increase in ED with respect to SD (Fig. 8). The correlation between SD and ED is high (R2= 0.74) (Fig. 9b). This implies that if

solely the SD was used to infer relative atmospheric CO2 changes, the reconstructed CO2 changes might be partly disturbed by ED (and consequently SD) changes related to changes in epidermal cell expansion caused by e.g. sunlight intensity, water availability, salinity, and

soil nutrient deficiency. SI is preferable for a CO2 reconstruction as there is a very low correlation (R 2= 0.14) between SI and ED (Fig. 9d) and a higher correlation between SI and SD (R2= 0.63) (Fig. 9c). The lack of correlation between SI and ED implies that changes in SD rather than epidermal cell expansion was the cause for changes in the SI. Level k2o is positioned within the very dark organic-rich shales and it shows significantly higher SD and ED values (Fig. 8). Experimental leaves grown under high humidity levels show a lower SD and ED, caused by an increased cell expansion (Kürschner, 1996). Stomata can be packed more densely because of drought (e.g., Woodward, 1987; Royer, 2001). The high SD and ED in level k2o are probably the result of adaptation to drought stress. All the leaves from this level also show strongly cutinized cell walls and papillae in the middle of the epidermal cells. Based on the strong cuticularisation it was suggested by Antevs (1914) that L. ottonis is a succulent. In contrast, extraordinary thick cuticles are reported from Lepidopteris callipteroides, an early Triassic seed fern, but which lived in a humid environment (Retallack, 2002b). Another pronounced feature of L. ottonis are the papillae covering the stomata,

132 CHAPTER 6

Level 3 140 1 1a k2c w z58 SI w z53 k2a k2o w z57 k2b 130 k2b 3 SI simulation 1 constant ED wz57 120 SI simulation 2 k2o constant SD 110

k2a ] 100 2 m m wz53 / 90 [ n

wz58 S 80

k2c 70

1a 60

1 50 A B R2 = 0.7364 40 3 4 5 6 7 8 800 1000 1200 1400 1600 1800 2000 SI [%] ED [n/mm 2]

7.0 7.0

6.5 6.5

6.0 6.0

5.5 5.5 ] ] [ % [ %

I 5.0

I 5.0 S S

4.5 4.5

4.0 4.0

3.5 3.5 2 C R = 0.6257 D R2 = 0.1416 3.0 3.0 50 70 90 110 130 150 800 1000 1200 1400 1600 1800 2000 SD [n/mm2] ED [n/mm 2]

Figure 9 (color version on p.214): a) simulated SI values with a constant ED and a constant SD compared to the original SI values b) correlation between SD and ED for each leaf within a fossil level c) correlation between SI and SD for each leaf within a fossil level d) correlation between SI and ED for each leaf within a fossil level

133 CHAPTER 6

speculated to be a mechanism for reducing transpiration in drought-prone soils (Retallack, 2002b). However, papillae overarching the stomata, sunken stomata and other xeromorph characteristics may also indicate adaptation to stress factors other than just a dry climate (Haworth and McElwain, 2008; Pott et al., 2008). One suggestion is that the papillae were helpful in reducing transpiration in soils low in nutrients (peinomorphic adaptation) (Retal- lack, 2002b). Physiological drought may also be caused by strongly permeable or osmotic soils, wind exposure, saline environments, low pH values, or an epiphytic way of life (Pott et al., 2008). McElwain et al. (2007) suggested that L. ottonis had a vine- or liana-like way of living based on similar anatomical and morphological characteristics common to among modern angiosperm vines and lianas (Krings et al., 2003). We did not find indications to confirm this interpretation. We propose thatL . ottonis from Wüstenwelsberg had these xeromorph features because it lived in a humid and coal-producing swampy environment low in nutrients, rather than a dry environment. The presence of papillae can be a way of prevent- ing the formation of a water film on the leaf surfaces, which can adversely affect the uptake of

CO2 (Pott et al., 2008; Haworth and McElwain, 2008). Additionally, papillae can be helpful in the self-cleaning effect of the leaves with regard to dust, or atmospheric particles caused by local volcanism (Pott et al., 2008; Haworth and McElwain, 2008). A late Carnian flora from Lunz also showed adaptations (sunken stomata, papillate surfaces) to ecological conditions characteristic of peat swamps, e.g., physiological drought (Pott et al., 2008). The palynological content of the organic rich samples confirms our interpretation of a swampy environment as fern spores are the major constituent of the assemblages.

4.2 Reconstruction of palaeoatmospheric CO2

A comparison of our data with those published by Retallack (2001, 2009) appears to be of limited use as the age assessment of his data is rather vague. Moreover, stomatal counts have been made from previously published photographs which may be biased because of selection for good visibility of the epidermal cell patterns but are not necessarily representative for the stomatal density patterns of the fossil leaf material. The SI of L. ottonis ranges between 6.3% (± 1.9) and 8.7%, comparable to the youngest sample from Wüstenwelsberg. The other SI values from Wüstenwelsberg are significantly lower than Retallack’s values, which is also the case for G. taeniatus with 2.47% versus 7.2% ± 0.5. If we apply our SI of G. taeniatus to the

transfer functions based on ginkgoalean leaves obtained by Retallack (2001, 2009) odd CO2 values (respectively 7560 ppmv and 1575256 ppmv) are calculated. This is probably because our SI is outside the range of the calibrated SI values for the transfer functions. The same problem arises when using transfer functions proposed by Royer et al. (2001a), Beerling and Royer (2002a), and Royer (2003). Furthermore, these functions are more useful for the Cenozoic species, such as G. occidentalis and G. biloba (Royer, pers. comm.). For these reasons we suggest that the stomatal ratio method (e.g., McElwain and Chaloner, 1995; McElwain et

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Table 3: a) Corrected SI of Lepidopteris ottonis (SI ‘Ginkgoites’) used to reconstruct CO2 values with the stomatal ratio method. SI ‘Ginkgoites’ = SI Lepidopteris – 1.95. SR = SI NLE/SI fossil. SI NLE =

11.33%. CO2 = SR*600. Level k2o is positioned within the dark shale layer. b) corrected CO2 values based on the changes in SD. c) SI/SD based CO2 values

A Level SI Lepidopteris SI Ginkgoites SI ‘Ginkgoites’ SR CO2 3 4.43 2.47 4.58 2749.19 k2b 4.87 2.92 3.88 2329.89 wz57 5.07 3.12 3.63 2177.39 k2o 5.49 3.54 3.20 1920.66 k2a 5.11 3.16 3.58 2150.35 wz53 5.30 3.35 3.38 2029.19 wz58 4.40 2.45 4.63 2777.85 k2c 4.49 2.53 4.47 2684.38 1a 5.53 3.57 3.17 1901.67 1 6.21 4.26 2.66 1595.49

B Interval ΔSD ΔCO2 % SD change ΔCO2new 1 to wz58 58.28 1182.36 wz58 to wz53 7.81 748.66 13.41 158.53 wz58 to k2a 5.51 627.50 9.45 111.73 wz58 to k2o 29.30 857.19 50.26 594.30 wz58 to wz57 8.56 600.46 14.69 173.69 wz58 to k2b 3.45 447.97 5.92 69.98 wz58 to 3 7.65 28.66 13.12 155.11

C Level CO2 new 3 2622.74 k2b 2707.87 wz57 2604.16 k2o 2183.55 k2a 2666.12 wz53 2619.33 wz58 2777.85 k2c 2684.38 1a 1901.67 1 1595.49

135 CHAPTER 6

al., 1999) is the most appropriate for our fossil data.

Ginkgoalean leaves proved to be suitable for CO2 reconstruction studies (e.g., McElwain et al., 1999; Chen et al., 2001; Retallack, 2001; Beerling and Royer, 2002b; Sun et al., 2008). The stomatal ratio (SR) was calculated by dividing the SI of the NLE by the SI of the fossil (Table 3a). We use Ginkgo biloba as the NLE of G. taeniatus with an SI of 11.33% for a preindustrial

Depth (m) Fossil leaves level 3

13 Triletes beds Triletes

8 k2b wz57 k2o k2a 7 wz53 Contorta beds k2cwz58 6 Triassic 5

1 1a

0 2

Postera beds? SI based CO2 reconstruction 1 1 SI/SD based CO2 reconstruction

0 1500 2000 2500 3000

CO2 [ppmv]

Figure 10 (color version on p. 215): atmospheric CO2 changes in the Wüstenwelsberg section based on the SI value of Ginkgoites taeniatus from the uppermost level 3 and on the corrected Lepidopteris ottonis SI (SI ‘Ginkgoites’, Table 3a) values from the other levels. The grey band

shows the general CO2 trend.

136 CHAPTER 6

CO2 value of 300 ppm (McElwain et al., 1999). McElwain and Chaloner (1995) standardized the stomatal ratios of Late Carboniferous conifers against CO2 estimates based on a long term C cycle model (McElwain et al., 1999). With the use of this Carboniferous standard, 1SR = 2RCO2

= 600 ppm CO2 (McElwain et al., 1999), a CO2 value of 2750 ppmv was reconstructed for level 3 (Table 3a; Fig. 10). Unfortunately, L. ottonis has no NLE because the order to which it belonged (Peltaspermales) went extinct at the Triassic-Jurassic boundary (e.g., McElwain et al., 2007). However, in level 3 both G. taeniatus and L. ottonis are present together which means that at a

CO2 value of 2750 ppmv L. ottonis had a SI of 4.43 (Table 3a). In order to infer CO2 concentrations from L. ottonis in the lower part of the section we assume that the difference between the SI from G. taeniatus and the SI from L. ottonis (1.95) maintained constant. With this corrected ‘Ginkgoites’ SI (SI ‘Ginkgoites’ = SI Lepidopteris – 1.95) we calculated the relative changes in

CO2 (Table 3a, Fig. 10). In general, there is an increase of ~1150 ppmv (Fig. 10).

An abrupt enigmatic CO2 perturbation seems to occur in the middle of the Contorta beds. One possibility is that this is a short term event with a ~750 ppmv drop in CO2. However, there are no indications for such an event from the carbon isotope record (Ruhl and Kürsch- ner, in prep.) or from other end-Triassic records. The question rises if the SI changes in this interval really reflect CO2 dynamics. Therefore, we checked the trends in SD and we note that in the lower part of the section a decline in SI from 6.2 to 4.4 (ΔCO2 of 1182) is associated with a decline in SD from 121 to 62 (~50%). By contrast, the increase in SI in the Contorta beds from 4.4 to 5.3 (ΔCO2 of 748) is only associated by an increase in SD from 62 to 70 (~10%).

This decline in SD during the CO2 increase is not proportional to the increase in SD during the CO2 decrease of a similar order. The observed disproportion is caused by the changes in ED (Fig. 8) and therefore we assume that the SI changes in this part of the record do not reflect CO2 changes. Because the ED is 10 to 20 times higher than SD, relative small changes in ED can cause large changes in the SI values. This ED sensitivity is a disadvantage when using SI values from L. ottonis as a proxy for CO2 changes. If we focus only on the regression between wz58 and wz53 (Fig. 9d) it is obvious that this would give a different regression line between ED and SI (i.e. ED would increase when SI decreases, dashed oval in Fig. 9d). This confirms that ED is indeed an influencing factor of the SI change between sample wz58 and wz53. To correct for the changes in ED, we reconstructed new CO2 values for samples wz53 to 3 standardized on the changes in SD and CO2 between level 1 and wz58 (Table 3b and c). This combined SI/SD reconstruction (Fig. 10) still shows one outlier: sample k2o from the dark organic rich shales. As mentioned above, it must be taken in mind that these SI fluctuations in the organic rich layer can be influenced by a local stressed swamp environment

(increased SD and ED) rather than a global CO2 signal caused by CAMP volcanism and/or methane hydrate. A study on stomatal responses of modern Polystichum munitum (sword fern) growing near the Kilauea volcano on Hawaii showed that, besides CO2, also volcanogenic SO2 can have a significant effect on leaf stomata (Tanner et al., 2007). However, at present it is not possible to recognize an SO2 effect in deep-time leaf records.

137 CHAPTER 6

The reconstructed end-Triassic CO2 value of ~2750 ppmv from Wüstenwelsberg is slight-

ly higher than the values reconstructed from Greenland and Sweden, where CO2 increased from 600 to 2100-2400 ppmv across the Triassic-Jurassic boundary (McElwain et al., 1999). There is palynological evidence that the uppermost sample from the Wüstenwelsberg section is Jurassic in age, which suggests that level 3 correlates with the transition from the Lepidopt- eris zone to the Thaumatopteris zone (McElwain et al., 1999, McElwain et al., 2007). Our study

shows that already prior to the Triassic-Jurassic transition CO2 values were high (~1600, and increasing to ~2800 ppmv in the Contorta beds) (Fig. 10). This is in line with the carbon isotope study on pedogenic carbonate nodules by Cleveland et al. (2008) which shows in-

creased Rhaetian levels (>1500 ppmv) and at least two periods of extreme CO2 levels (~3000 ppmv) preceding the Triassic-Jurassic boundary. Unfortunately, these North American data lack a detailed chronological resolution so a one-to-one correlation is at present not possible.

5. Conclusions

The extinct seed fern Lepidopteris ottonis appears to be useful in reconstructing relative CO2 changes. The intra-and interpinnule variability of the SI is insignificant. Despite the fact that L. ottonis shows various xeromorphic features, we suggest this is an adaptation to a stressed (swamp) environment rather than just a dry environment. With the use of co-occurring

Ginkgoites taeniatus leaves from the youngest fossil level, an elevated CO2 concentration up to 2750 ppmv was reconstructed for the Triassic-Jurassic boundary interval. In order to infer

CO2 concentrations from L. ottonis in the lower part of the section we made a corrected

‘Ginkgoites’ SI and estimated the relative changes in CO2. A remarkable result is that already

before the latest Triassic elevated atmospheric CO2 concentrations were present, which is in line with a previous carbon-isotope study on pedogenic carbonate nodules. It is important that future work will focus on palaeomagnetic, palynological and chemostratigraphic correlation of the different Triassic-Jurassic boundary sections in- and outside the German Triassic Basin to

get a grip on the global pattern of these end-Triassic high CO2 values.

Acknowledgements

We acknowledge funding from the ‘High Potential’ stimulation program of Utrecht Univer- sity. We are grateful to Stefan Schmeissner and Günter Dütsch for showing us around in the field, providing us with data from the recovered macroflora and sending usLepidopteris ottonis leaves for cuticular analysis. Martijn Deenen and Micha Ruhl are thanked for their help during the field work. We thank Rike Wagner-Cremer for her help with the stomatal frequency analysis and Henk Visscher for constructive comments on an earlier version of the manuscript.

138

CHAPTER 7

Vegetation history, diversity patterns, and climate change across the Triassic-Jurassic boundary

High-resolution palynological datasets from Triassic-Jurassic (T-J) boundary beds of two key sections in Europe (Hochalplgraben, Austria and St. Audrie’s Bay, UK) were analysed to reconstruct changes in vegetation, biodiversity and climate. In Hochalplgraben, a hardwood gymnosperm forest with conifers and seed ferns is replaced by vegetation with dominant ferns, club mosses and liverworts. This goes along with an increased diversification of spore types during the latest Rhaetian. Multivariate statistical analysis reveals a trend to wetter conditions across the T-J boundary in Hochalplgraben. The vegetation changes in St. Audrie’s Bay are markedly different. Here, a mixed gymnosperm forest is replaced by monotonous vegetation consisting mainly of Cheirolepid- iaceae (80-100%). This change is caused by a transition to a warmer and more arid climate. The observed diversity decrease in St. Audrie’s Bay affirms this interpretation. Neither of the sections demonstrates a distinctive floral mass extinction. A compilation of T-J boundary sections across the world demonstrates the presence of Cheirolepidiaceae dominated forests in the Pangaean interior and an increase in spore producing plants near the Tethys Ocean. We propose that the floral differentiation reflected in the T-J palynological record is the indirect result of Central Atlantic Magmatic Province volcanism. The increase in greenhouse gases caused a warmer climate and an enhanced thermal contrast between the continent and the seas. Consequently, the monsoon system got stronger and induced a drier continental interior and more intensive rainfall near the margins of the Tethys Ocean.

140 CHAPTER 7

1. Introduction

The transition from the Triassic to the Jurassic period is characterized by a major biotic crisis in the marine and terrestrial realm (e.g., Hallam, 2002; Olsen et al., 2002; Tanner et al., 2004). However, the severity of this crisis, especially for the terrestrial realm, is disputed (Hallam, 2002; Lucas and Tanner, 2008). Palynological analysis is a useful method in unravelling past vegetation patterns and reconstructing climate change (e.g., Barrón et al., 2006; Galfetti et al., 2007, Chapter 3). Palynological records across the Triassic-Jurassic (T-J) transition, however, are controversially interpreted because of the paucity of sections with a sufficient time resolution and/or well established chronostratigraphic framework. Evidence for a prominent end-Triassic extinction event in the plant fossil record is ambiguous. A major extinction of 60% of sporomorph taxa followed by a sharp spore spike at the T-J boundary is claimed in the Newark Basin, USA (Fowell and Olsen, 1993; Fowell et al., 1994; Olsen et al., 2002). By contrast, most palynological studies from Europe show gradual changes in assemblages at the T-J transition (e.g., Warrington, 1974; Morbey, 1975; Lund, 1977; Schuurman, 1979; Achilles, 1981; Kürschner et al., 2007; Bonis et al., 2009a). Also the T-J plant macrofossil record is equivocal. Quantitative macrobotanical data from East Greenland showed that Triassic forests with high diversity communities were replaced by lower diversity forests and that there was a gradual extinction prior to the T-J boundary (McElwain and Punyasena, 2007; McElwain et al., 2007). Although the late Triassic event in Greenland did not induce mass extinction of plant families, it is accompanied by major and abrupt changes in floral ecology and diversity (McElwain et al., 2009). By contrast, the palynological record from Greenland shows a different pattern, with no major diversity or assemblage changes and no conclusive evidence for an extinction event (Raunsgaard Pedersen and Lund, 1980; Koppelhus, 1997). In order to clarify the controversy, the present paper is aimed at a reconstruction of vegetation history, diversity patterns and inferred climate changes across the T-J transition on the basis of high-resolution quantitative palynological datasets from palaeogeographically contrasting settings in Europe. Comparison of records from Hochalplgraben (Austria) and St. Audrie’s Bay (UK) with each other and with various boundary sections across the world may contribute to better understanding of the nature and rate of T-J vegetation changes, as well as the environmental or climate changes that could have driven this turnover.

141 CHAPTER 7

2. Materials and methods 2.1 Studied sections

The present study concentrates on two European T-J boundary key sections (Fig. 1): Hochalplgraben in the Northern Calcareous Alps, Austria (47°28’20’’N, 11°24’42’’E) and St. Audrie’s Bay, Southwest UK (51°11’N, -10˚ -5˚ 0˚ 5˚ 10˚ 15˚ 20˚ 60˚ 60˚ 3°17’W). A description of the Hochalpl- graben section, together with a high resolution palynological study is presented in Bonis et al. 2009a. Hochalplgraben shows similar 55˚ 55˚ palynological assemblages as the Tiefen- graben section (Kürschner et al., 2007) and

St. Audrie’s Bay the Kuhjoch section (Bonis et al., 2009a). 50˚ 50˚ The latter is recently approved as the Global

Hochalplgraben Stratotype Section and Point (GSSP) for the base of the Jurassic (Von Hillebrandt et al., 45˚ 45˚ 2007). St. Audrie’s Bay is a classic outer- Alpine T-J boundary section (Warrington et 40˚ 40˚ al., 1994; Warrington et al., 2008). Investi-

km gated samples for a detailed palynological 0 200 400 35˚ 35˚ study come from the Westbury Formation, -10˚ -5˚ 0˚ 5˚ 10˚ 15˚ 20˚ the Lilstock Formation (Cotham Member Figure 1: Present day location of the Hochalplgraben and Langport Member) and the lower part (47°28’20’’N, 11°24’42’’E) and St. Audrie’s Bay section of the Blue Lias Formation (Chapter 5). (51°11’N, 3°17’W).

2.2 Vegetation reconstruction

The majority of the pollen and spores of land plants (sporomorphs) from the Hochalplgraben and St. Audrie’s Bay sections may be classified (Appendix 1) in terms of their botanical affinity (Schulz, 1967; Balme, 1995; Abbink, 1998; Hubbard and Boulter, 2000; Herngreen, 2005a,b; Raine et al., 2005; Barrón et al., 2006; Lindström and Erlström, 2006; Ziaja, 2006; Traverse, 2007; Van Konijnenburg-Van Cittert, pers. comm.) These affinities were used to reconstruct vegetation development across the T-J boundary.

142 CHAPTER 7

2.3 Multivariate statistical analysis

A Detrended Correspondence Analysis (DCA) was carried out on the relative abundances of sporomorphs (with a square root transformation of species data and down weighting of rare species) to determine the gradient length of the first axis. The gradient length of the first axis is 2.026 SD (standard deviations) for Hochalplgraben and 2.393 SD for St. Audrie’s Bay. Because the gradient lengths of the datasets did not exceed 3 SD, a linear ordination method, Principal Components Analysis (PCA), was used to make a summary of the relative pollen and spore abundances datasets (Lepš and Šmilauer, 2003). Both PCA’s were done with a square- root transformation of the species data and the data were centred by variables (taxa).

2.4 Diversity analysis

The qualitative sporomorph datasets from Hochalplgraben and St. Audrie’s Bay were used to carry out a diversity analysis with the computer program PAST (Hammer et al., 2001). Data were subjected to the range-through assumption (absences between first and last appearance are treated as presences) and possibly reworked taxa were rejected from the diversity analysis. Additionally, the amount of pollen and spore taxa present per sample (species richness) was calculated. The quantitative datasets were used for a rarefaction analysis (Birks and Line, 1992) standardized on a pollen sum of 207 grains with 95% confidence intervals. This is an intrapolation technique making it possible to estimate how many species would have been found had the sample been smaller than it actually was (Raup, 1975). In this way we can compare estimated diversities at a constant sample size.

3. Results 3.1 Vegetation reconstruction

Vegetation patterns across the T-J boundary were reconstructed on the basis of known botanical affinities of the sporomorphs. It has been demonstrated in studies on modern palynology – vegetation relationships that modern pollen diversity reflects the diversity of the surrounding vegetation and that fossil pollen diversity may provide an important proxy to reconstruct changes in the plant diversity (e.g., Weng et al., 2007; Lézine et al., 2009; Pelánk- ová and Chytrý, 2009). However, dealing with Mesozoic material, one factor that is difficult to include is the variation in sporomorph production by the parent plants. For example, as Cheirolepidiaceae were most probably wind pollinators they produced a large amount of pollen (Classopollis) per individual tree (Alvin, 1982; Ziaja, 2006). However, for the majority

143 CHAPTER 7

of Triassic and Jurassic plant species the way of pollination is not known and can only be inferred from the appearance of the plant. For some sporomorph taxa the botanical affinity is even still unknown because the sporomorphs have never been found in situ (e.g., Trachysporites spp. and Ovalipollis spp.). This should be kept in mind when interpreting changes in vegetation composition based on sporomorph counts. Furthermore, the relative abundance of pollen and spores can be influenced by sea level change (e.g. Abbink et al., 2004). During high sea level one would expect an increase of pollen with a high buoyancy (e.g., bisaccates). Spores would be more abundant at low sea level as they are relatively heavy and more difficult to transport. This is known as the Neves effect (e.g., Traverse, 2007). Influence of sea level changes on the sporomorph composition in St. Audrie’s Bay is described in detail in Chapter 5. Figure 2 shows the relative abundances of plant groups in the Hochalplgraben section. In the Kössen Formation the vegetation can be described as a hardwood gymnosperm forest with conifers and seed ferns. The most dominant conifers are Cheirolepidiaceae, the parent plants of Classopollis pollen. At the transition from the Kössen Formation to the Tiefengraben Formation spore producing plants increase considerably. Most abundant are liverworts, club mosses and different fern types like Dicksoniaceae and Cyatheaceae (both tree ferns), Schizaeaceae (climbing ferns), Matoniaceae and Osmundaceae. Cheirolepidiaceae are decreasing while Caytoniales (seed ferns) on the other hand are increasing. A narrow peak (>90%) of Cheirolepidiaceae is present at 550 cm. Above the Schattwald beds, Cheirolepid- iaceae decrease and the seed ferns almost disappear. Abundant vegetation groups are different fern types and Selaginellales (spike mosses). ‘Trachysporites producing ferns’ comprise a major part within the ‘Other ferns’ group. Above 800 cm (the earliest Jurassic), the vegetation is relatively stable, consisting mainly of liverworts and ferns. Gymnosperms are no longer an abundant element of the vegetation and the relative amount of pollen of Cheirolepidiaceae decreased to values of around 10%. The vegetation changes in St. Audrie’s Bay are quite different from the change described from Hochalplgraben (Fig. 3). Till the top of the Westbury Formation major components are Cheirolepidiaceae, other gymnosperms, and liverworts. The lower part of the Lilstock Formation (Cotham Member) consists of Caytoniales, Taxodiaceae, and an alternating amount of Cheirolepidiaceae and different fern types (e.g., Schizaeaceae, Matoniaceae, Osmundaceae). Around 1250 cm horsetails increase in abundance. The most striking feature in the upper Lilstock Formation (Langport Member) is the acme (>80%) of club mosses (Selaginellales). Alternating fern and Cheirolepidiaceae abundance make up the rest of the Langport Member. The Blue Lias Formation shows a monotonous vegetation consisting of Cheirolepidiaceae (80-100%) combined with club mosses (Selaginellales).

144

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Limestone

Depth (cm) Tiefengraben Member Tiefengraben Eiberg Mb Eiberg

Kendlbach Formation Kendlbach Fm Kössen

Lithology Figure 2: on their botanical affinity (Appendix 1). Jurassic Triassic

145

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146 CHAPTER 7

3.2 Multivariate statistical analysis

The results from Hochalplgraben and St. Audrie’s Bay are displayed as the species scores on the first and second axis of PCA ordination diagrams (Fig. 4). These two main ordination axes are the dimensions through the dataset which explain the largest variance in species composition and can be translated in terms of the environmental and/or climatic gradient that controls the dataset. Figure 4a shows the spread of the sporomorph taxa from Hochalplgraben in the PCA ordination diagram. The first axis explains 43.5% of the total variance within the dataset, the second axis explains 19.7%. On the positive side of the first axis, various pollen taxa have a high score (e.g., Classopollis meyeriana, Classopollis torosus, Vitreisporites pallidus + bjuvensis and Ovalipollis pseudoalatus). Spores like Trachysporites fuscus, Ricciisporites tuberculatus and Heliosporites reissingeri have a high negative score on the first axis. These spores are produced by moisture-loving plants as ferns and liverworts. Cheirolepidiaceae, which produced Classo- pollis pollen, became dominant or even mono-dominant during increasing aridification that limits the areal of the moisture-loving plants (Vakhrameev, 1987). They preferred a subtropical to tropical, somewhat arid climate (Vakhrameev, 1981; 1991), although there are some species which may have had a coastal habitat (e.g. Batten, 1974; Watson; 1988; Abbink, 1998; see also Chapter 3). However, based on the position of the spores relative to the position of Classopollis, the first axis may then be interpreted as a ratio between sporomorph types indicative of relatively wet and relatively dry conditions. C. meyeriana also has a very high score on the positive side of the second axis. Taxa with a negative score on the second axis are C. torosus, R. tuberculatus, Vitreisporites pallidus + bjuvensis, Convolutispora microrugulata and Deltoidospora spp. Cheirolepidiaceous conifers were regarded as thermophillous (Vakhrameev, 1981; Alvin, 1982). Studies from Greenland (Raunsgaard Pedersen and Lund, 1980; Koppel- hus, 1997) and Siberia (Rovnina, 1972) report that Classopollis is rare and if it occurs in the record, it is mostly C. torosus. The occurrence of mainly C. torosus in high latitude records suggests that this pollen type was produced by a Cheirolepidiaceae species better adapted to colder conditions than the parent plant of C. meyeriana. Therefore, the second axis is interpreted to represent a ratio between sporomorph types indicative of relatively cold versus relatively warm conditions. The high negative score on the second axis of Vitreisporites confirms this interpretation, as this taxon is also known from Greenland (Raunsgaard Pedersen and Lund, 1980).

Figure 3 (left): Vegetation history of the St. Audrie’s Bay section. The sporomorphs are classified based on their botanical affinity (Appendix 1). Lithology after Hesselbo et al. (2004).

147 CHAPTER 7

A: Hochalplgraben

7.5 + 6.0 Classopollis meyeriana

4.5

3.0 d (temperature) e n i 1.5 Trachysporites fuscus Heliosporites reissingeri a l x p e

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6.0 4.8 3.6 2.4 1.2 0.0 -1.2 -2.4 -3.6

PCA axis 2, 17.8% explained (humidity) + -

Figure 4: PCA ordination diagram of the pollen and spore taxa in a) Hochalplgraben and b) St. Audrie’s Bay.

The first and second PCA axes from St. Audrie’s Bay explain respectively 46.5% and 17.8% of the total variance within the dataset (Fig. 4b). The position of the taxa in the ordination plot is similar to Hochalplgraben. To facilitate a direct comparison, the axes from

148 CHAPTER 7

St. Audrie’s Bay are reflected (Fig. 4b).C . meyeriana has a high negative score on the first axis, while C. torosus, O. pseudoalatus, Rhaetipollis germanicus and R. tuberculatus score high on the positive side of the first axis. On the second axis, spore taxa likeH . reissingeri, Deltoidospora spp., Conbaculatisporites spp., Trachysporites spp., Todisporites minor+major and Acanthotriletes varius have a positive score, while various pollen taxa (e.g., C. meyeriana, C. torosus, O. pseudoalatus, R. germanicus and Granuloperculatipollis rudis) have a high negative score. The first axis is interpreted as a relative temperature axis and the second as a humidity axis using the same way of reasoning as discussed above for Hochalplgraben.

3.3 Diversity

All three diversity records (range through, species richness and rarefaction) show a similar trend for both sections (Fig. 5). There is an almost fourfold diversity increase in Hochal- plgraben from the Eiberg Member to the top of the Schattwald beds (Fig. 5a). Highest diversity is reached in the Schattwald beds with 80-90 taxa in the range through record. In St. Audrie’s Bay, the diversity decreases from the Westbury Formation into the Blue Lias Formation (Fig. 5b). The highest diversity in the Westbury Formation is reached with 61 taxa. The increasing and decreasing diversity at the base and top of the records is probably affected by the edge effect, which is an exaggerated high concentration of first occurrences at the beginning and a high concentration of last occurrences at the end of a record (Boltovskoy, 1988; Jaramillo, 2002). However, the parallel trend with the species richness record implies that the range-through record shows true changes in the diversity.

4. Discussion 4.1 Climate changes in Hochalplgraben and St. Audrie’s Bay

The PCA ordination diagrams of Hochalplgraben and St. Audrie’s Bay show a comparable taxa distribution (Fig. 4). This is an important result as it implies that despite the different vegetation history in both realms, the sporomorph assemblages have a similar composition. In Hochalplgraben, axis 1 is the humidity axis and axis 2 the temperature axis while this is the opposite in St. Audrie’s Bay. This is explained by the observation that the main vegetation change in Hochalplgraben is from Cheirolepidiaceae to spore producing plants (Fig. 2), probably related to the change in humidity levels (e.g., Vakhrameev, 1981; 1987; 1991). Such a major change between Cheirolepidiaceae and spore producing plants is not present in St. Audrie’s Bay (Fig. 3). Instead, there are changes between different conifer species where relative temperature might be a more important factor.

149 CHAPTER 7 5 3 CIE 0 3 initial 5 2 x FO C. thiergartii 0 x 2 x 5 95% limits 1 x x 0 1 Diversity (rarefaction) n pollen and spore taxa 5 x 0 0 7 0 6 0 5 0 4 0 3 0 2 Species richness 0 1 n pollen and spore taxa Diversity (range through) 0 5 2 - 7 [‰] 2 - org C 13 9 2 - δ δ 1 3 -

0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 70 60 50 4 30 20 10 00 9 9 80 70 60 50 4 30 20 10 00 800 40 cm

2 2 2 2 2 2 2 2 1 1 1 1 1 1 1 1 1 1 80 70 60 50 30 20 10 0 2

Williton Mb Williton Member Member

Blue Lias Formation Lias Blue Formation Westbury Langport Cotham Formation

Lilstock Fm Lilstock Blue Anchor Anchor Blue

B: St. Audrie’s Bay Audrie’s B: St.

Triassic Jurassic CIE 5 4 0 initial 4 5 3 FO C. thiergartii 0 3 5 2 0 x 2 x 95% limits 5 1 x 0 x 1 n pollen and spore taxa Diversity (rarefaction) 5 x 0 0 9 0 8 0 7 0 6 0 5 0 4 0 3 0 Species richness 2 0 1 n pollen and spore taxa Diversity (range through) 0 -24 -26 [‰] org -28 C 13 δ δ -30 -32 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0

9 8 7 6 5 4 3 2 1 8 7 6 5 4 3 2 1 0

1 1 1 1 1 1 1 1 1 Schattwald beds Schattwald

Tiefengraben Member Tiefengraben Eiberg Mb Eiberg

Kendlbach Formation Kendlbach Fm Kössen

A: Hochalplgraben Jurassic Triassic

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The sample scores on the first and the second axes from the PCA plotted through time result in relative humidity and temperature records for Hochalplgraben (Fig. 6a) and St. Audrie’s Bay (Fig. 6b). The initial negative carbon isotope excursion (CIE) (e.g., Hesselbo et al., 2002; Ruhl et al., 2009) and the first occurrence (FO) ofCerebropollenites thiergartii are used as correlation lines. A minor hiatus is present in the lower part of the Tiefengraben Member (Bonis et al., 2009a) from Hochalplgraben. This corresponds to the interval between 1450 and 1850 cm in St. Audrie’s Bay (Fig. 6b, grey interval in δ13C curve). Our correlation between both realms is in agreement with Hallam’s (1981) suggestion based on Schuurman’s (1977, 1979) data: that the Zlambach Beds, most of the Kössen Beds and most of the Westbury Formation (together with the topmost part of the underlying Mercia Mudstone Group) are age-equivalent, while the topmost Kössen Beds correspond with part of the Lilstock Forma- tion till the upper part of the Cotham Member. Prior to the initial CIE the vegetation was dominated by conifers (Fig. 2 and 3) and the climate was relative dry and warm in the Tethys realm and relatively dry and ‘cold’ in the north-western European realm (Fig. 6). It must be noted that the temperature gradient reconstructed by the PCA was probably small as a warm climate predominated the T-J boundary interval (e.g., Frakes, 1979; Sellwood and Valdes, 2007). The initial CIE coincides with peak temperatures and a relatively dry period. This is followed by cooling and increasing humidity in both sections. In the Tiefengraben Member (from 550 cm) in Hochalplgraben and in the Blue Lias Formation (from 1850 cm) in St. Audrie’s Bay, vegetation and climate patterns are completely different. Although warming is visible in both sections, Hochal- plgraben is characterized by an increasingly wetter climate while drier climate prevails in St. Audrie’s Bay (Fig. 6). This is visible in the vegetation diagram as the transition from a conifer-dominated assemblage to a vegetation type consisting mainly of ferns and liverworts in Hochalplgraben. By contrast, assemblages in St. Audrie’s Bay were completely dominated by Cheirolepidiaceae during this time. Distinct changes in high resolution δ18O records of fossil oysters from the UK (Korte et al., 2009) correspond with our reconstructed temperature axis from St. Audrie’s Bay. Both proxy records show a warming trend from the Triassic to the Jurassic. The fossil oyster record is not influenced by sea level changes. The similar climate signal from both the oyster and the sporomorph records suggests that the changes in sporomorph assemblages reflect true climate induced vegetation changes and that it is not biased by sea level changes. Therefore, the sporomorph record is a useful proxy for climate change (i.e., temperature, humidity).

Figure 5 (left): Palynological diversity patterns from a) Hochalplgraben and b) St. Audrie’s Bay. Rarefaction analysis is standardized on a pollen sum of 207 grains. Hochalplgraben

13 13 δ Corg data from Ruhl et al. (2009) and St. Audrie’s Bay δ Corg data from Hesselbo et al. (2002).

151 CHAPTER 7 CIE 0 . initial 2 + FO C. thiergartii 0 . 1 0 humidity . 0 - Sample score axis 2 (17.8% explained) 0 . 1 - 0 . 1 - + 0 . 0 0 . 1 temperature - Sample score axis 1 (46.5% explained) 0 . 2 5 2 - 7 [‰] 2 - org C 9 13 2 - δ δ 1 3 -

0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 70 60 50 4 30 20 10 00 9 9 00 10 20 30 4 50 60 70 80 800 40 cm

2 2 2 2 2 2 2 2 1 0 10 20 30

50 60 70 80

1 1 1 1 1 1 1 1 1 2

Williton Mb Williton Member Member

Blue Lias Formation Lias Blue Formation Westbury Langport Cotham

Formation

B: St. Audrie’s Bay Audrie’s B: St. Lilstock Fm Lilstock Blue Anchor Anchor Blue

Triassic Jurassic CIE 0 . initial 2 + FO C. thiergartii 0 . 1 0 . 0 temperature - 0 Sample score axis 2 (19.7% explained) . 1 - 0 . 1 - + 6 . 0 - 2 . 0 - 2 . 0 humidity 6 . - 0 Sample score axis 1 (43.5% explained) 0 . 1 -24 -26 [‰] org C -28 13 δ δ -30 -32 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0

9 8 7 6 5 4 3 2 1 8 7 6 5 4 3 2 1 0

1 1 1 1 1 1 1 1 1 Schattwald beds Schattwald

Tiefengraben Member Tiefengraben Eiberg Mb Eiberg

Kendlbach Formation Kendlbach Fm Kössen

A: Hochalplgraben

Jurassic Triassic

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4.2 Diversity patterns in Hochalplgraben and St. Audrie’s Bay

The first important result is that all diversity methods (qualitative: range through and species richness, and quantitative: rarefaction) show a similar pattern (Fig. 5). This implies that in this case even a relatively low sporomorph sum (~200) is sufficient to get a general idea about diversity trends. However, for a better estimation of the number of taxa per sample it is preferred to carry out a qualitative diversity analysis as the number of taxa can be a double to triple times higher (Fig. 5). Besides true first or last occurrences, taxa can also be temporarily absent due to changing climate conditions. Therefore, the number of taxa per sample (species richness) was plotted next to the range-through diversity method. Both methods show parallel trends. An almost fourfold increase in pollen and spore diversity occurs across the T-J boundary (Fig. 5a) in Hochalplgraben, which is linked to a climate change from relatively hot and arid to more humid conditions (Fig. 6a). In contrast to Hochalplgraben, diversity drasti- cally decreases from the latest Triassic to the earliest Jurassic in the St. Audrie’s Bay section (Fig. 5b) caused by a change to a warm an arid climate (Fig. 6b). Whereas the diversity increases to 80-90 taxa in the Schattwald beds, the diversity in the time-equivalent beds from St. Audrie’s Bay is much lower, about 50 taxa (Fig. 5). This can be explained by the different palaeogeographic positions of the sections: Hochalplgraben was situated in a tropical summerwet biome while St. Audrie’s Bay was located in a warm temperate biome (e.g., Chapter 3; Rees et al., 2000; Willis and McElwain, 2002). Our palynological diversity results from St. Audrie’s Bay correspond with the bivalve study by Mander et al. (2008), describing a loss of taxonomic richness from the Upper Westbury Formation to the lower Lilstock Formation but little convincing evidence of a catastrophic marine extinction. Figure 7 shows the diversity, extinction and origination in Hochalplgraben and St. Audrie’s Bay separately for the pollen and spores. The originations at the beginning and the extinctions at the end of the records are partly influenced by the edge-effect. However, true disappearances in the end of the St. Audrie’s Bay section (Fig. 7f) are observed as C. meyeriana becomes the monodominant species (Chapter 5). The most striking character in the Hochal- plgraben record is the gradual diversification of spores above the Kössen Formation and the maximum diversity in the Schattwald beds (Fig. 7a). However, most of these spores are known to have much longer stratigraphic ranges than recorded in these sections and were already present earlier in the Rhaetian (Bonis et al., 2009a). The virtual lack of spores in the Kössen Formation is probably caused by non-favourable environmental conditions limiting spore- producing plant growth during deposition (Bonis et al., 2009a): relatively arid and warm (Fig. 6a). The sudden diversity decrease at 550 cm in Hochalplgraben is not an extinction event as a

Figure 6: (left) PCA sample scores for axis 1 and 2 of a) Hochalplgraben b)

13 St. Audrie’s Bay. Hochalplgraben δ Corg data from Ruhl et al. (2009) and St. 13 Audrie’s Bay δ Corg data from Hesselbo et al. (2002).

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Diversity (range through) Origination Extinction

1000 1000 1000

900 900 900

800 800 800

700 700 700

600 600 600

500 500 500 Depth (cm) 400 400 400

300 300 300

200 200 200

100 100 100 A B C 0 0 0 0 10 20 30 40 50 60 70 0 1 2 3 4 5 6 7 8 9 10 11 12 13 0 1 2 3 4 5 6 7 8 9 n taxa n taxa n taxa

3000 3000 3000

2500 2500 2500

2000 2000 2000

1500 1500 1500 Depth (cm)

1000 1000 1000

500 500 500 pollen spores D E F 0 0 0 0 10 20 30 40 0 1 2 3 4 5 6 7 8 9 10 0 1 2 3 4 5 6 7 n taxa n taxa n taxa

Figure 7: Palynological diversity (range-through), origination, and extinction records from Hochalplgraben (a, b, c) and St. Audrie’s Bay (d, e, f).

hiatus disturbs the record. There are no evident extinction horizons present in the Hochal- plgraben and St. Audrie’s Bay sections (Fig. 7). Some end-Triassic taxa in St. Audrie’s Bay which disappear coincident with the initial negative CIE are Rhaetipollis germanicus, Ovalipollis pseudoalatus and Lunatisporites rhaeticus (Chapter 5). A possible explanation is that these pollen types were produced by conifers which could not cope with the change to a warmer and/or wetter climate (Fig. 6b). The diversity increase in Hochalplgraben is in sharp contrast to floral diversity data from other regions. The contrasting trends in floral diversity likely reflect regional differences in

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environmental stress, climatic changes and different palaeogeographic positions during the end-Triassic biotic crisis. In the Eocene, a similar increase in spore abundance and diversity reflects increased rainfall in a tropical climate (Jaramillo, 2002). This might be a Cenozoic analogy of the changes we see in Hochalplgraben (Fig. 5a and 6a), a site likely located in the tropical summerwet biome during deposition of the Tiefengraben Member (Chapter 3). Although the position of the T-J boundary in the North American basins is still under discussion (e.g., Hounslow et al., 2004; Kozur and Weems, 2005; Lucas and Tanner, 2007b; Cirilli et al., 2009) there is a clear palynological turnover and a strong sporomorph diversity decrease of about 60% in the Passaic Formation below the Jacksonwald Basalt (Fowell and Olsen, 1993). It is possible that this extinction is an artefact caused by a hiatus at this level (Van Veen, 1995; Kozur and Weems, 2005). The floral record from Greenland is ambiguous. A quantitative macrobotanical study from Greenland shows a decrease in standing species richness by about 85% (McElwain et al., 2007). Although the late Triassic event in East Greenland did not induce mass extinction of plant families, it accompanied major and abrupt change in their ecology and diversity (McElwain et al., 2009). However, it should be noted that macrofossil records (and derived diversity changes) are largely biased by taphonomical processes. Sheet splay or crevasse splay deposits are the primary facies in which the plant fossils occur (McEl- wain et al., 2007), implying a mix between the autochtonous and allochtonous vegetation. On the other hand, abandoned channels would be expected to preserve predominantly parautoch- tonous plant communities growing in close proximity to the channel (McElwain et al. 2007). In contrast to the macrofossil record, the Greenland sporomorph record does not show pronounced assemblage changes or an extinction event (Raunsgaard Pedersen and Lund, 1980; Koppelhus, 1997). The Newark Basin is thus far the only realm which shows a major extinction or diversity decrease. The diversity based on a range chart compiled from several Triassic studies and reviews also shows only a minor decline in sporomorph diversity (~20%) across the T-J boundary (Kürschner and Herngreen, in press). This is mainly the result of a decrease in the numbers of spore species (Kürschner and Herngreen, in press). Furthermore, it should be taken into account that this T-J diversity curve concerns stages and that typical middle Rhaetian sporomorphs like vesicate and bisaccate pollen are included in the diversity calculation. These are not included in the present study, which only includes the latest Rhaetian and earliest Hettangian. To summarize, we suggest that there is only a minor qualitative palynological extinction event across the T-J boundary, limited to the interior of Pangaea. Instead, the transition interval is characterized by climate-induced quantitative changes in the sporomorph assemblages.

4.3 Global Triassic-Jurassic boundary vegetation patterns and climate

A compilation of palynological data from the T-J transition interval gives insight in the global distribution of palynofloral assemblages (Fig. 8, Appendix 2). T-J boundary sections from the

155 CHAPTER 7

A: Latest Rhaetian

Classopollis dominated assemblage ? spore dominated assemblage bisaccate dominated assemblage mixed assemblage ? no clear stratigraphic framework ? ? Tethys ocean

CAMP

? ?

B: Earliest Hettangian

? Classopollis dominated assemblage spore dominated assemblage bisaccate dominated assemblage mixed assemblage ? ? ? no clear stratigraphic framework ? Tethys ocean

CAMP

? ?

Cratonic landmasses Marginal marine - Fluviolacustrine Deep ocean

Figure 8 (color version on p. 216): Position of sporomorph assemblages in a) the latest Rhaetian and b) the earliest Hettangian. (Map modified from Quan et al., 2008)

southern hemisphere with abundant sporomorphs are scarce and often lack a high resolution. Therefore, we focus on the northern hemisphere. It is obvious that the latest Rhaetian was dominated by mixed assemblages consisting of pollen and spores in various abundances. In the earliest Hettangian, Classopollis dominated assemblages spread across the interior of

156 CHAPTER 7

Pangaea, in areas affected by the basaltic volcanism of the Central Atlantic Magmatic Prov- ince (CAMP). Sections close to the Tethys Ocean show spore dominated assemblages. During the late Triassic, the Tethys, on the eastern edge of Pangaea, experienced a monsoonal climate because of the thermal contrast between the large continental plates along the equator and the sea (e.g., Parrish 1993; Satterley, 1996; Sellwood and Valdes 2007). Evidence for Upper Triassic monsoonal activity has been found, e.g., in the continental Chinle Formation, Colorado Plateau (Dubiel et al., 1991), from carbonate platforms in the Southern Alps (Mutti and Weissert, 1995), and from playa cycles in Germany (Reinhardt and Ricken, 2000; Vollmer et al., 2008). The monsoonal activity influences precipitation patterns, and consequently the floral distribution and development. Additionally, CAMP volcanism may have caused major climate change and substantial environmental disturbance, for example by the increase of atmospheric CO2 concentration and associated warming (Marzoli et al., 1999; McElwain et al. 1999, 2009; Huynh and Poulsen, 2005; Chapter 3, 4, 6). We propose that the observed changes in palynofloral distribution and development are the indirect result of CAMP volcanism. The increase in greenhouse gases caused a warmer climate and an enhanced thermal contrast between the continent and the seas. In addition to warming by greenhouses gases, the large area of dark basalt could have amplified warming by a negative albedo-effect. Consequently, the monsoon system got stronger and induced a drier interior and more intensive rainfall near the margins of the Tethys Ocean. While some palynological studies report a change to a more humid climate (i.e., increase in fern spores) across the T-J boundary (Bonis et al., 2009a; Götz et al. 2009), a change to more arid conditions on both sides of the North Atlantic rift during the earliest Jurassic is reflected by a palynofloral change to monotonous assemblages dominated by Classopollis meyeriana (e.g., Fowell et al., 1993; Van Veen, 1995; Whiteside et al., 2007; this study). Vakhrameev (1981) proposed the following classification based on the relative abundance of Classopollis: 1-10% - temperate climatic conditions, 20-50% - warm subtropical climate, >60% - arid climate. T-J boundary sporomorph records studied by Hubbard and Boulter (2000) show pronounced and rapid climatic fluctuations culminating in a dramatic and protracted cold event near the stage boundary. This cold event is ambiguous as some botanical affinities were misinterpreted (e.g.,C . meyeriana in the relatively cool group, see also Chapter 3). It should be noted that some species of the Cheirolepidiaceae group (e.g. Classopollis torosus) may be indicative of coastal habitats (e.g. Batten, 1974; Watson; 1988; Abbink, 1998; see also Chapter 3). We propose that Classopollis meyeriana/Cheirolepidiaceae dominance indicates a drier climate in the earliest Hettangian and that plants were capable of withstanding enhanced seasonality. ‘In general, one can conclude that that Cheirolepidiaceae were drought resistant, thermophillous shrubs and trees that required at least a subtropical climate’ (Francis, 1973; Vakhrameev, 1991 in Abbink, 1998). Further evidence of a dry Pangaean interior lies in the change from late Triassic fluvial-lacustrine facies to early Jurassic eolian dune and interdune facies in south-western U.S.A. (e.g., Tanner and Lucas, 2007).

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On the other hand, increased rainfall along the Tethys margin explains the shift to the Hettangian spore-dominated assemblages in Hochalplgraben. Wetter climate is also reflected in the lithology. Reefs were numerous in the latest Triassic Tethys (e.g., Stanley, 2003; Seuss et al., 2005; Kiessling et al., 2007).These need clear waters and therefore, runoff was probably negligible. The change to a wetter climate (Fig. 6) coincides with carbonate break up (reef extinction) (Hallam, 2002; Hautmann, 2004; Seuss et al., 2005) and an increasing amount of silty marls, indicating enhanced runoff. Also the spore shift in the Tatra Mountains, Slovakia, is interpreted to display a sudden increase in humidity most probably caused by the volcanic activity of the CAMP (Ruckwied et al., 2009). We agree that volcanic activity was indeed associated with changes in oceanic and atmospheric circulation patterns, regionally resulting in increasing precipitation and/or humidity (Ruckwied et al., 2009). A recent study from the Germanic Basin showed a severe and wide ranging vegetation shift from gymnosperm forests to pioneer assemblages across the T-J boundary, linked to CAMP volcanism (Van de Schoot- brugge et al., 2009). However, the presented data are not straightforward as the sections may contain unconformities and the sporomorph changes could be biased by marked facies changes (Chapter 5, Kürschner and Herngreen, in press and references therein). Higher latitude records show differences in relative abundances and/or composition between the latest Rhaetian and earliest Hettangian but dominance by spores or Classopollis does not occur (Appendix 2). This is probably related to the higher latitude where a warm temperate climate prevailed (e.g., biome map fig. 5.14 in Willis and McElwain, 2002) and the impact of the monsoon system was less severe. Continental T-J boundary deposits from the Boreal realm (the North-West of the West-Siberian lowland) show sporomorph assemblages consisting of a variety of spores and pollen types (Rovnina, 1972). There are no marked assemblage changes across the boundary. A remarkable difference compared to other T-J records is the high abundance of monosulcate pollen produced by Ginkgoaceae, Bennettitales, and Peltaspermales (Rovnina, 1972), implying relatively ‘cold’ and humid conditions in the highest latitudes during the T-J transition interval. This is confirmed by the very low abundance of Classopollis. The different distribution of palynological assemblages across the T-J boundary follows the climate changes inferred by Pangaean climate models. Mesozoic palaeoclimate reconstructions generated on a general circulation model by Sellwood and Valdes (2007) shows that the world was predominantly warm and that rainfall often focused over the oceans, leaving major desert expanses on the continental areas. There was no ice present on the high palaeolatitudes (Frakes et al., 1992; Satterley, 1996; Hallam and Wignall, 1999). A series of sensitivity experiments was done with a late Triassic numerical coupled ocean-atmosphere

climate model to predict extreme environmental condition with 2-8x pre-industrial CO2

levels (Huynh and Poulsen, 2005). On land, increasing CO2 caused extreme heating, intense seasonal fluctuations of surface temperatures, an increase in the number and severity of hot days and days without precipitation, and an exponential rise in the land surface area experi- encing affected by heat and aridity (Huynh and Poulsen, 2005). Sensitivity experiments by

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Kutzbach and Gallimore (1989) also indicated a strong seasonality in low and middle latitudes, with seasonal intense precipitation and aridity. All of the experiments simulate dry conditions in the tropics (except along the east and west coasts), seasonally dry conditions in the mid-latitude continental interiors, and year-round moist conditions only in middle or high latitudes (Kutzbach and Gallimore, 1989). These model results are in agreement with the observed change to monotonous Cheirolepidiaceae vegetation in the continental interior of Pangaea and the fern-dominated assemblages near the Tethys Ocean.

5. Conclusions

Instead of a major palynofloral extinction event, the T-J boundary is characterized by climate- induced quantitative changes in the sporomorph assemblages. We propose that the changes in palynofloral distribution and development are the indirect result of CAMP volcanism. The increase in greenhouse gases caused a warmer climate and an enhanced thermal contrast between the continent and the seas. Consequently, the monsoon system got stronger and induced a drier interior and more intensive rainfall near the margins of the Tethys Ocean. The latest Rhaetian was dominated by mixed sporomorph assemblages consisting of pollen and spores in various abundances. Increased rainfall along the Tethys margin explains the shift to the diverse Hettangian spore-dominated assemblages in Hochalplgraben. In contrast, the interior of Pangaea became more arid and warmer. This can be seen in e.g. the St. Audrie’s Bay section, where low diversity Classopollis meyeriana assemblages dominated the record. Our results are in agreement with the Pangaean climate changes inferred by climate model studies where the world was predominantly warm with seasonal intense aridity and precipitation.

Acknowledgements

We acknowledge funding from the “High Potential” stimulation program of Utrecht University. We thank J. van Tongeren and N. Welters for laboratory assistance. S. Hesselbo is thanked for providing us with samples from St. Audrie’s Bay. M. Ruhl, M. Deenen, L. Krystyn, A. Von Hillebrandt, W. Krijgsman and M. Hounslow are thanked for their collaboration in the field and useful discussions. We gratefully acknowledge H. Visscher and H. van Konijnenburg - van Cittert for their comments on an earlier version of the manuscript.

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Appendix 1

Botanical affinities of the sporomorphs found in the Hochalplgraben (H) and St. Audrie’s Bay (S) section

Plant groups Detailed botanical affinity H S POLLEN Cheirolepidiaceaen Conifers Coniferophyta - Cheirolepidiaceae Classopollis kedangensis x Classopollis meyeriana x x Classopollis murphyi x x Classopollis sp. x Classopollis torosus x x Granuloperticulatipollis rudis x

Other conifers Other Coniferophyta Araucariacites australis Araucariaceae x x Pinuspollenites minimus Pinaceae x x Quadraeculina anellaeformis Podocarpaceae x x Cerebropollenites thiergartii Taxodiaceae x Perinopollenites elatoides Taxodiaceae x x Tsugaepollenites pseudomassulae Taxodiaceae x x Lunatisporites rhaeticus Voltziaceae x x Platysaccus spp. Voltziaceae x x Triadispora sp. Voltziaceae x Enzonalasporites vigens Glyptolepis x Florinites pellucidus Cordaites x

Seed ferns Pteridospermophyta Vitreisporites bjuvensis Caytoniales x x Vitreisporites pallidus Caytoniales x x Vitreisporites spp. Caytoniales x Alisporites diaphanus Corystospermales x x Alisporites radialis Corystospermales x x Alisporites robustus Corystospermales x x Striatoabieites aytugii x Vesicaspora fuscus x x

Other gymnosperms Chasmatosporites apertus Cycadophyta - Cycadales x x Cycadopites spp. Cycadophyta, Ginkgoales, Peltaspermales x x Ephedripites spp. Gnetales x x Eucommiidites troedsonii probably Gnetales x x Lagenella martini x Ovalipollis pseudoalatus x x Rhaetipollis germanicus x x

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Plant groups Detailed botanical affinity H S SPORES Mosses Bryophyta Annulispora folliculosa x Camarozonosporites aulosenensis x Nevesisporites bigranulatus x x Rogalskaisporites cicatricosus x Plant groups Detailed botanical affinity H S Stereisporites australis x POLLEN Stereisporites punctatus x Cheirolepidiaceaen Conifers Coniferophyta - Cheirolepidiaceae Stereisporites radiatus x Classopollis kedangensis x Stereisporites sp. div. x Classopollis meyeriana x x Classopollis murphyi x x Liverworts Hepatophyta Classopollis sp. x Porcellispora longdonensis x x Classopollis torosus x x Ricciisporites tuberculatus x x Granuloperticulatipollis rudis x Horsetails Sphenophyta Other conifers Other Coniferophyta Calamospora tener Equisetales x x Araucariacites australis Araucariaceae x x Pinuspollenites minimus Pinaceae x x Clubmosses Lycophyta Quadraeculina anellaeformis Podocarpaceae x x Aratrisporites minimus Isoetales x x Cerebropollenites thiergartii Taxodiaceae x Aratrisporites parvispinosus Isoetales x Perinopollenites elatoides Taxodiaceae x x Aratrisporites spp. Isoetales x Tsugaepollenites pseudomassulae Taxodiaceae x x Lycopodiacidites rhaeticus Lycopodiales x x Lunatisporites rhaeticus Voltziaceae x x Lycopodiacidites rugulatus Lycopodiales x x Platysaccus spp. Voltziaceae x x Densosporites fissus Selaginellales x x Triadispora sp. Voltziaceae x Foveosporites spp. Selaginellales x x Enzonalasporites vigens Glyptolepis x Heliosporites reissingeri Selaginellales x x Florinites pellucidus Cordaites x Limbosporites lundbladii Selaginellales x x Uvaesporites argentaeformis Selaginellales x Seed ferns Pteridospermophyta Uvaesporites microverrucatus Selaginellales x Vitreisporites bjuvensis Caytoniales x x cf. Uvaesporites sp. Selaginellales x Vitreisporites pallidus Caytoniales x x Densoisporites nejburgii Pleuromeia x x Vitreisporites spp. Caytoniales x Acanthotriletes varius x x Alisporites diaphanus Corystospermales x x Camarozonosporites laevigatus x x Alisporites radialis Corystospermales x x Camarozonosporites rudis x Alisporites robustus Corystospermales x x Carnisporites anteriscus x Striatoabieites aytugii x Carnisporites lecythus x Vesicaspora fuscus x x Carnisporites leviornatus x Carnisporites megaspiniger x Other gymnosperms Carnisporites spiniger x Chasmatosporites apertus Cycadophyta - Cycadales x x Carnisporites sp. x Cycadopites spp. Cycadophyta, Ginkgoales, Peltaspermales x x Carnisporites sp. div. x Ephedripites spp. Gnetales x x Cingulizonates rhaeticus x Eucommiidites troedsonii probably Gnetales x x Densosporites irregularis x Lagenella martini x Leptolepidites spp. x x Ovalipollis pseudoalatus x x Retitriletes clavatoides x Rhaetipollis germanicus x x

161 CHAPTER 7

Plant groups Detailed botanical affinity H S Retitriletes gracilis x Retitriletes sp. div. x Tigrisporites microrugulatus x Triancoraesporites ancorae x Triancoraesporites reticulatus x

Ferns Pteridophyta Baculatisporites spp. Osmundales - Osmundaceae x x Conbaculatisporites spp. Osmundales - Osmundaceae x x Todisporites sp. div. Osmundales - Osmundaceae x x Perinosporites thuringiacus Filicales - Cyatheaceae x x Zebrasporites interscriptus Filicales - Cyatheaceae x x Zebrasporites laevigatus Filicales - Cyatheaceae x x Converrucosisporites luebbenensis Filicales - Dipteridaceae* x x Deltoidospora spp. Filicales - Dicksoniaceae, Cyatheaceae, x x Dipteridaceae, Matoniaceae* Thymospora canaliculata Filicales - Marattiaceae?** x Concavisporites spp. Filicales - Matoniaceae* x x Polypodiisporites ipsviciensis Filicales - Schizaeaceae x x Polypodiisporites polymicroforatus Filicales - Schizaeaceae x x Ischyosporites variegatus Filicales - Schizaeaceae x x Kyrtomisporis laevigatus x x Kyrtomisporis speciosus x x Lophotriletes verrucosus x x Trachysporites fuscus x x Verrucosisporites cheneyi x Verrucosisporites sp. x

Other spore producing plants Asseretospora gyrata x x Cosmosporites elegans x x cf. Cyclotriletes sp. x x Echinitosporites iliacoides x cf. Guthoerlisporites x Neochomotriletes triangularis x x Paraklukisporites foraminis x Platyptera trilingua x Polycingulatisporites bicollateralis x

Selagosporis mesozoicus x Semiretisporis gothae x

* in figure 2 and 3 these have been grouped together ** because of the very sparse abundance, added to ‘Other ferns’ in figure 2 and 3

162 CHAPTER 7

Retitriletes gracilis x Retitriletes sp. div. x Tigrisporites microrugulatus x Triancoraesporites ancorae x Triancoraesporites reticulatus x

Ferns Pteridophyta Baculatisporites spp. Osmundales - Osmundaceae x x Conbaculatisporites spp. Osmundales - Osmundaceae x x Todisporites sp. div. Osmundales - Osmundaceae x x Cirilli et al. (1994) Cirilli et al. Reference Chapter 1 (2009), Bonis et al. (2008) et al. Warrington Chapter 5, (2009) de Schootbrugge van et al. Chapter 6 Lund (2003) Dybkjær (1991) Dybkjær (1988) Lund (1977) Vigran (2009) Hochuli and Lindström and Erlström (2006) (2009) de Schootbrugge van et al. and Lund (1980); Raunsgaard Pedersen Koppelhus (1997) (2007) Galli et al. Perinosporites thuringiacus Filicales - Cyatheaceae x x Zebrasporites interscriptus Filicales - Cyatheaceae x x Zebrasporites laevigatus Filicales - Cyatheaceae x x Converrucosisporites luebbenensis Filicales - Dipteridaceae* x x Deltoidospora spp. Filicales - Dicksoniaceae, Cyatheaceae, x x Dipteridaceae, Matoniaceae* Thymospora canaliculata Filicales - Marattiaceae?** x Concavisporites spp. Filicales - Matoniaceae* x x

Polypodiisporites ipsviciensis Filicales - Schizaeaceae x x Pinuspollenites ) Polypodiisporites polymicroforatus Filicales - Schizaeaceae x x Ischyosporites variegatus Filicales - Schizaeaceae x x Kyrtomisporis laevigatus x x

Kyrtomisporis speciosus x x Earliest Hettangian palynoflora Deltoidospora, Ricciisporites, mainly spores: Trachysporites Heliosporites, meyeriana mainly C. Selaginellales Pinales, assemblage: mixed Classopollis , assemblage (e.g. mixed , Deltoidospora maxima Pinuspollenites minimus, ornamented spores are rare pollen bisaccate spores dominate, pollen only a minor part and non-saccate abundant torosus , increase in C. pollen saccate Pinuspollenites minimus , trilete spores (20-50%) dominate, common presence of Pinuspollenites, trilete Chasmatosporites spores present , , elatoides by Perinopollenites dominated Ricciisporites, Deltoidospora tree conifers, seed ferns, assemblage, mixed cycads/ginkgos/ground ferns, ferns pollen common fern spores, acme of Heliosporitesand other pollen , data) and spore taxa (no quantitative Lophotriletes verrucosus x x Trachysporites fuscus x x Verrucosisporites cheneyi x Verrucosisporites sp. x , Ovalipollis Ricciisporites

Other spore producing plants Asseretospora gyrata x x Cosmosporites elegans x x Tsugaepollenites cf. Cyclotriletes sp. x x Deltoidospora

Echinitosporites iliacoides x Vitreisporites spores , Monosulcites ) cf. Guthoerlisporites x Neochomotriletes triangularis x x Paraklukisporites foraminis x C. torosus , meyeriana , C. C. palynoflora: mixed , Ovalipollis torous , C. meyeriana, C. palynoflora: mixed Ricciisporites Ovalipollis, Gymnosperm forests ( Classopollis, , Ovalipollis Ricciisporites (probably local), dominates assemblage mixed data) spores (no quantitative dominance of C.torosus and common occurrence of Ricciisporites mainly trilete spores dominant (78-87%), and Baculatisporites Deltoidospora by Polypodiisporites , dominated assemblages mixed , elatoides by Perinopollenites dominated Ricciisporites , cycads/ginkgos, assemablage: mixed groundconifers tree ferns, ferns, frequent Ricciisporitescommon fern , pollen ( Rhaetipollis ) spores, palynoflora e.g., Rhaetipollis mixed data) Classopollis (no quantitative germanicus, by e.g. Classopollis , Dominated and rich of trilete variation spores , Triancoraesporites Latest Rhaetian palynoflora Platyptera trilingua x Polycingulatisporites bicollateralis x Selagosporis mesozoicus x Semiretisporis gothae x x 2 x ppendi Eiberg Basin (Austria) (Southern Bay UK) Audrie’s St. Germany (GermanicWüstenwelsberg Basin) (NW Germanic Eitzendorf 8 well Basin) Danish Subbasin (Danmark) Danish Subbasin Gassum borehole, (Danmark) North Sea Basin Barents Sea Scania (southern Sweden) Hollviken (southern Sweden) Land (East Greenland) Jameson westernAlps (Italy) Southern NorthernAppenines (Italy) Location A earliestsporomorphHettangianassemblages and latestRhaetian of Compilation

163 CHAPTER 7 Hankel (1987) Whiteside (1995); Traverse and Fowell (2007) et al. (2009) Cirilli et al. (1975) Traverse Cornet and (1994) et al. Fowell Kürschner(in prep.) and Batenburg (2006) Waanders Cornet and Marzoli et al. (2004) Marzoli et al. (1986) Adloff et al. Götz et al. (2009) Götz et al. Ruckwied and Götz (2009) Schuurman (1977) (1995) Rauscher et al. Cittert Konijnenburg-van Muir and van (1970) (2006) Barrón et al. (2006) Gómez et al. (1977) et al. Palain (2007) Whiteside et al. Ziaja (2006) (1983) Orlowska-Zwolinska Cyathidites (no Cyathidites Aratrisporites Trachysporites predominant Classopollis chataeunovi meyeriana never C. mainly Classopollis, C.torosus and more than 10%, murphyi dominant C. mainly Classopollis meyeriana (>95%) mainly Classopollis meyeriana (95%) by Classopollis meyeriana dominated by Classopollis meyeriana dominated mainly Classopollis meyeriana , Deltoidospora increase in spores: Concavisporites , Classopollis dominates 99% Classpollis harrisii over in general C. mainly Classopollis , torosus . C. meyeriana over dominated torosus meyeriana and C. C. by Classopollis (80-94% dominated till 100%) mainly Classopollis meyeriana C. torosus , by C. dominated Concavisporites minimus, , data) quantitative Concavisporites, Deltoidospora, are dominant Cyathidites

Ovalipollis, spores Ovalipollis, , bisaccates densus , P. Cyathidites Ricciisporites , overwhelming dominance of bisaccates overwhelming higher up mostly dominance of Classopollis , bisaccates and assemblage with monosaccates mixed bisaccates by Classpollis meyeriana dominated dominated by trilete dominated spores mixed assemblage, high abundance of high abundance assemblage, mixed Ricciisporites Classopollis abundant assemblages, mixed Classopollis , coniferous forests meyeriana with less by C. dominated torosus common C. diverse assemblage murphya , meyeriana and C. C. abundant multiple grainsdensus of Patinasporites Classopollis , abundant Ricciisporites, predominant ,Deltoidospora Concavisporites , Rhaetipollis, high amount of Classopollis, and numerous trilete Ovalipollis spores Fundy Basin (Canada) Fundy Basin (Canada) Hartford Basin (eastern USA) Basin (eastern Newark USA) western USA Group (western USA) Glen Canyon Tripolitania (north western Libya) Tripolitania Basin (Tanzania) Luwegu Tatra mountains (Slovakia) Tatra Northeastern France and southern Luxemburg Basin (France) Paris Carentan Basin (northern France) Asturias (northern Spain) Iberian range (Spain) Portugal Argana Basin (Morocco) Atlas basin (Morocco) Central High Holy Cross mountains (Poland) Holy Cross mountains (Poland) Basin Polish Cs ő vár section (northern Hungary)

164 CHAPTER 7 Rovnina (1972) Rovnina Lu and Deng (2005) Zhang and Grant-Mackie (2001) de Jersey and Raine (1990) (2008) Bomfleur et al. (2005) Grice et al. (1987) Helby et al. Semenova (1970) Semenova Embry and Suneby (1994) (2007) Yaroshenko , Baculatisporites , mixed assemblage, e.g., spores, spores, e.g., assemblage, mixed bisaccates monosulcates, , mainly Asseretospora fern spores (57.14%), Crassitudisporites and Dictyophyllidites than pteridophytes more abundant gymnospermous pollen Classopollis present and mainly spores, in the upper part abundant of the earliest Hettangian zone Classopollis dominated by Classopollis, dominated torosus especially C. Classopollis abundant mixed assemblage of spores, bisaccates bisaccates assemblage of spores, mixed and other pollen grains mixed assemblage: bisaccates, bisaccates, assemblage: mixed Osmundacidites , Lycopodiacidites Matonisporites , , Dictyophyllidites mainly coniferous bisaccates mainly coniferous bisaccates gymnospermous pteridophytes are abundant, pollen grains occupy second place mainly spores Alisporites dominated Falcisporites miospores, Triassic diverse australis mixed assemblage, e.g., spores, spores, e.g., assemblage, mixed bisaccates monosulcates, bisaccate and monosulcate pollen and monosulcate bisaccate dominated by Ricciisporites dominated (up to 60%), pollen grains not numerous spores, Ricciisporites , mixed assemblage: abundant Ricciisporites abundant , assemblage: mixed Rhaetipollis and Limbosporites Junggar Basin (China) New Zealand Hettangian may New Zealand (lower be not present) Victoria Land (Antarctica ) North Australia NW Australia North-west of the West Siberian West North-west of the lowland Western Ciscaucasia (Georgia) Western Donets Basin (Ukraine) Sverdrup Basin (Arctic Canada)

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Algemene introductie en samenvatting

De Trias-Jura (T-J) grens, ~201.58 miljoen jaar geleden, staat algemeen bekend als een periode waarin een van de vijf grote massa extincties in de geschiedenis van de aarde plaatsvond. Het is echter een van de minst onderzochte extincties, omdat complete en goed geda- teerde secties die dit tijdsinterval bevatten schaars zijn. Om de eind-Trias massa extinctie te bestuderen, werd er een multidisciplinair project opgezet: ‘Earth’s and life’s history: from core to biosphere (CoBi)’. Met verschillende disciplines bestaande uit palaeomagnetisme, cyclostratigrafie, geochemie en biogeologie, worden vragen over de timing, oorzaak en patronen van het extinctie interval onderzocht. Dit proefschrift richt zich op de biogeologische kant van het project. Er worden twee technieken gebruikt om palaeomilieuveranderingen te reconstrueren: 1) palynologie en 2) huidmondjesanalyse op fossiele bladeren.

Verklaringen voor de biologische omslag tijdens het laat Trias bestaan uit geleidelijke en catastrofale mechanismen. Mariene extincties kunnen gerelateerd worden aan het verlies van leefomgeving op het continentaal plat tijdens zeespiegeldaling in het laat Trias, maar dit kan niet het uitsterven verklaren van soorten op het land (bv. tetrapoden). Een vaak voorgesteld mechanisme is het grootschalige vulkanisme (uitstromen van basalt) van de ‘Central Atlantic Magmatic Province (CAMP)’, gerelateerd aan het uiteenvallen van Pangaea. Het CAMP basalt vulkanisme kan geassocieerd worden met de volgende effecten: snelle wereldwijde opwarming, een anoxische oceaan of grotere nutriënten toevoer in de oceaan (of beide), een calcificatie crisis, en een verschuiving naar negatievere koolstofisotopen waarden. Het T-J grens interval is gekarakteriseerd door twee duidelijke perturbaties in de stabiele koolstofisotopen profielen van secties binnen en buiten de Tethys. Een relatief kortdurende daling en teruggang van de verhouding tussen zware en lichte koolstofisotopen 13( C/12C) gedurende het laat Trias (‘initial’ carbon isotope excursion, CIE), gaat vooraf aan een meer geleidelijke verschuiving naar negatievere waarden (‘main’ CIE) aan de basis van de Jura. Het vrijkomen

van grote hoeveelheden koolstofdioxide (CO2) in de atmosfeer door het CAMP vulkanisme veroorzaakte een klimaatverandering en verstoring van ecosystemen. Het vrijkomen van methaan uit gashydraten is een andere belangrijke gebeurtenis waarvan gesuggereerd wordt dat het geassocieerd is met het vulkanisme in het laat Trias. Een alternatief catastrofaal mechanisme dat de extinctie veroorzaakt zou kunnen hebben, is een meteoriet inslag, maar tot op heden is hier geen overtuigend bewijs voor gevonden.

Voorbeelden van biotische verstoringen in het laat Trias zijn: het uitsterven van tetrapoden, abrupte en aanzienlijke veranderingen in de samenstelling van brachiopoden en bivalven leefgemeenschappen, het uiteindelijke uitsterven van conodonten waarbij sporadische

188 SAMENVATTING

overlevenden voorkomen in het Hettangien, een wereldwijde soortenverandering in radio- lariën, het uiteindelijke uitsterven van de al laag-diverse ammonieten leefgemeenschappen, en het ineenstorten van rif ecosystemen. De grootschaligheid en het exacte patroon (abrupt, stapsgewijs, of geleidelijk) van de extincties in het laat Trias worden echter betwist.

Hoewel de T-J overgang gekarakteriseerd wordt door extincties in mariene ecosystemen, is bewijs voor een vegetatieomslag op het land niet eenduidig. Een recent palynologisch onderzoek van het Germaanse bekken laat een duidelijke vegetatieverandering zien op de T-J grens, veroorzaakt door het CAMP vulkanisme. Een aanzienlijke extinctie van 60% sporomorfen taxa, gevolgd door een scherpe sporenpiek op de T-J grens, is gevonden in het Newark bekken in Noord-Amerika. Deze sporenpiek wordt gevolgd door een coniferen (Cheirolepidiaceae) gedomineerde palynoflora welke door sommigen gebruikt wordt om de basis van de Jura in het Newark bekken vast te leggen. In tegenstelling tot de Noord-Amerikaanse data, laten de meeste palynologische studies van Europese secties geleidelijke veranderingen in de soortensamenstelling zien op de T-J grens. Ook het T-J macropaleobotanische archief is tegenstrijdig. Kwantitatieve macrobotanische data uit Oost Groenland laten zien, dat bossen in het Trias met een hoge diversiteit vervangen werden door bossen met een lagere diversiteit en dat er een geleidelijke extinctie plaatsvond al vóór de T-J grens. Palynologische profielen uit Groenland laten geen duidelijke veranderingen zien in diversiteit of soortensamenstelling, noch enig overtuigend bewijs voor een extinctie.

Palynologische gegevens van de T-J grens worden controversieel bediscussieerd omdat secties met een toereikende resolutie en/of goed stratigrafisch kader schaars zijn. Verder zijn veel profielen kwalitatief (aanwezig/afwezig zijn van taxa) bestudeerd. Daarom is het belangrijk om hoge-resolutie kwantitatieve studies uit te voeren. Het voorkomen van goed bewaarde ammonieten en palynomorfen in belangrijke T-J grens secties uit de Noordelijke Kalkalpen (Oostenrijk) en het zuiden van Engeland maken een integratie van terrestrische microflorale gebeurtenissen in een marien biostratigrafisch kader mogelijk. De palynologische data in dit proefschrift zijn gebruikt om 1) een solide palynostratigrafisch kader voor het T-J grens interval te verkrijgen, 2) een reconstructie te maken van vegetatie- en klimaatveranderingen, en 3) meer inzicht te krijgen in de orde van grootte en de aard van de vegetatieveranderingen op de T-J grens.

De Kuhjoch sectie in de Karwendel syncline (Noordelijke Kalkalpen, Oostenrijk) is recentelijk gekozen als de ‘Global boundary Stratotype Section and Point (GSSP)’ voor de basis van de Jura. Het eerste voorkomen van de ammoniet Psiloceras spelae tirolicum is gekozen als primaire markering voor het herkennen van de T-J grens. Hoofdstuk 1 beschrijft de resultaten van een gedetailleerde kwantitatieve palynologische en koolstofisotopen studie aan de T-J grens van de nabijgelegen (5 km) Hochalplgraben sectie, evenals de eerste data van de Kuhjoch sectie. Beide secties werden afgezet in het Eiberg bekken in de westelijke Tethys. Het eerste voorko-

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men van Cerebropollenites thiergartii is bruikbaar als een palynologische markering voor de basis van de Jura omdat het de enige palynomorf is waarvan het eerste voorkomen ongeveer gelijk valt met dat van Psiloceras spelae tirolicum. De laat Trias ‘initial’ negatieve CIE is te zien op de overgang van de Kössen Formatie naar de Kendlbach Formatie. De palynologische en koolstofisotopen profielen van verschillende secties binnen het Eiberg bekken (Hochalplgra- ben, Kuhjoch and Tiefengraben) komen erg goed overeen en het verkregen palynostratigrafis- che schema kan gebruikt worden voor gedetailleerde locale en regionale correlaties.

Hoofdstuk 2 richt zich op de oorzaak en aard van de ‘black shale’ afzettingen in het Eiberg bekken. Deze vallen samen met de ‘initial’ negatieve CIE, een massaal voorkomen van groenalgen, en duidelijke vegetatieveranderingen. Een verhoogde toevoer van coniferen pollen (Cheirolepidiaceae) en hogere totale organische koolstofconcentraties worden gevolgd door een groenalgen piek (Cymatiosphaera). Een model wordt gepresenteerd waarin een toename in terrestrisch organisch materiaal gerelateerd wordt aan een toenemende seizoenaliteit en toenemende erosie van het achterland. Een sterkere toevoer van zoetwater leidde tot een lager zoutgehalte van het oppervlaktewater en het massale voorkomen van groenalgen. Stratificatie van de waterkolom zou anoxische diepwater condities en ‘black shale’ afzetting tot gevolg hebben gedurende de ‘initial’ CIE aan de basis van de Kendlbach Formatie.

De vegetatieveranderingen die samenvallen met de laat Trias ‘initial’ CIE zijn het onderwerp van Hoofdstuk 3. Hoge resolutie palynologische data van Hochalplgraben en Kuhjoch zijn gebruikt om relatieve veranderingen in temperatuur en vochtigheidsgraad te ontrafelen met behulp van multivariate statistische analysetechnieken. Voor de eerste keer worden relatieve klimaatveranderingen direct gekoppeld aan de ‘initial’ CIE. Het sporomorfen profiel laat zien dat een vegetatie gedomineerd door houtige sclerophylle coniferen vervangen werd door een gemengde sporenplanten (bv. varens, mossen en levermossen) en gymnospermen vegetatie. Snelle opwarming valt samen met de verschuiving van koolstof isotopen naar negatievere waarden. De CIE valt ook samen met een trend van een relatief droog naar een vochtiger klimaat. De snelle wereldwijde opwarming gedurende de CIE is waarschijnlijk het resultaat

van CO2 uitstoot door het CAMP vulkanisme en het daaraan gerelateerde vrijkomen van methaan. Deze opwarming had een noordwaartse verschuiving van de tropische gordel op het land aangrenzend aan de westerse Tethys tot gevolg.

Of veranderingen in de T-J koolstofkringloop een wereldwijd signaal zijn, wordt betwist vanwege de verschillen in grootte van de CIE tussen locaties en de mogelijke invloed van veranderingen in de herkomst van het organisch materiaal (terrestrisch en/of marien) op de totale organische koolstofisotopen samenstelling. Daarom wordt inHoofdstuk 4 de werkelijke grootte van de T-J koolstofkringloop veranderingen geschat op basis van koolstofisotopen metingen van specifieke organische verbindingen (n-alkanen) uit de waslaagjes op bladeren van landplanten. Data van Kuhjoch laten een 5-6‰ negatieve CIE zien, met koolstof isotopen

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waarden die 2-3 ‰ lager zijn dan voorheen aangenomen werd. Het blijkt dat eerdere verklaringen, uitsluitend gebaseerd op vulkanisme, niet langer alleen verantwoordelijk kunnen zijn voor deze negatievere waarden. De grootte en snelheid van de veranderingen in de koolstofisotopen waarden impliceert dat tot aan ~6900-8200 Gt 13C verminderd koolstof vanuit het methaan hydraten reservoir naar de atmosfeer en oceanen getransporteerd moet zijn. De klimaatveranderingen die gereconstrueerd zijn door middel van palynologie worden bediscussieerd met betrekking tot veranderingen in de koolstofkringloop tijdens de ‘initial’ CIE. Het blijkt dat de laat Trias koolstofkringloop veranderingen samenvallen met een sterke opwarming en een versterkte hydrologische kringloop. Deze data bevestigen het causaal verband tussen het grootschalige uitstoten van methaan, klimaatverandering en de extincties op het T-J grens interval.

Een andere belangrijke T-J grens sectie is St. Audrie’s Bay in Somerset, in het zuidwesten van de UK. Hoofdstuk 5 presenteert een hoge resolutie kwantitatieve palynologische studie van deze sectie, welke een overgangsinterval laat zien met vier uitgesproken sporenpieken tijdens het laat Trias. Regelmatige cyclische toenames in de palynomorfen concentraties kunnen gekoppeld worden aan perioden van versterkte runoff, samenvallend met de orbitale excentriciteit cyclus. De sporenpieken zijn waarschijnlijk gerelateerd aan de variabiliteit in de intensiteit van de moesson veroorzaakt door de orbitale precessie cyclus. Er is geen bewijs voor een wereldwijde sporenpiek in het laat Trias die, analoog aan de sporenpiek op de Krijt-Tertiair grens, gerelateerd is aan een catastrofale massa extinctie. Klimaatverandering is een meer aannemelijk mechanisme voor de toenames in de hoeveelheid sporen.

Wereldwijde opwarming door de massale CO2 uitstoot tijdens de afzetting van de CAMP is een van de belangrijkste klimaatveranderingen die gesuggereerd is voor het T-J grens interval.

Fossiele bladeren zijn bruikbaar als biosensor van palaeoatmosferische CO2 waarden vanwege de omgekeerde relatie tussen de atmosferische CO2 concentratie en de ontwikkeling van het aantal stomata (huidmondjes) per bladoppervlak tijdens de groei van het blad. In Hoofdstuk 6 worden fluctuaties in CO2 concentraties gereconstrueerd door middel van stomata analyse op één plantensoort: de zaadvaren Lepidopteris ottonis. De stomata-index laat geen significante variatie zien binnen één pinnule van een blad en tussen verschillende pinnulen van een blad, waardoor het een bruikbare proxy is om CO2 veranderingen in het verleden te reconstrueren. Een dalende stomata-index en stomata dichtheid van de basis naar de top in de Wüstenwels- berg sectie (Bavaria, Germany) impliceert stijgende CO2 concentraties gedurende de T-J overgang. Aanvullende data van de hoeveelheid stomata van fossiele ginkgoales bladeren

(Ginkgoites taeniatus) suggereert een palaeoatmosferische CO2 waarde van 2750 ppmv voor de laat Trias.

De hoge resolutie kwantitatieve palynologische gegevens van Hochalplgraben en St. Audrie’s Bay (Hoofdstuk 1 en 5) worden gebruikt om veranderingen in vegetatie, biodiversiteit en

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klimaat te reconstrueren in Hoofdstuk 7. Hochalplgraben laat een verandering zien van gymnospermen naar sporenplanten, en een toename in de sporendiversiteit tijdens het laatste Rhaetien. Multivariate statistische analyses laten een trend zien naar een vochtiger klimaat tijdens de T-J grens. In tegenstelling tot deze gegevens wordt in St. Audrie’s Bay een gemengd gymnospermen bos vervangen door een monotone vegetatie voornamelijk bestaande uit Cheirolepidiaceae coniferen. Dit valt samen met een dalende palynologische diversiteit en een overgang naar een warmer en droger klimaat. Geen van beide secties laat een duidelijk massaal uitsterven van planten zien. Een compilatie van diverse T-J grens secties laat zien dat de Cheiro-lepidiaceae gedomineerde bossen voorkomen in het binnenland van Pangaea en dat er een toename in sporenplanten is in de nabijheid van de Tethys. Deze zichtbare differentiatie in de T-J palynologische datasets is het indirecte resultaat van het CAMP vulkanisme. De toename in broeikasgassen veroorzaakte een warmer klimaat en een toenemend thermaal contrast tussen land en zee. Hierdoor werd het moesson systeem sterker, het binnenland van Pangaea droger en nam de regenval toe aan de rand van de Tethys.

N.B. De hoofdstukken uit dit proefschrift zijn/worden gepubliceerd in verschillende wetenschappelijke tijdschriften waardoor enige herhaling onvermijdelijk is. De datasets waarop dit proefschrift gebaseerd is kunnen worden opgevraagd bij de auteur.

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Verklarende woordenlijst:

Black shale: donkere organisch rijke sedimenten Cyclostratigrafie: de studie van cyclische patronen in gesteentelagen Excentriciteit: mate voor de afwijking van de aardbaan ten opzichte van een cirkel GSSP: Global boundary Stratotype Section and Point, een internationaal overeengekomen stratigrafische grens tussen twee tijdperken Gymnospermen: naaktzadigen, planten waarvan de zaden onbedekt zijn Isotopen: atomen van hetzelfde element (bv. koolstof: C) met een verschillend aantal neutronen in de kern waardoor ze een verschillende massa (12C en 13C) hebben Massa extinctie: het massale uitsterven van organismen in een relatief korte tijdsduur Palaeo: van het Griekse ‘palaios’: oud Palaeomagnetisme: de studie van het aardmagnetisch veld zoals vastgelegd in gesteenten Palynologie: de studie van pollen, sporen, en bepaalde andere microfossielen Palynomorf: microscopische organische resten die overblijven na het oplossen van sediment met behulp van verschillende zuren Pangaea: een supercontinent dat ontstond tijdens het Perm, en uiteen begon te vallen in de continenten Gondwana (in het Zuiden) en Laurazië (in het Noorden) tijdens de laat Trias Precessie: de variatie in de stand van de aardas Proxy: een meetbare grootheid die gebruikt kan worden om niet-meetbare gegevens uit het verleden te reconstrueren (bv. jaarringen van bomen om temperatuurveranderingen te reconstrueren) Sectie: opeenvolging van gesteenten Sediment: afzettingsgesteente Sporomorf: pollenkorrels en sporen van planten Stratificatie: gelaagdheid in de waterkolom door verschillen in zoutgehalte en/of temperatuur over de waterkolom Stratigrafie: het beschrijven, benoemen en correleren van gesteentelagen Stomata dichtheid: het aantal huidmondjes op een bepaald bladoppervlak Stomata-index: het percentage huidmondjes ten opzichte van alle epidermis cellen Tethys: een niet meer bestaande oceaan die zich bevond tussen de oude continenten Gond- wana (in het Zuiden) en Laurazië (in het Noorden).

193 Dankwoord

Dankwoord

En dan is het nu toch echt tijd voor het dankwoord… De vier jaar die ik aan mijn proefschrift heb gewerkt gingen als een speer voorbij. Ieder jaar leek korter te duren dan het vorige. En nu zijn er nog maar een paar weken over… Maar laat ik bij het ‘begin’ beginnen. In het 4e jaar van mijn studie Aardwetenschappen kwam ik voor het eerst in aanraking met Trias-Jura palynologie tijdens mijn doctoraalonderzoek bij Wolfram Kürschner. En er is van alles onder de microscoop voorbij gekomen: hamburgers, pannenkoeken, Mickey Mouse, hoedjes, ufo’s, flapjes, en bolletjes. Door dit onderzoek besloot ik twee jaar later te beginnen aan een promotieonderzoek binnen het ‘CoBi’ project: ‘Earth’s and life’s history: from core to biosphere’. Wolfram en Wout Krijgsman, bedankt voor het opzetten van dit project. Wout, bedankt voor de samenwerking in het veld en voor de wetenschappelijke discussies. Andy Lotter en Cor Langereis, bedankt dat ook jullie me de kans gaven om aan dit project deel te nemen. Ik wil Andy, mijn promotor, ook graag bedanken voor de gesprekken in het begin van mijn promo- tietraject en voor het vertrouwen dat alles goed zou komen. Wolfram, mijn co-promotor, bedankt dat je vele uren met me door de microscoop hebt gekeken en me enthousiast hebt gemaakt voor het palynologie vak. Ook bedankt voor alle dagelijkse hulp tijdens mijn promo- tie, de leuke veldwerken en congressen, en voor het opbeuren als ik soms te zenuwachtig werd. Een belangrijk onderdeel tijdens mijn onderzoek waren de veldwerken. Deze waren nooit zo leuk en succesvol geweest zonder mijn ‘CoBi partners in crime’: Martijn en Micha. Ik heb er veel goede herinneringen aan! Martijn, bedankt voor alle gezelligheid, wetenschappelijke gesprekken, en minder wetenschappelijke gesprekken in de afgelopen tijd. Micha, jij bent vier jaar lang mijn kamergenoot geweest. En dat niet alleen, ook discussiepartner, proofreader, mede-auteur, congresmaatje, psycholoog, straks paranimf, en bovenal een goede vriend. Bedankt voor alles! The Austrian fieldtrips would have been far more difficult without the help of Leopold Krystyn, Axel von Hillebrandt, their colleagues and students. Leo, thank you for the collaboration, the scientific discussions and your hospitality in Vienna. I want to thank Michael Szurlies for his help during the Austrian fieldtrip. In Germany, Stefan Schmeissner showed us ‘the place to be’ for collecting fossil leaves. Stefan and Günter Dütsch also provided us with very important Lepidopteris leaves. Vielen Dank! I want to thank Mark Hounslow for his guidance in the field during one of the UK fieldtrips. Steve Hesselbo, thank you for providing the samples which I used for the St. Audrie’s Bay palynological dataset. Als ik niet op veldwerk was, liep ik ergens rond bij het Laboratorium voor Palaeobotanie en Palynologie (beter bekend als Pal en Pal, en wie zijn nu toch die Paul en Paul?). Hier heb ik een erg goede tijd gehad, mede door de leuke en hechte groep AiO’s. Judith, jij bent nog een tijd kamergenoot geweest, bedankt voor je luisterend oor en gezelligheid. Katya, you were my officemate during the most busy time. Thanks for your happy face each morning, the nice conversations, and the (Russian) chocolate, which always came exactly at the right moment!

194 Dankwoord

Frederike, Maarten, Emi, Peter S., Peter B., Sander, Emmy, Gianluca, Wade, Qing and Luke: we did a lot of nice things together: table tennis, squash, walking in the gardens, dancing, whisky tasting, sushidinners, chocolatedinners…. I will miss you! Ook alle andere (ex) Pal en Pal collega’s wil ik bedanken: Rike, Henk, Appy, Timme, Francesca, Jeroen, Oliver, Walter, Boris, Johan, Roel, Marjolein, Leonard, Jan, Natasja, Ton, Hans, en Zwier, bedankt voor de goede sfeer en de hulp op allerlei vlakken. Henk Visscher heeft veel tijd gestoken in het doorlezen van de manuscripten, erg bedankt daarvoor. Han van Konijnenburg - van Cittert is nauw betrokken geweest bij het huidmondjes gedeelte van mijn onderzoek, bedankt voor alle hulp! De ‘Trias-Jura’ studenten: Tammo, Martin, Maurits, en Mischa, ik vond het leuk om met jullie door de microscoop te kijken, en het is leuk om te zien dat de Trias-Jura palynologie nog wel even door gaat! I also want to thank the colleagues from Paleomagnetism, Stratigraphy and Paleontology, and Geochemistry for the nice chats from time to time. En dan is er natuurlijk nog een leven naast de universiteit. Allereerst wil ik graag mijn lieve ouders, Appie en Tilly, bedanken voor alle steun en fijne weekenden. Van jullie heb ik mijn interesse voor de natuur meegekregen en daar begon het eigenlijk mee… Ook Katrijn (grote zus), Marijn, Rosa en Margot, bedankt voor het altijd vragen hoe het ging! Joost, Olga, Anna, Nick, en Rosa, bedankt voor de steun, en de ontspannende weken in februari. En Nick, super bedankt voor het mooie ontwerp van het boek! Yvette en Kim, heel erg bedankt voor alle fijne gesprekken en lieve kaartjes. Nu zien jullie eindelijk het eindresultaat van wat ik nu heb gedaan al die tijd, wat soms toch een ver-van-jullie-bed show was. Fijn dat jullie altijd op de hoogte wilden blijven. Nelly en Hanna, van het ‘oude’ biogeologieclubje, bedankt voor de interesse en gezelligheid. Thibault and Cécile: it’s good to have friends in Friesland, let the frost begin! Marijn(tje), we hebben heel wat restaurantjes van binnen gezien waar we elkaar op de hoogte houden, hopelijk blijven we dat nog lang doen. Limke, bedankt voor de etentjes en voor het gezellige en relaxte weekendje Ardennen op het hoogtepunt van de drukte, dat deed me goed! Alle (ex)Himalaya’s, bedankt voor de leuke avondjes en uitjes! Anja, ik vond het erg fijn om zo’n goede vriendin dicht in de buurt te hebben bij wie ik altijd binnen kon lopen, bedankt voor alle wandelingen, gesprekken, theepauzes, bakavondjes en dat je mijn paranimf wil zijn. Joeri Lief hebber… Lieve Joeri, ik ben erg blij met jou als vriend. We hebben het altijd leuk samen, en we hebben veel mooie dingen gezien de afgelopen vier jaar. Niet alleen op vakantie, maar ook thuis in de Himalaya is het altijd fijn en daar wil ik je voor bedanken. Bedankt dat je er altijd voor me bent!

195 Curriculum vitae

Curriculum Vitae

Nina Rosa Bonis werd geboren op 13 februari 1983 te Oosterhout, Noord-Brabant. In 2000 behaalde zij haar vwo-diploma aan het Cambreur College te Dongen waarna ze in datzelfde jaar Aardwetenschappen ging studeren aan de Universiteit Utrecht. Na het eerste jaar volgde ze de studierichting Biogeologie. Haar doctoraal onderzoek bij de vakgroep Palaeoecologie werd begeleid door Dr. Wolfram M. Kürschner en betrof de palynologie van de Trias-Jura overgang in de Tiefengraben sectie in de Noordelijke Kalkalpen in Oostenrijk. Hierna ging ze een stage doen bij TNO Bouw en Ondergrond onder begeleiding van Dr. Frans Bunnik. Het onderwerp van dit onderzoek was ‘palaeomilieuaspecten van de Maas: klimaatgestuurde veranderingen in vegetatie en rivierdynamiek’. In 2005 studeerde ze ‘met genoegen’ af. Van januari 2006 tot en met december 2009 was zij werkzaam als promovenda bij de vakgroep Palaeoecologie (Laboratorium voor Palaeobotanie en Palynologie), Universiteit Utrecht. Haar onderzoek valt binnen het multidisciplinaire project ‘Earth’s and life’s history: from core to biosphere (CoBi)’, geïnitieerd door Dr. Wolfram M. Kürschner en Dr. Wout Krijgsman, en gefinancierd door het ‘High Potential’ programma van de Universiteit Utrecht. Haar promotor is Prof. dr. André F. Lotter, en co-promotor is Dr. Wolfram M. Kürschner. De resultaten van haar promotieonderzoek zijn beschreven in dit proefschrift.

196 Publications

Publications

Bonis, N.R., Ruhl, M. and Kürschner, W.M. (in press) Climate change driven black shale deposition during the end-Triassic in the western Tethys. Palaeogeography, Palaeoclimatology, Palaeoecology Bonis, N.R., Kürschner, W.M. and Krystyn, L. (2009) A detailed palynological study of the Triassic-Jurassic transition in key sections of the Eiberg Basin (Northern Calcareous Alps, Austria). Review of Palaeobotany and Palynology 156, 376-400 Kürschner, W.M., Bonis, N.R. and Krystyn, L. (2007) Carbon-isotope stratigraphy and palynostratigraphy of the Triassic-Jurassic transition in the Tiefengraben section - Northern Calcareous Alps (Austria). Palaeogeography, Palaeoclimatology, Palaeoecology 244, 257-280

Hillebrandt, A. v., Krystyn, L., Kürschner, W. M., Bonis, N. R., Ruhl. N. R. and Urlichs, M. with contributions by Bown, P., Kment, K., McRoberts, C., Simms, M. and Tomašových, A. (2009) A candidate GSSP for the base of the Jurassic in the Northern Calcareous Alps (Kuhjoch section; Karwendel Mountains, Tyrol, Austria. International Subcommission on Jurassic Stratigraphy, Triassic/Jurassic Boundary Working Group Ballot 2008, 44 pp.

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Depth Sample score Sample score A 13 (cm) δ Corg [‰] axis 1: humidity axis 2: T Spores Plant groups 70

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Cheirolepidiacean Seedferns Mosses & Liverworts Ferns conifers Plant groups Other spore Other conifers Other Gymnosperms Fern allies producing plants

13 Figure 3: Chronostratigraphy, lithostratigraphy, δ Corg data (Ruhl et al., 2009), PCA sample scores for axis 1 and 2, relative abundance of spores and the relative abundance of plant groups of Kuhjoch (a) and Hochalplgraben (b). The grey shaded band emphasizes the most striking changes.

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Figure 2: n-alkane biomarker C-isotope and climate proxy-records. The combined n-C25 to n-C35 odd-chain-length (green) and n-C17 13 to n-C23 (blue) δ Cn-alkane signature from the end-Triassic mass extinction interval in Kuhjoch and Hochalplgraben. The 5-6‰ negative CIE coincides with strong relative warming and enhanced hydrological cycling based on statistical analysis of palynological data.

13 13 Supplementary Figure 2 (right): Individual high molecular weight odd-carbon numbered δ C n-alkane signatures and δ CTOC signatures (in grey) from (a) Hochalplgraben and (b) Kuhjoch. Individual low to middle molecular weight odd-carbon numbered δ13C 13 n-alkane signatures and δ CTOC signatures (in grey) from (c) Hochalplgraben and (d) Kuhjoch. The calculated Average Chain Length (ACL) compared to Carbon Preference Index (CPI) values from (e) Hochalplgraben and (f) Kuhjoch. The C-isotope composition of individual n-alkanes per sample from (g) Hochalplgraben and (h) Kuhjoch.

206 color figure ch 4 13 13 δ Cn-alkanes δ Cn-alkanes -40-39 -38-37-36 -35 -34-33-32 -31-30-29-28-27 -26-25-24 -38 -37-36 -35 -34 -33-32 -31 -30-29 -28 -27-26 -25 -24 280 40 n-C25 n-C25 n-C 270 27 30 n-C27 n-C29 n-C29 n-C 260 31 20 n-C31 n-C33 n-C33 n-C35 n-C35 250 13 10 13 δ CTOC δ CTOC

240 0

230 -10 Stratigraphic position (cm)

240 0

270 30

260 20

250 10

240 0

230 -10 Stratigraphic position (cm) 220 -20 E F 210 -30 26 27 28 29 30 26 27 28 29

ACLC25-C33 ACLC25-C33

-25 -26 -27 -28 -29 -30 -31 -32

-33 -34

C n -alkanes per sample -36 -35 13 δ -38 -37 G H -40 -39 17 19 21 23 25 27 29 31 33 35 17 19 21 23 25 27 29 31 33 35 n-alkane c-chainlength n-alkane c-chainlength

Color figure ch 5

Peak

p S 0 5 3 2 1 4b 4a 0

s 0 h

p 0 r

o 4

l 0

a 0 p

l 2

o T 0 0 0 5

/g dry sediment]

3 palynomorphs

Aquatic 2

estrial palynomorphs estrial

e T 0 5 1 0

s 5 e

p S 0 0 0 3 Concentration of palynomorphs [grains x 10 0 0 2

0 n

l 1

o P

e r o 0

0 p S 1 0 8 0

p r 0

o 4

n Mudcracks

o concentrations Shell e

l n Percentage [%]

P 2 a

c i

a 0

u 0

q A 1

h 0

p r 8

o 0

n 6

l Dolostone 'Beef' calcite

a 4

i r t s

Percentage [%]

e r

r 2 e T 5 2 - Sand Limestone 7 [‰] 2 - org

C Carbon isotopes Carbon 9 13 2 - δ δ 1 3 - fall

SB TST & sea-level change sea-level & HST? RSLC

FSST-LST Sequence stratigraphy Sequence TST Dark grey mudstone Dark grey shale laminated Dark grey

rise

Lithology Height (m) Height

9 8 7 6 5 4 3 2 1 0

18 13 12 11 10 28 27 26 25 24 23 22 21 20 19 17 16 15 14 Palynological sample sample Palynological 2

189 143 135 134 133 132 131 130 129 127 125 123 122 121 120 119 118 116 115 114 113 112 111 110 109 108 107 106 105 98 87 50 43 34 23 14 80 70

290 273 263 254 239 229 213 206 201 197 194 192 168 165 161 159 155 151 149 191 141 102 179 170 Bed Number Bed

– 8 7 6 5 4 3 2 3 6 7 9 1 2 4 5 8 16

18 22 23 24 25

42 29 25 12 11 12 13 15 16 17 10

Williton Mr Williton Member Member

Blue Lias Formation Lias Blue Westbury Formation Westbury Langport Cotham

Formation Lithostratigraphy

Blue Anchor Anchor Blue Lilstock Fm Lilstock marl grey Pale mudstone Medium grey

Rhaetian Hettangian Stage 208

color figure ch 5

zones

Palynomorph assemblage Palynomorph

n s

. i SAB2 SAB3

Spore peak number peak Spore

p s

u c s

p s

u 5 3 2

estrial palynomorph sum [n] sum palynomorph estrial

a 4b 4a r

s e

s i

e m

t e

2 1 2 8 1 6 1 1 8 4 5 0 7 2 6 6 6 9 4 2 6

7 6 9

i e

r 5 8 4 3 7 6 1 7 9 3 9 1 1 0 9 6 9 2 0 4 7 T 3 2 5 t

i 3 2 3 3 2 2 3 1 3 3 3 2 3 3 2 2 2 3 4 2 3

r o

p s o

p s u c s Other Spores Other

n n

i t

r r

p s y

a e

n s

h c

s c

p s

e . s

r b

r T

p s e

o o t r

p s

o r C

p i

e p s

o i

S t

p s

r T

r i v o

p s a c C

t n

a o

u c

i r

g 0

i s s i

e 0 r

i 0 s r

u o 4 i

p s

a v

o 0 i

s 2

s t

p t l

i 0 a r

l t p s

4 o

u c

r t

i s

o n 0 n

e 2 p s

u a c

n t

d s

d i

t g 0

t i

n 2

o o

r p s a

e 0

i c

c 2

e i

p s

e c 0 o

s 2 o

p s u

P o

s 0 i

i a

l 2

a m a

s t s

t u

i 0 t

a r

n 2

e u

Polypodiisporites polymicroforatus Polypodiisporites

s t

s o a

i c

u c

p s

Other Pollen Other

s e

t u t

r e n

r o

a c

e P

u h

i r

p s

. s

A e

p l

a t

p s e

i r

p s

n s o

m s

s u

n p s

a t

t s

u s s

i C p

h u

n n

p i

u l

u c

e o

i a

l d

l a

l p

a c y

r s

d s

e r

o C

e s s

e p

u i

P i t l

o i l

p s

i i

p s

t i

n s

l e

r l p

l u c t

i i

u c

n V

i t p

a r

e e

r p

i s

l n

u s

s T e v

a u l u

t j r

l G p

a t

o e

e 0 t

i d a

r 2 u

e s

p s

p 0

t 2 l i

V o

l 0

a v 2 O

u s

r 0

t 6

i l

l 0

o 4 p

0 o s s

a 2 l C 0 0

1 a

a 0

after Hesselbo et i

r 8

org y

e 0

m 6

i l

l 0 13

o 4 p δ

0 o s s

a 2 l

C e

o 0

p 0 S 1 0 8 0 6 0

e ercentage [%] 0 l

l 2 o P P

3 0 Aquatic palynomorphs Aquatic 0

errestrial palynomorphs errestrial 0 T 1

0 /g dry sediment]

ollen 3 P 8 6

4 Spores 2 0 Concentration [grains x 10 4 2 - 5 2 - [%] 6 2 - org 7

C - Carbon isotopes Carbon 8 2 Initial shift 13 - δ 9 2 - 0 3 - record and the lithology from the spore peak interval in the Lilstock org Chronostratigraphy, lithology, sea-level changes and and sea-levellithology,Chronostratigraphy, changes C 13 δ

14 13 12 Palynological sample Palynological 98

130 129 127 125 123 122 121 120 119 118 116 115 114 113 112 111 110 109 108 107 106 105 102

Enlargement oftherelative terrestrial palynomorph abundances concentrationsand in

Bed number Bed 2 3 4

Cotham Mbr Cotham Langport Mbr Langport

Lithostratigraphy Lilstock Fm Lilstock Figure 4: the to relation Figure 2 (above): Formation.The shaded portion ofthecurves times5 ais exaggeration oftheabundances plotted in 2).(seeFig. numbered areSpore peaks black. al. (2004), palynological sample positions, terrestrial:aquatic palynomorph ratio, pollen:spore ratio, palynomorph terrestrial:aquatic positions, sample (2004), palynological al. ratioconcentrationsand ofpalynomorphs through theTriassic–Jurassic transition theinSt. (1-5). number a and band shaded a with indicated section.areSporepeaks Bay Audrie’s

209 Color figure ch 5

with 470cm periodicity n

1.6E+12 a 2600 90% confidence interval i g

n 2500

1.4E+12 a t t 2400 e

1.2E+12 H 2300

1.E+12 n 2200 o i t 8.E+11 a 2100 m Relative power r

o 2000 F

6.E+11 s 1900 a i L 4.E+11

e 1800 u l

2.E+11 B 1700 1600 0.E+00 0.000 0.004 0.008 0.012 0.016 0.020 1500 t r r e o b p 1400 m g m F

Cycles/cm e n

a M k L c B 80000 r e o 1300 m t b a s h l m t i e Power spectrum: Spore abundance [%] o L C M 1200 Precession peak, with 100cm periodicity n a

70000 i

90% confidence interval t 1100 e a

60000 h 1000 R n

o 900 i t 50000 a 800 m r

o 700 F 40000 y r 600 u b t

s 500

Relative power 30000 e

W 400 20000 300 r 200 r

10000 M

n o n o h i o c t t i

l 100 a n l i A

m W r e

0 o

u 0 l F 0.00 0.04 0.08 0.12 0.16 0.20 B -31 -29 -27 -25 0 100 200 0 20 40 60 80 100 -1.0 0.0 1.0 2.0 3 Cycles/cm δ 13Corg [‰] [grains x 10 /g dry sediment] Percentage [%] - Humidity +

Figure 7: Power spectra of (a) the terrestrial palynomorph concentration with a main periodicity of ~470 cm and (b) the relative spore abundance with main periodicity of ~100cm. Gaussian band-pass filters reflect periodic changes in both proxy records (c) and are tentatively assigned to the astronomical ~100 kyr eccentricity and ~20kyr precession cycles.

Figure 2 (right): a) Lithology of the Wüstenwelsberg section, position of the fossil levels and position of the palynology samples. For practical reason, three different locations within a few meters distance from each other were chosen for sampling. The base of each location is indicated with a 0 in the lithological column. The asterisk indicates the position where a hole was excavated to sample the lowermost clay layer with abundant Lepidopteris ottonis leaves. b) Detail of level 2 with the position of all the layers.

210 color figure ch 6 B A * k2b wz57 k2o k2a wz53 wz58 k2c level 2 detail level

9 8 7 6 ‘Hauptton’ Fine sands Fine sands Medium grey shales Dark grey shales Light grey shales Sandstone

Lithology Fossil leaves levels leaves Fossil 3 2 1 Palynology sample 1a *

4 3 2 1 0 6 5 4 3 2 1 0

6 5 4 3 2 1 9 8 7 0 0 2 1 1 1 1 1 1

Depth (m)

‘Hauptton’ ‘Hauptsandstein’

Triletes beds Triletes Contorta beds Contorta Postera beds? Postera

Jurassic Triassic transition interval transition

211 Color figure ch 6

cm cm A B

C D

50 μm 50 μm

E F

50 μm 50 μm

Figure 3: a) A Lepidopteris ottonis leaf from level 1. b) Fragments of Ginkgoites taeniatus leaves from level 3. c) abaxial cuticle surface of Lepidopteris ottonis from level 1, slide 1-A-A. d) abaxial cuticle surface of Ginkgoites taeniatus from level 3, slide Gi-3-A. e) adaxial cuticle surface of Lepidopteris ottonis from level 1, slide 1-A-A. f) adaxial cuticle surface of Ginkgoites taeniatus from level 3, slide Gi-3-A.

212 color figure ch 6

3.5 A B 8

7 3.0

6 2.5 SI [%] SI [%]

2.0 4

3 1.5 0 2 4 6 8 10 12 14 16 18 0 2 4 6 8 10 12 14 16 Field of view Field of view 1-A-A 1-B-A 1a-A-A WZ53-A-A WZ53-B-A Gi-3-A Gi-3-B Gi-3-Ca Gi-3-D k2a-C-B WZ57-B-A k2b-D-A 3-A-A

Figure 5: a) Cumulative mean SI of Lepidopteris ottonis b) Cumulative mean SI of Ginkgoites taeniatus

213 Color figure ch 6

Level 3 140 1 1a k2c w z58 SI w z53 k2a k2o w z57 k2b 130 k2b 3 SI simulation 1 constant ED wz57 120 SI simulation 2 k2o constant SD 110

k2a ] 100 2 m m wz53 / 90 [ n

wz58 S 80

k2c 70

1a 60

1 50 A B R2 = 0.7364 40 3 4 5 6 7 8 800 1000 1200 1400 1600 1800 2000 SI [%] ED [n/mm 2]

7.0 7.0

6.5 6.5

6.0 6.0

5.5 5.5 ] ] [ % [ %

I 5.0

I 5.0 S S

4.5 4.5

4.0 4.0

3.5 3.5 2 C R = 0.6257 D R2 = 0.1416 3.0 3.0 50 70 90 110 130 150 800 1000 1200 1400 1600 1800 2000 SD [n/mm2] ED [n/mm 2]

Figure 9: a) simulated SI values with a constant ED and a constant SD compared to the original SI values b) correlation between SD and ED for each leaf within a fossil level c) correlation between SI and SD for each leaf within a fossil level d) correlation between SI and ED for each leaf within a fossil level

214 color figure ch 6

Depth (m) Fossil leaves level 3

13 Triletes beds Triletes

8 k2b wz57 k2o k2a 7 wz53 Contorta beds k2cwz58 6 Triassic 5

1 1a

0 2

Postera beds? SI based CO2 reconstruction 1 1 SI/SD based CO2 reconstruction

0 1500 2000 2500 3000

CO2 [ppmv]

Figure 10: atmospheric CO2 changes in the Wüstenwelsberg section based on the SI value of Ginkgoites taeniatus from the uppermost level 3 and on the corrected Lepidopteris ottonis SI (SI

‘Ginkgoites’, Table 3a) values from the other levels. The grey band shows the general CO2 trend.

215 Color figure ch 7

A: Latest Rhaetian

Classopollis dominated assemblage ? spore dominated assemblage bisaccate dominated assemblage mixed assemblage ? no clear stratigraphic framework ? ? Tethys ocean

CAMP

? ?

B: Earliest Hettangian

? Classopollis dominated assemblage spore dominated assemblage bisaccate dominated assemblage mixed assemblage ? ? ? no clear stratigraphic framework ? Tethys ocean

CAMP

? ?

Cratonic landmasses Marginal marine - Fluviolacustrine Deep ocean

Figure 8: Position of sporomorph assemblages in a) the latest Rhaetian and b) the earliest Hettangian. (Map modified from Quan et al., 2008)

216 CHAPTER 1 FIGURES

Pollen Spores

e s s a s is si l si s si l n n en e ssu tu s era n s ve s a e n t s e su s u e s s g s t se la ri b s s s m ma j i s s rmi s is s s vi s (R) la o l s la b tu tu rmi o b i u e u ra s . si d s i s tu tu s a e su l u e a s s cu co a s o d s s s n rt s fo s t ri u p n u i p a tu icu t e d o cu u ico g u ii l ri a tri l f a u tu + i (R) e id li e la e co t e ru t d cu l a t la la i in ti a b icu n l rg ns s u a i u su e n p la s icu icu so i p to a mu . g s n a . s sp n a la ti v. a u g e s sa a s a g l e su g ca n a a s u se u t d g a a ra ll i p rcu n a g s p iv. r u s o s g b i t g e a cu e co sp o re s s zo ri e u g ru u a o n co t ri ro p a id n e e s (R)u . l st e n s . scu s vi d e ri e d te d iv. e d rscri li t ia l s te e t t jb . imu g h ci . ri ci ra n ] su g v. o ll a i u . s . e u n sp e p ssi u e e ssu sp t g ri vi n d a . v. e yra ru ri rh a ri il . rvi ri t so s ri e a p le n ot e p u s ig e n ye s i d a n tro d yt p e p sp a a mi s b i f t fi . n va ri n o e . rh i t s s n o a icu te o ri o n u e cro irre g h e b rru s ro st te d u rma . rh i ci a t s s s scu tu sp ri la s sp te s o o la lu sp d in g e va e s sp s (R)p ll ri o te g sp mi sp sp t t rg t me to re i . p e s rt s lu s sp ri sp te s a e fu te re s o s te l sp s s sp s . it s it ca sp te e o o sp me e sp s in s mi s s s s ri s ve a n n se s g sp e te l e o i e in it ri s ra s te s e ri s te sp o e te s s ra d e d o ri s t s f o tisp s l te l s te s e ri s e e o te s t u is is is le sp p e s t ma i e it s s ccu n t l n ra t o e o e ri sp ri t o e ra e ti ra n te t s a te sp te o t o e le e sp n tri si tri te ri e s t o ri t t ri es e e ll ll ll l s t li ra ri id p e te sp e le ci e t t o e o ri t o l a o ri ri te n ri s ci ri ci ra n t ri t ra e la ori ri i t te ri o ri ri isp o t t co o o o o lis ri l o a i i to it l a cu lle o d sp ri sp ri o ri tri l o o o e ori sp ia o ia o ri t ri o u co o id o ri o sp o o le ri d l p te l o o o ll s b p p o ri e n o o o ra sp isp o sp o sp zo ri t to zo sp o o o ra zo g sp mo ra p o ri ti sp o in a p p p e ri o ip sp e te a ri o tysa p o sp i vi sp t cu lli ro sp sp o lizo sp le sp od sp d sp to sp ro n o rru isp te sp o sp sp sp ska sp tri t a o p sp t isp ti n mmi i o d d a o ca ra n isp id sp ysp co mi la isp th a o i re p po ro i sp sp li co o le sp re mi o o l si o sp s o g li i e d a e n t e sma n d sp ca e ca o o n o so mo n b ra isp u tri ra yo in ri ycl ri u n g ch ve so yp o ri sp so o n a h e re t sso sso sso sp re a a n g co ri a h a Pl ri u a u si ll n lt li ch a n cu rn la n rce ma b d g re ti b ymo ma t C t n a ma lyci la o n n t pt smo t ri n mi ve ri g ve p o la la la su i va t h ri lo ri h yca . u n o icci e e ra ri e a a o a e mb o in e e e h a . ri a e o e a o ra ig e o o e va

u a i yco yco e o Local palynomorph assemblages Regional zonation Kürschner et al. (2007) Depth (cm) C C C T Al O Vi R T L L Eu F St Ep C C cf Pe Ara Q Pi Ve CerebropollenitesPo Polypodiisporites thiergartiiC R polymicroforatusD H T T PolypodiisporitesKyrt D B ipsviciensisa C C Aca C Po C Z L T C St R Z Asse L Isch L T C Ech Ara cf Ara An T C Po Se N C D Pl L C A T D Se Kyrt F Pe R N L U Sp Terrestrial palynomorphC suONISS [ 1800

1700

1600 H4b TPi 1500

1400 b 1300

1200 Jurassic 1100

1000 H4a

900 TH Kendlbach Formation Tiefengraben Member Tiefengraben 800

700 a H3 600

500

400 H2 RPo

Schattwald beds 300 Triassic 200

100 H1 RL Eiberg Mb Kössen Fm 0 20 40 60 80 100 20 40 60 20 20 20 20 20 40 20 40 20 20 40 20 400 2 4 6 8 10 12 1416 18 20 Total sum of squares Percentage [%]

Figure 4: Relative abundances [%] of pollen and spores through the Triassic-Jurassic transition in the Hochalplgraben section, and their relationship to the palynological zonation scheme from the Tiefengraben section (Kürschner et al., 2007). RL=Rhaetipollis- Limbosporites zone, RPo=Rhaetipollis-Porcellispora zone, TH=Trachysporites-Heliosporites zone, TPi= Trachysporites-Pinuspollenites zone (note that the TPo zone of Tiefengraben is absent at Hochalplgraben). The presence of Cerebropollenites thiergartii and Ischyosporites variegatus found after qualitative analysis is indicated by dots.

Pollen Spores

e s si s la li en " s n ra ssu s i s s e te s i ra s ri su s b e a s r m i ma s rt o s s u a s u b ll s s e s s d o u s rmi a u . si t rmi a l co yi i s t i ra u tu s nsi ig s u su s ru s d rt s e fo s isp t p n ri la fo t u ri s s e d u la e ii i s co ct e s cu tu t s h a tu s e is id e rg t t la e e u e la g t n ru t u lu d co i n p tu u in s s ti la tu n a si is cu u l cu o s mu e e lu s . sp s . n g g a r n su a e la n rg a s n icu b e a rg d s li n e a s u a rp a s l n ll ti s p ra i t lla u i i d p u p o n t s e i cu ca o scu t s u s u l ri a t s sa a u e n sp a a g u g ri a o e u se a n a n n s th . n . r rcu s ri d ru n ig li tra ci ci ri ch e ra te jb b e e lo h p rscri rn b u imu a ra g e h g su o ve p i id p s s st l e a u s i vo p e scu e sp te g e n ci a s e t g ig ri e d s a t t ra . e o e t g st le rh ri ru vi mo ye d u ti n a ll s . e li u e n h mi st sp n n fu b ri va . sp n issi cro issu i a s s e li t yra s ri ru b o n n mu te ri cu o f t vi n u e cyt e o ro u j a i rh a e p t a a rma s a u te . d o sp e u s o p s o rg f n n te te ci n g e o lu rh o li g n n ro a es va s a o b l rt p it ri i e e a p s b i s sp C t s t e s te l re a s sp ig ri sp a t s s sp s mi ri s l i sp i le se b mi me s it le te l lyn me t se s s s n sp o d s g it ia te n t ra e s t sp te sp s mi s e ca le o s ra ra ri sp e e B te s o e sp fo s s s s s s s s te d s i s a s s p te s te n lin d i ro le s rme t h ra t e ri ti s ri ra s e t s s ri ri g s o o o t t si t te se t ti ri e e e s e su e e te ri te s d ri t p li li te rcu ma ri te lle e sp ra is e n l ccu e o e o o ri t o le e o o te e t ri te te ra t o te n ri ri e ri ri sp a a o t t t te t t t ci ri te ci o e l l l is ri e a o ri o it o s ci ll ll s le s o it f d lig o ri la ri t ri t ri o ri ri o isp s ri sp sp o o co o o s e n l ra ri ri ri ri ri s ri ri ri o ia o ri ia d o o ll o p ll o p p t te a o o cu e l te p p e n " sp sp o isp t ri sp sp o ri o o o mo e o si zo d te ra u o isp o o o o o te o o o o sp in a p p o lo e e o ri p e t o tysa ri d i o t cu o o vi li o sp sp o sp t to isp sp rru in isp e izo g t e sp d sp d ri p sp isp sp d ri ti op ra ri ri ro d A id ysp sp la a th l sp sp sp lska le co ro t si sp l co l n isp sp sp sp sp l sp sp isp o sp o mi s st li i u n t i a sma o ca e n o sp o b a e n n o mo i b sp ca sp e so i sp ch a omi tri isp re u ve re so o tri u l re i i isp i tri ri i p yo i p o re sso sso n e a g a u a ad u re Pl le le e lt la ch n n i n lio ri n rn la o g ti rn ve n n ti n g yci u re re ra re i t rn re smo rn [total counts n] a a va tre ra g n tre isp ri u isp n isp . h l l cci cu d rce g l ymo rru ma rn o mb a n l n mi e e b rn e t e o rre l l a u su yca h h e yp e a ra i o o o e i va e a h e o e a a o e o e i e ri i e a e a o yco a yco e ürschner et al. (2007) C C O Vi G L L Vi T C C Al Ara R Pe Q Al Pi Al C cf Ep Po Po PolypodiisporitesT polymicroforatusD C T R Ba C Aca T C Po H T U D C Va T N R Kyrt PolypodiisporitesR C A ipsviciensissse Ve C C N C Sp D L R T C Po An Se St St Z C St R Ara C St C L Isch C L Kyrt Sp T CONISS Local palynomorph assemblages Depth (cm) Regional zonation K

1900

1800

1700 1600 K4 TPi 1500

1400

1300

1200 b Jurassic 1100

1000

900

800 Kendlbach Formation Tiefengraben Member Tiefengraben 700 K3 TH

600 a

500

400

300

200

100

Schattwald K2 RPo

Triassic 0

-100 -200 K1 RL

Eiberg Mb -300 Kössen Fm

-400 20 40 60 80 100 20 20 20 20 20 20 20 20 20 40 60 20 500 2 4 6 Percentage [%] Total sum of squares

Figure 5: Relative abundances [%] of pollen and spores through the Triassic-Jurassic transition in the Kuhjoch section, and their relationship to the palynological zonation scheme from the Tiefengraben section (Kürschner et al., 2007). RL=Rhaetipollis- Limbosporites zone, RPo=Rhaetipollis-Porcellispora zone, TH=Trachysporites-Heliosporites zone, TPi= Trachysporites-Pinuspollenites zone (note that the TPo zone of Tiefengraben is absent at Kuhjoch). The presence of Cerebropollenites thiergartii and Ischyosporites variegatus found after qualitative analysis is indicated by dots. Dinof lagellate cy sts Acritarchs Prasinophy tes ii s ic v o ta a is m t ti s ic s u e s r g t j i d lu a n e o n n o p CHAPTER 1 FIGURES i a i m m e i e ly t in h i u s . ts l r g c m s ts p v o . e s t . x n a s u e s p . fa p p ia d y s p la a n ri i o y s p t . . a s d in c te s u l ia p id k c s e f r a ri ] e s l s a b r e m d c sp e ll s n t te a u ya ll a m e m t u m n s a e e te s la y r c n e u a iu la i u i u s h a y m t l s h fe c t sw i h n l id i s c te p rm h h u n e h p i o o n in p i e tr h h is i s e p p s u g c o in c g o ia d s d g s c c d n io p s o o la r n o to m s o to o la y a r o a t s io n c f ta i m y e u s c s e f h h ta h m a o e i l o ri s ra tr a a e p i lv o cr ry ri tt s m r L s ta in c ra o o h e u a le a in i e c y a y te f. ra o Depth D A P F B R B S D C V D M V A T T C P c P Aquatict (cm) [ 1800

1700

1600

1500

1400 b 1300

1200 Jurassic 1100

1000

900

Kendlbach Formation 800 Tiefengraben Member Tiefengraben

700 a

600

500

400 beds

Schattwald 300 Relative abundances [%] of aquatic palynomorphs Triassic Figure 6: 200 through the Triassic-Jurassic transition in the Hochalplgraben 100

section Eiberg Mb Kössen Fm 20 40 60 80 100 20 20 40 60 80 20 40 60 80 20 40 20 40 60 80 100 20 20 40 60 80 100 400 800 Percentage [%]

Dinoflagellate cysts Acritarchs Prasinophytes

ii s ic v ] o a n a is it s ic s m rt ts t j iu a n g e o n p u in a m m e y o in h u ii s l c ts l r c m g type As o l s t x s a u n e p a a y s a ri n i a la o t a i t c te l p ia id l l k e r a sp. id to e s l s sp. u b r a te m spp. d e ll r [ t te a u ya m a e ll n m u n a s cf. faveoluse e la y r c n u a e o iu i i h u s sp. a m l s h fe c i sw h t n id s p c rm te h u e h p i o o in p n m i tr h s is e i p s g c o in c g d ia s o u d s c io d p n s c la r n o to o s to m a o y r t o s a io ti f ta i m y e c s s u e e h ta a h o m e a o ri s ra tr a p e i a lv cr ri m tt r s L u in c ra o o h a u le e a i c y y te a f. q Depth D A P F B R D S C B cf. B V M A C T P T c A (cm) 1900

1800

1700

1600

1500

1400 1300

1200 b Jurassic 1100

1000

900

800 Kendlbach Formation Tiefengraben Member Tiefengraben 700

600 a 500

400

300

200

Schatt wald 100

Triassic 0

-100

Figure 7: Relative abundances [%] of aquatic palynomorphs -200

Eiberg Mb -300

through the Triassic-Jurassic transition in the Kuhjoch section Kössen Fm -400 20 40 60 80 100 20 20 40 20 20 40 60 80 20 40 60 20 40 60 80 100 200 600 1000 Percentage [%] CHAPTER 6 FIGURE

Pollen Spores e s la tu u a Ricciisporites) is s r s s . s fo is u cl m is a s o s s s s s tu a s r d m u s s r u is s u n o a is i x a it u s fo u o u u c s n . u t u e ic l s y (e c t t s s e s e r d r s cc s rt s i o i p s s t s la ri s ii c s i tr u n e ti s r la e is tu id lis a t u a is u jo u a u e o m in m s p u u la s u u e tu d ia u ic a tic ns e n m e lu pa r s o n c s ll l a s c s l . . r ly a e s c ic u s g g a a g at s s v e n s e u o cu o n la t ra ia s e i u e ol se im o ti u p u p p e o p r d is t g su o u n g bl in s g u u s ta ic a re a o u h s m ha e ly r e a la t r u d A n n n p s p s m in r e d a n . r p p ig p is fo i s r e u o c r i i d r u i ri m ip a c h . g d c c u r v o e sp v o e s e s in a a . n ti tu u is s ic a li s ra . p e s s n rv to te te a r b ru ro s v n u s v a i yr s t p s le g is s . c x a a p b d u d u y o d e m h p a a s s c l te m h l te g p p n pi s a s a ri n rh s b r c is e u th fis e v in s g e o p ite e n il fu p es ris a . . f i a tu n j u s a e r e p r p p l u te s ia ri s m r a ri p s te s s s te p te v o a e i e i e la l la s m te it g s r s o c sp it p l p f id r a b e ite s m to ty e a s a u b i fu d o e s s p o s s a ite te ri ri a p s it g v m r s s s s s e ri ra r is s o e l ra s r u p c r e s s s n te rm s g i s lin ic o n a p it u e s te s r ra r ri s o s o cl is s te id s s s e e e i t s o o o r e p it a te s o t m a s s e a us te n e p i is is s d rt r le ra r s n c it e sp i t o o o te p te p t te a c e e s e ite it it it r e te p p o it s r r g ri e p e y s u a h m ri le it s lle c ll ll fo ite lli s ite u e s l o s o e c r it o iid e te p po p p ri s ri s s la ri n a t t te it r r or or o ril ri s sp is p r e o o s o rit is d iu n e c h p cc u o l r lli o ia o o e p o e p c p e o p e t ll a o r t d ri s s is s o ii o ki te u o o i ile ile ri r o o p p sp t o ii o a is o a p p te p o s in in o it s p s o s o o o p r p p d o p it o e o it p s it a o s p o a m in o o o v ti p d p u e c p iz d r r o o p p s s i o p d t k t p r s lis e s p o d g n i s io c c Depth (m) sp p n s sp p o a o o n d ti r d ra ul r ae a r m p y tis sp m m n p id a la s o s l il a s l o ot ot p sp s s o o m th s o re ls re is co o l il y s c s o to a od io t o ti i r e re i li n c s s e a e o a d n o g ic o s s h a i s o e is o m c u i p ri k tr b i u p h h is o ra o n s o n ri p e a i e n m e tr h e u re . palynomorphc e m h e a y a e ll o re a ri u s s ll c a sp c a a sp u s sp a u c n re a c ll d lt la n c rn ly t ra ti n rn g o p p r li b b ri n rt a t ly s g m r a s rc ti c a rr o rr p a s tt L m tr u 6 icci th o p it v e ra la la o y h li y u r li s e li h in ra u it h u o o e a o a a o ra a e o a in yc o o ig e e im e e y c ra o s o e te ri o o e ra v e p e a h a y f. y o q R O P S V O P A C C P C R A C Q G A T V A C P B L V C E P T D C C B C P A P R C C C L L L T H Z L P D K A A P A R S S T C P R T U V S T D R T T c C B A 398 15 5

0 278 0 143 4 14 132 2 101 1 13

11 109 1

8 114 0 244 0 7 164 13

6 115 0

5 115 0

4 106 0

3 100 0 2 110 1 1 104 0 0 100 100 20 40 20 40 20 20 20 20 20 40 60 20 40 20 20 20 Percentage (%)

Figure 4: Relative abundance of terrestrial palynomorphs and the presence-absence of aquatic palynomorphs in the Wüstenwelsberg section.