Climatic and environmental dynamics during the Valanginian carbon isotope event - Evidence from

geochemistry and palynology

Ariane Kujau

Dissertation

zur Erlangung des Grades eines Doktorsder Naturwissenschaften an der Fakultät für

Geowissenschaften der Ruhr-Universität Bochum

Climatic and environmental dynamics during the Valanginian carbon isotope event

- Evidence from geochemistry and palynology

vorgelegt von

Ariane Kujaugeboren am 2. April 1983

in Nordhorn

Bochum im Juni 2012

La Charce outcrop section in SE France (own photograph).

“I hear the ancient footsteps like the motion of the seaSometimes I turn, there‘s someone there, other times it‘s only meI am hanging in the balance of the reality of manLike every sparrow falling, like every grain of sand“

Bob Dylan (1981)

Eidesstattliche Erklärung

Hiermit erkläre ich an Eides statt, dass ich die vorliegende Arbeit selbstständig angefertigt sowie die benutzten Quellen und Hilfsmittel vollständig angegeben habe. Soweit Zitate oder Abbildungen anderer Werke im Wortlaut oder dem Sinn nach entnommen wurden, wurde dieses in jedem Einzelfall als Entlehnung kenntlich gemacht. Die vorliegende Dissertation wurde in dieser oder ähnlicher Form bei keiner anderen Fakultät oder Hochschule zur Prüfung vorgelegt.

Bochum, Juni 2012 Ariane Kujau

Table of contents

I Table of contentsV AbstractIX Kurzfassung1. Introduction 1 1.1 Carbon and the Climate system 1 1.2 Cretaceous perturbations of the carbon cycle 2 1.3 Vegetation and the Climate system 5 1.4 The Climate System of the Early Cretaceous 8 1.5 Aims and objectives of this study 10 1.6 Outline of this manuscript 11 References 132. Methods 21 2.1 Field work 21 2.2 Carbon and oxygen isotope analysis 21 2.3 Organic carbon and total organic carbon content 22 2.4 Carbonate carbon and total organic carbon content 22 2.5 RockEval pyrolysis 22 2.6 Biomarker analysis 22 2.7 Palynology 23 2.8 Calcarous nannofossil biostratigraphy 23 2.9 References 233. No evidence for anoxia during the Valanginina carbon isotope event - An organic-geochemical study from the Vocontian Basin, SE France 25 Abstract 25 Keywords 25 3.1 Introduction 26 3.2 Geological setting 27 3.3 Methods 29 3.4 Results 30 3.5 Discussion of Results 36 3.6 Conclusions 42 Acknowledgements 43 References 444. Reconstructing Valanginian (Early Cretaceous) mid-latitude vegetation and climate dynamics based on spore-pollen assemblages 53 Abstract 53 Keywords 53 4.1 Introduction 54 4.2 Geological setting and stratigraphy of studied sections 55 4.3 Material and Methods 57

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4.4 Results 57 4.5 Discussion 68 4.6 Conclusions 74 Acknowledgements 75 References 805. Fluctuations in sea-level and terrestrial input at the NW Tethys and the Carpathian seaway during the Valanginina (Early Cretaceous) - Evidence from palynofacies and n-alkanes 89 5.1 Introduction 90 5.2 Geologic setting and stratigraphy of studied sections 91 5.3 Methods 93 5.4 Results 96 5.5 Discussion 98 5.6 Conclusions 102 Acknowledgements 102 References 1036. Linking changes in pCO2 with environmental and climate dynamics during the Valanginian (Early Cretaceous) 113 Abstract 113 Keywords 113 6.1 Introduction 114 6.2 Geologic setting 115 6.3 Material and Methods 115 6.4.1 The significance of the ∆δ13C record for reconstructing trends in Valanginian pCO2 117 6.4.2 Comparison with existing carbon isotope records 119 6.5.1 Climate settings durng the initiation of the CIE 121 6.5.2 Causes and consequences of changes in Valanginian pCO2 123 6.5.3 Implications for triggers of the Valanginian CIE 126 6.6 Conclusions 127 Acknowledgements 127 References 1287. Synthesis and future perspectives 135 7.1 The applicability of the chosen approach 135 7.2 A marine or terrestrial trigger for the Valanginian CIE? 135 7.3 Fluctuations in the atmospheric carbon content 136 7.4 Future perspectives 137Acknowledgements 139Appendix 141Curriculum vitae 175

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Table of contents

IV

Abstract

The Cretaceous (~145.5-65.5 Ma), previously regarded as a time period of stable environmental conditions under a greenhouse climate, is today known to have experienced recurrent severe environmental and climatic perturbations. A stage of the Early Cretaceous that only recently receives closer attention is represented by the Valanginian (~144.5-133.9 Ma). This time interval is characterized by a distinct positive carbon isotope excursion (CIE), revealing a perturbation of the global carbon cycle, accompanied by severe changes in the ocean-climate system. This phenomenon has previously been observed for a number of time intervals during the Cretaceous. The Valanginian CIE, however, was the first of these positive CIEs. A major difference to the stratigraphically younger CIEs (e.g. OAE1a, OAE1b, OAE2) is the absence of enhanced marine organic matter (OM) accumulation (e.g. in the form of black shales). This OM accumulation was assigned to the widespread establishment of marine anoxia, probably in combination with a relatively rapid intrusion of CO2 into the ocean-atmosphere system by volcanic activity. Up to now, no comparable marine carbon sinks could be identified for the Valanginian. A change in the focus of investigation towards terrestrial environments could provide new insights into the carbon cycle and may reveal major triggers for the Valanginian event. This study provides a detailed investigaition of paleoenvironmental and climatic dynamics accompanying the Valanginian CIE, with a special focus on the terrestrial realm. The chosen approach uses a combination of geochemical, chemo- and biostratigraphic, as well as palynologic investigations to reconstruct environmental and climatic change during the Valanginian CIE. It provides a more detailed understanding of the lead and lag of events within this complex interval of Earth history. The role of the terrestrial realm as well as the interplay with the marine sphere and changes therein are investigated.This is done by investigating Valanginian marine sediments from two study sites, one located within the northwestern marginal marine Tethys and another one located within a seaway representing a passage towards the Boreal realm. Studying the corresponding archives allows for an investigation of changes within marine and adjacent continental environments. The Tethyan site represents a stratigraphically well constraint setting that allows for a detailed comparison of the collected results with existing data sets from other Valanginian successions. The study site located in the seaway connecting the Tethys with the Boreal realm is to date stratigraphically not yet well constraint. Within this study, a correlation with the Tethyan site is based on bio- and chemostratigraphy.A high-resolution carbonate carbon isotope record in combination with data on carbonate and organic carbon contents is established for the Vocontian Basin (southeast France). The carbon isotope record forms the basis for a correlation with existing Valanginian records. For an investigation of the OM, biomarker data sets are established, which represent the very first for the Valanginian Tethys. Varying abundances of isoprenoids, n-alkanes, steranes, and hopanes reflect the occurrence of source organisms for these biomarkers (e.g. land-plants, dinoflagellates, and cyanobacteria). Based on these, the state of oxygenation of the depositional environment is assessed. However, except for a few cm-thick layers (the so called Barrande layers) stratigraphically located before the initiation of the CIE, the establishment of anoxic or euxinic conditions can, at least for this basin, not be affirmed.With regard to changes in terrestrial environments spore-pollen records from marine sediments are assembled,

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reflecting variations in the paleo-vegetation on adjacent continents. One record has been established for the Vocontian Basin and one for the Mid-Polish Trough (Carpathian seaway, central Poland). The vegetation adjacent to the Vocontian Basin was dominated by drought-resistant cheirolepidiaceans and that around the Mid-Polish Trough by araucarians/cupressaceans, at both sites associated with high pteridophyte abundances. A gradual trend to more humid conditions is observed towards the lower/upper Valanginian boundary, which represents the interval of the initiation of the positive CIE, at the Vocontan Basin interrupted by a short-termed drying. Following this phase of locally different vegetation patterns, a supra-regional phase of particularly humid conditions is observed. Subsequently, palynological data point to the establishment of comparatively arid conditions at the Vocontian Basin. A reconstruction of changes in sea-level and terrestrial input for both sites is done based on changes in the abundance and distribution of the palynofacies. The palynofacies results are flanked by data sets on n-alkane ratios interpreted to reflect changes in land-plant input. The resulting data point to the establishment of a high sea-level around the lower/upper Valanginian boundary interval, followed by a phase of sea-level lowering, persisting until the earliest Hauterivian.Finally, the importance of fluctuations in atmospheric pCO2 during the Valanginian CIE is assessed based on a record of relative pCO2 changes. Here, an approach comparing changes in carbonate and in organic carbon isotope records is used. The results reveal an increasing trend for atmospheric pCO2 accompanying the initiation of the CIE, which is followed by a rapid drawdown in combination with a proposed cooling interval. This phase of pCO2-drawdown is characterized by rather humid conditions in mid-latitude settings, probably associated with enhanced continental carbon storage. Continental carbon storage is thereby highlighted as an important trigger for the Valanginian CIE.The multi-proxy approach used to decipher the nature of the Valanginian CIE allows the conclusion that marine carbon storage was most probably of only minor importance for the observed perturbations in the carbon cycle. Changes on continents in the form of moisture variations can be correlated with the CIE. This can be seen as an indication for a terrestrial carbon storage as an important trigger for the Valanginian CIE. This study provides an important step towards a more detailed understanding of the Valanginian CIE beyond changes in the marine realm.

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Abstract

VIII

Kurzfassung

Die Kreidezeit (~145.5-65.5 Mio) wurde lange als ein Zeitintervall angesehen, das von einem Treibhausklima und stabilen Umweltbedingungen gekennzeichnet war. Heute ist jedoch bekannt, dass die Kreide von wiederkehrenden enormen Umwälzungen in Klima und Umwelt geprägt war. Ein Abschnitt der frühen Kreidezeit, der erst in letzter Zeit gesteigerte Aufmerksamkeit erfährt, ist das Valangin (~144.5-133.9 Mio.) Das Valangin zeichnet sich durch eine positive Kohlenstoffisotopenexkursion (CIE) aus, die auf eine Störung im Kohlenstoffkreislauf hindeutet, welche von starken Veränderungen im Ozean-Klima System begleitet wurde. Dieses Phänomen ist für zahlreiche Abschnitte der Kreidezeit bekannt, wobei die CIE im Valangin die früheste derartige Störung darstellt. Die CIE im Valangin unterscheidet sich von jüngeren Exkursionen (z.B. OAE1a, OAE1b, OAE2) dadurch, dass im Valangin keine erhöhte Akkumulation von organischem Material (OM) im marinem Raum erfolgte (z.B. in Form von Schwarzschiefern). Derartige OM-Akkumulationen in jüngeren Zeitscheiben wurden mit der weit verbreiteten Ausbildung von anoxischen Bedingungen im marinen Raum in Verbindung gebracht. Diese standen vermutlich im Zusammenhang mit einer relativ abrupten Intrusion von CO2 in das Ozean-Atmosphäre System, durch vulkanische Aktivität. Bis heute konnten vergleichbare marine Senken für Kohlenstoff für das Valangin nicht nachgewiesen werden.Eine Erforschung dieses Ereignisses aus kontinentaler Sicht könnte neue Einsichten in den Kohlenstoffkreislauf und Hinweise für die Ursachen der CIE im Valangin liefern. Die vorliegende Arbeit umfasst eine detaillierte Untersuchung von Paläoumwelt- und Klimadynamiken, die die CIE im Valangin begleitet haben. Ein besonderer Fokus der Arbeit liegt auf dem terrestrischen Raum. Der gewählte Ansatz beruht auf einer Kombination von Geochemie, Chemo- und Biostratigraphie, sowie von palynologischen Untersuchungen. Diese erlauben es den Umwelt- und Klimawandel während der CIE im Valangin zu rekonstruieren und damit ein verbessertes Verständnis von der Reihenfolge der auftretenden Veränderungen während dieses komplexen Intervalls der Erdgeschichte zu erhalten. Die Rolle des terrestrischen Raums und das Zusammenspiel zwischen marinem und terrestrischen Raum werden untersucht.Die Arbeit basiert auf der Untersuchung von valanginzeitlichen Sedimenten zweier Lokalitäten, eine davon in einem randmarinen Gebiet in der nordwestlichen Tethys gelegen, die andere in einem Seeweg, der den borealen Raum mit der Tethys verband. Die Untersuchung der jeweiligen Archive erlaubt es sowohl Veränderungen im marinen Raum als auch solche auf den angrenzenden Kontinenten zu erforschen. Die Lokalität in der Tethys ist stratigraphisch sehr gut untersucht, eine Voraussetzung dafür, bereits bestehendes Datenmaterial zum Valangin hochauflösend mit den neu gewonnenen Daten zu vergleichen. Da für die zweite Lokalität keine gut abgesicherte Stratigraphie besteht, erfolgt die Korrelation mit dem Vokontischen Becken bio- und chemostratigraphisch. Für das Vokontische Becken (Südwest-Frankreich) wird eine hochauflösende Kohlenstoffisotopenkurve mit organischen und karbonatischen Kohlenstoffmessungen kombiniert. Die Kohlenstoffisotopenkurve bildet die Grundlage für eine Korrelation mit bestehenden Datensets zum Valangin. Zur Untersuchung des OM werden Biomarkeruntersuchungen durchgeführt, welche die erste derartige Untersuchung für die Tethys im Valangin darstellen. Variierende Häufigkeiten in Isoprenoiden, n-Alkanen, Steranen und Hopanen reflektieren das Auftreten von Organismen, welche die Quelle dieser Biomarker sind (z.B. Landpflanzen, Dinoflagellaten und Cyanobakterien). So können Abschätzungen zur Durchlüftung des Ablagerungsraums

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gemacht werden. Mit Ausnahme von zentimeterdünnen Lagen (den sogenannten Barrande Lagen), welche vor Beginn der CIE abgelagert wurden, können jedoch zumindest für dieses Becken keine anoxischen oder euxinischen Bedingungen nachgewiesen werden. Terrestrische Systeme werden anhand von Sporen-Pollen Kurven aus marinen Sedimenten rekonstruiert, welche die Paleo-Vegetation auf den angrenzenden Kontinenten reflektieren. Eine Kurve wurde für das Vokontische Becken erstellt und eine für den Mittel-Polnischen Trog (Karpathischer Seeweg, Zentral Polen). Die das Vokontische Becken umgebende Vegetation wurde von xerophytischen Cheirolepediaceen dominiert und die den Mittel-Polnischen Trog umgebende Vegetation von Araucariaceen/Cypressen. Beide Lokalitäten zeigen hohe Abundanzen von Pteridophyten. Für den Mittel-Polnischen Trog kann für den Zeitraum der beginnenden CIE an der Grenze zwischen unterem und oberem Valangin ein Trend zu einer graduell zunehmenden Feuchte festgestellt werden. Gleichzeitig ist die Vegetation um das Vokontische Becken von einer trockenen Phase geprägt. Im Anschluss an diese Phase von lokal unterschiedlichen Entwicklungen in der Vegetation zeigt sich die Etablierung überregional besonders feuchter Bedingungen. Daran schließt sich im Vokontischen Becken eine eher trockene Phase an.Basierend auf der Untersuchung der Palynofazies werden für beide Lokalitäten Schwankungen im Meeresspiegel und terrestrischer Eintrag rekonstruiert. N-Alkan-Datensets in Form von Biomarkerverhältnissen werden im Hinblick auf Veränderungen im Landpflanzeneintrag interpretiert um die Palynofazies-Daten zu unterstützen. Die Daten weisen auf einen erhöhten Meeresspiegel während der Grenze zwischen unterem und oberem Valangin hin, gefolgt von einer Phase mit niedrigem Meeresspiegel, welche bis ins unterste Hauterive anhält. Schließlich wird die Bedeutung von Schwankungen im atmosphärischen CO2-Gehalt während der CIE im Valangin anhand der Untersuchung von relativen pCO2 Veränderungen ermittelt. Hierzu dient der Vergleich von Veränderungen in karbonatischen und organischen Kohlenstoffisotopenkurven. Die Ergebnisse lassen einen zunehmenden Trend im atmosphärischen CO2-Gehalt während des Beginns der CIE erkennen, gefolgt von einem rapiden Absinken, das zeitgleich zu einem propagierten Abkühlungsereignis stattfindet. Diese Phase von abnehmendem CO2-Gehalt ist charakterisiert durch feuchte Bedingungen in den mittleren Breiten, ein Hinweis für die kontinentale Speicherung von Kohlenstoff. Damit ist diese kontinentale Speicherung von Kohlenstoff als ein wichtiger Auslöser für die CIE im Valangin identifiziert.Der Multi-Proxy-Ansatz der auf das Valangin Ereignis angewandt wurde erlaubt die Schlussfolgerung, dass die marine Speicherung von Kohlenstoff vermutlich von untergeordneter Bedeutung für die festgestellten Veränderungen im Kohlenstoffkreislauf war. Veränderungen auf den Kontinenten, insbesondere von Feuchteverhältnissen, lassen sich sehr gut mit der CIE korrelieren. Dies wird als Hinweis dafür gesehen, dass eine potentielle kontinentale Kohlenstoffspeicherung ein wichtiger auslösender Faktor für die CIE war. Diese Studie stellt einen wichtigen Schritt in Richtung eines detaillierteren Verständnisses der CIE im Valangin dar, welches über Veränderungen im marinen Raum hinausgeht.

Kurzfassung

XI

1. Introduction

1.1 Carbon and the climate systemIn 1896 Arrhenius calculated for the first time the impact of human emissions of CO2 on global warming. In 1930 a climatic warming trend was discovered but ascribed to natural cycles in weather patterns. It was not until 1938 when the hobby-meteorologist Guy Stewart Callendar gave a talk in front of the Royal Meteorological Society in London on weather statistics that the view of scientists on manmade global warming, by emitting carbon into the atmosphere, changed (Weart, 2008). The important role of carbon and its gaseous forms within the climate system is today widely accepted (e.g. Vaughan, 2007, and references therein). The modern climate setting with approximately 380 ppm atmospheric pCO2 (partial pressure CO2; Foster et al., 2007) and an average global mean temperature of ~13.9°C (NOAA, 2012) is rather exceptional in the Phanerozoic, and represents comparatively low atmospheric pCO2 concentrations and temperatures, accompanied by glaciated poles, only established for about the last 34 Ma (Zachos et al., 2008). If current greenhouse gas emissions would remain constant, by 2060 ~430 ppm pCO2 would be reached and temperatures would rise by about 1.3°C. Regarding rapid economic growth, e.g. in China, this scenario is, however, unlikely (Davis et al., 2010). An increase in pCO2 up to 600 ppm in 2100 and a warming by about 2.4° to 4.6°C is more likely (Meehl et al., 2007). The investigation of potential impacts an enhanced level of atmospheric pCO2 has on climate and the environment is therefore of high interest. Most of the time during the history of the earth atmospheric pCO2 has been much higher than today (Figure 1.1). An understanding of the operation mode of the climate system under high levels of pCO2 and the role of the carbon cycle within climate change is a basic requirement for understanding future global warming, needed for adaption human society has to face with. There do exist numerical models of the earth’s climate system to simulate the effects of greenhouse gas emissions, but they need to be validated (Williams et al., 2007). This can be done by investigating the geologic past via marine or terrestrial archives that provide paleo-records of climate-proxies. Proxy data allow assessing climate parameters like temperature, salinity, moisture levels, main wind direction, productivity, the amount of distinct elements and gases in water and atmosphere, etc., e.g. from sediments (e.g. Vaughan, 2007). Atmospheric pCO2 is controlled by the carbon cycle, which operates on different time scales. The short-term cycle, working on day to month scale, is the atmosphere-biosphere cycle mainly controlled by organic carbon via photosynthesis and respiration, and soil carbon take up due to plant decay. The intermediate cycle operates on years to millennia scale and is the atmosphere-ocean cycle, based on mainly inorganic (carbonate) carbon. It also includes marine and terrestrial methane clathtrades and swamps, mires, etc. The long-term or geologic carbon cycle works over millions of years and includes storage and release of carbon in and from rocks as well as by magmatic intrusions (e.g. Skelton et al., 2003).A better understanding of the relationship between pCO2, the carbon cycle, and climatic and environmental change requires a long term observation of these parameters, longer than the recorded history, which can be provided by the paleo-records. Much research on climate and the carbon cycle has been done. But considering the many uncertainties in steering factors and sub-cycles this research remains an enduring challenge. To

1

understand how and why the past was different from the present and what this may tell us about a future earth under greenhouse conditions with high pCO2 and temperatures, while today the earth is still in an icehouse mode (e.g. Takashima et al., 2006; Williams et al., 2007), one needs to “read” the paleo-records.

1.2 Cretaceous perturbations of the carbon cycleThe reconstruction of changes in atmospheric pCO2 and its connection to changes in climate and environment of ancient times provides a valuable fund for interpretations of recent and estimations of future climate change and its consequences for life on earth. An interesting analogue for a time of greenhouse conditions with high pCO2 and temperatures is represented by the era of the Mesozoic, especially the Cretaceous period (with average global mean temp. around 6°-9°C warmer than today, and pCO2 of ~600-1800 ppm, respectively; e.g. Sellwood and Valdes, 2007; Willis and McElwain, 2002). For this time interval a high number of paleo-records is available, providing more or less continuous climate-proxy records, especially since the beginning of the Deep Sea Ocean Drilling Project in the 1970s (e.g. Takashima et al., 2006). This makes this interval of earth history a perfect frame to study paleoenvironmental and climatic settings, interactions, and changes under high pCO2, which is of value considering the rising pCO2 levels predicted for the near future. Paleo-records revealed that the Cretaceous was a period that underwent several distinct environmental perturbations like the dawn of the angiosperms (flowering plants), the opening of the South Atlantic and further spreading of the North Atlantic, the formation of gigantic volcanic provinces, the rise and fall of enormous carbonate platform ecosystems, and finally the fall of the dinosaurs (e.g. Skelton et al., 2003; Sellwood and Valdes, 2007). Regarding carbon cycling one immediately ends up with focusing on time intervals of the Cretaceous known as Oceanic Anoxic Events (OAEs). These intervals represent phases of severe perturbations of the global carbon cycle evidenced in high amplitude positive and minor negative shifts in the carbon isotope records (δ13C, Jenkyns, 2010; Figure 1.2). They are defined as phases of enhanced, more than locally expressed storage of organic matter (OM) in marine pelagic and hemipelagic sediments. This sequesters 12C, since OM is 13C-depleted, in so called “black shales”, leaving the active carbon cycle 13C-enriched expressed by a positive shift to heavier δ13C values in carbon bearing archives (e.g. Schlanger and Jenkyns, 1976; Scholle and Arthur, 1980). These time intervals have been named OAEs since the occurrence of extended anoxic layers in oceans is regarded as the major explanation for the carbon cycle perturbations, in connection with either enhanced productivity or preservation or a combination of both,

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Chapter 1

leading to enhanced OM accumulation (Leckie et al., 2002 and references therein).Major Cretaceous OAEs are the Early Aptian OAE1a or “Selli Event” (e.g. Erba et al., 1999), the Aptian-Albian boundary OAE1b (e.g. Herrle et al., 2004), the Late Albian OAE1c and OAE1d, the End-Cenomanian OAE2, and the Coniacian-Santonian boundary OAE3 (e.g. Arthur et al., 1990). The question of triggers for expanded anoxic layers, enhanced OM storage, and black shale formation was answered in two ways by two major hypotheses: By the stagnant ocean model, with reduced mixing of the water column under an enhanced stratification and preservation, leading to bottom water anoxia (e.g. Schlanger and Jenkyns, 1976; Rullkötter, 2000), and by the expanded oxygen minimum model, under an enhanced primary production in the photic zone leading to water column anoxia, e.g. under upwelling conditions (e.g. Erbacher et al., 2001; Hasegawa, 2003; Takashima et al., 2006). The major initial trigger was supposed to be a rise in temperatures under rapid intrusion of CO2 into the active carbon cycle from volcanic or methanogenic sources (Jenkyns, 2010).

Introduction

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Fig. 1.2. Lower to Mid Cretaceous occurrence of black shales and perturbations in the δ13C record (modified from Westermann et al., 2010) compared to sea-level, ocean crust production, and platform drowning (modified from Takashima et al., 2006). Discrepancies in age intervals of the Valanginian compared to the text are based on the fact that different age models are used.

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Each OAE can, however, possibly be regarded as a to a certain degree unique event with individual driving factors (Erbacher et al., 1996). Recently, the imperative co-occurrence of black shales and OAEs has been challenged. Accordingly, OAEs must not lead to black shale formation in every case (Föllmi, 2012). Föllmi (2012) proposes the term EEC (episode of environmental change) for these events. This, however, raises the question of what would then define an OAE. Maybe the question should be the other way round if positive perturbations of the carbon cycle are in every case caused by an OAE. This approach would keep the connection of black shales and OAEs and would add the dimension of non-OAE carbon isotope excursions (CIEs). The Valanginian (~144.5-133.9 Ma; Ogg and Ogg, 2008) was the first time interval of the Early Cretaceous showing a severe global carbon cycle anomaly (e.g. Weissert and Erba, 2004; Föllmi et al., 2006). This is reflected in a positive excursion of the δ13C records, globally preserved in Corg and Ccarb containing sediments, of ~1.5 to 2.5‰ (marine carbonates, e.g. Lini et al., 1992; Hennig et al., 1999; Wortmann and Weissert, 2000), ~2 to 3‰ (marine organic matter, e.g. Lini et al., 1992, Wortmann and Weissert, 2000), and ~4 to 5‰ (terrestrial organic matter, e.g. Gröcke et al., 2005; Nunn et al., 2010). The Valanginian positive CIE was recently interpreted as being the result of an OAE and consequently the interval was named “Weissert OAE” (Lini et al., 1992; Erba et al., 2004) after the pioneering work of Helmi Weissert on this event (Weissert, 1989). However, the occurrence of widespread deposition of black shales or organic rich deposits comparable to younger Cretaceous OAEs cannot be confirmed to date (e.g. Weissert et al., 1998; Weissert and Erba, 2004; Westermann et al., 2010). “Standard” OAEs are described as starting with a rapid onset and being of short duration (< 1.0 Ma, Föllmi, 2012). They are supposed to be initiated by a CO2 intrusion into the ocean-atmosphere system by the formation of large igneous provinces (LIPs, Jenkyns, 2003), submarine volcanism (Leckie et al., 2002), or from methanogenic sources (Jenkyns, 2010). This would have caused global warming that accelerated the hydrologic cycle, thereby increased weathering and nutrient influx to oceans, intensified upwelling, and changed deep water formation, and finally led to OM accumulation and to a reorganization of the biosphere (e.g. Leckie et al., 2002, Takashima et al., 2006; Jenkyns 2010). The occurrence of short episodes of climatic cooling, punctuating global warmth, probably connected to pCO2 drawdown after OM storage, is as well thought to be a general phenomenon associated with OAEs (Weissert and Erba, 2004; McElwain et al., 2005; Takashima et al., 2006; Forster et al., 2007). They were characterized by either rising or falling sea-level (e.g. Leckie et al., 2002). Especially OAE1a and OAE2 are associated with declines in calcareous nannofossils. In general, recurrent carbonate platform drowning occurred during the Mesozoic, probably related to an expansion of anoxia within the euphotic zone and eutrophication (Takashima et al., 2006 and references therein). These intervals have been termed “biocalcification crisis” (Weissert and Erba, 2004). Enhanced ocean crust production, inducing biolimiting metals that probably changed ocean chemistry, was also regarded as an important factor (e.g. Takashima et al., 2006). Black shales often show high abundances of biomarkers of e.g. cyanobacteria (e.g. Kuypers et al., 2004) or green sulphur bacteria origin pointing to euxinic conditions (e.g. Damsté and Köster, 1998), and supporting the idea of expanded anoxia (Takashima et al., 2006). Biomarkers can be used for various approaches. They are defined as organic compounds composed of carbon, hydrogen, and other elements occurring in sediments, rocks, and crude oils with little to no change in structure compared to their precursor molecules in living organisms (Bianchi and Canuel, 2011 and references therein). The Valanginian differed from younger Cretaceous OAEs in terms of duration (longer, ~3 Ma e.g. Gréselle et

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Chapter 1

al., 2011), ocean crust production (lower, see Fig. 1.2), and OM storage (less). The role of volcanic activity is still unclear (e.g. Barbarin et al., 2012). It was, however, as well characterized as being connected to short term cooling (e.g. Weissert and Lini, 1991; Price 1999) and a carbonate production crisis, e.g. reflected in platform drowning (Föllmi et al., 2006; Fig. 1.2). Rare Valanginian marine organic rich deposits, e.g. known from ODP site 1213 in the Western Atlantic (probably connected to enhanced upwelling) and Europe, were ascribed to enhanced terrestrial input under an accelerated hydrologic cycle and rising sea-level (Brassell 2009 and references therein). In the past few years the occurrence of a Valanginian OAE has consequently been questioned, major forcing mechanisms remain an unknown (Westermann et al., 2010; Gréselle et al., 2011; Föllmi, 2012). This raises the question of alternative explanations or more effective or additional factors than just oceanic anoxia and marine carbon sequestration for the recorded perturbation in the carbon cycle. Alternative explanations have been proposed by van de Schootbrugge et al. (2000), Price and Mutterlose (2004), and Westermann et al. (2010) moving the terrestrial realm as a sink for 13C-depleted OM into the focus. Terrestrial records, e.g. in the form of lake sediments, paleosoils, or eolian dunes, that may prove these ideas are, however, rare. In general, terrestrial records are of low resolution, continuity, or preservation, and provide dating problems. Marginal marine deposits rich in terrestrial input can bridge this deficiency. The investigation of different intervals of carbon cycle perturbations probably provides different insights into the carbon-climate system, since they all had slightly or severely different causes. Investigating the Valanginian carbon cycle anomaly may provide another dimension of climate change whereby the role of the terrestrial realm within the climate system may become better understood.

1.3 Vegetation and the climate systemThe role of terrestrial plants within the global carbon cycle and deep time climate change has long been underestimated and so has the terrestrial realm in general received less attention while the oceans were in focus of attention (e.g. Gröcke et al., 2005; McElwain and Punyasena, 2007). Beerling (2008) wrote “But how many of us have stopped to wonder how remarkable plants are, how profoundly they have altered the history of life on Earth, and how critically they are involved in shaping its climate?” and thereby highlighted the use that can be gained from studying plants to get to know more about climate change. Carbon is the substrate of primary production, which represents the direct link between plants and the atmospheric carbon reservoir. The process of photosynthesis discriminates against 13C, which makes OM depleted in 13C (Farquhar et al., 1989; Sigman and Haug, 2003). The storage of terrestrial OM in swamps and mires, occasionally followed by coal formation and the oxidation of this matter thereby must have profound influences on the intermediate as well as long-term carbon cycle (e.g. Wissler, 2001; Beerling and Royer, 2002; Kurtz et al., 2004). Besides the fact that plants are an active part of the carbon cycle vegetation changes reflect changes in climate settings like the expansion of arid belts or the distribution of precipitation (e.g. Vakhrameev, 2010). Perturbations in the carbon cycle, if associated with Corg burial on land, should also be accompanied by environmental change (e.g. in moisture) and thereby should be reflected in changes of the vegetation structure. Furthermore, it was for example shown that there is a high correlation between pCO2 and the composition of floras, with angiosperms being of higher abundance under low levels of pCO2, and pteridophytes (fern plants) and gymnosperms (conifers, cycads, ginkgos) being comparatively more abundant under high pCO2 (Fig. 1.3; Willis and McElwain, 2002). This reveals that the investigation of both, the carbon-cycle and vegetation change provides interesting insights into the history of climate and environmental change.

5

Introduction

Plant fossils, like the impressive Cretaceous tree trunks of the arctic region (e.g. Skelton et al., 2003), could help reconstructing the Cretaceous vegetation and interactions of vegetation and climate, and therewith could provide information about the terrestrial realm and its importance for the ancient carbon cycle. But they are rare and in general do not provide continuous records over longer time spans but only snap-shots. This gap can be closed by palynology, which can be used to reconstruct vegetation history and deduced changes in climate. Palynology is the study of organic microfossils, exclusive of those destroyed by hydrochloric acid (HCL), i.e. those consisting of calcium carbonate (CaCO3), and those destroyed by hydrofluoric acid (HF), i.e. consisting of silicia (mainly radiolarians and diatoms; Traverse, 2007). The study of palynology can be separated into the study of organic walled cysts of marine origin, palynofacies, and spores and pollen. These are determined by microscopy, whereby qualitative and quantitative approaches are used. The study of organic walled marine cysts helps reconstructing marine environmental conditions or can be used for biostratigraphic purpose (especially dinoflagellate cysts, Traverse, 2007). Palynofacies changes can be used to reconstruct changes in the proportion of terrestrial to marine input and in their respective compositions (e.g. Tyson, 1995; Feist-Burkhardt and Götz, 2002; Götz et al., 2009). This reveals information on the sedimentary environment and the producing biosphere (Combaz, 1964). The study of spores and pollen provides information about changes in the composition of the ancient vegetation, by assigning them to parent plants. Spores and pollen for this matter (excluding e.g. spores of bacteria and fungi) are the medium of reproduction of non-seed producing land-plants, and of seed plants, respectively (Traverse, 2007; Fig. 1.4).Especially spore-pollen records derived from marginal marine archives rich in terrestrial input are of use for the study of deep-time environmental change, since marine archives have the advantage of being potentially longer ranging, of higher resolution, and are more easily datable since they e.g. bare microfossils of

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Chapter 1

Fig. 1.3. Relationsphip between CO2 and abundances (in %) of spore-pollen in fossil floras of angiosperms (left), gymnosperms (middle), and pteridophytes (right), with atmosperic CO2 during the Cretaceous (modified from Willis and McElwain, 2002).

biostratigraphic use or allow for precise chemostratigraphic correlation. These records reflect the composition of the vegetation of adjacent continents and changes therein and thereby provide information about climate and terrestrial environments. One of the first to describe Jurassic and Cretaceous vegetation based on macrofossils and palynology was the Russian Vsevolod A. Vakhrameev, who published a synthetic book of his

Spore Pollen Marine cyst

work on this matter in 1985 (Vakhrameev, 2010). A distinct feature he described for the Early Cretaceous is the occurrence of the now extinct group of the drought-resistant conifers Cheirolepediaceae, represented by the pollen Classopollis (Fig. 1.5), which declined towards the end of the Early

Cretaceous. Based on phytochoria (geographic areas of relatively uniform composition of the vegetation) he defined climatic belts for the Early Cretaceous (Fig. 1.6). Remarkable is the widespread area of semiarid to arid conditions, also spanning the equator, characterized by high occurrences of Classopollis (>50%). Vakhrameev describes a cooling and humidification, accompanying the decline in Cheirolepediaceae, for the northern hemisphere. One other profound feature regarding Cretaceous vegetation is the dawn of angiosperms, probably first occurring close to the equator and subsequently spreading polewards (Brenner, 1976; Crane and Lidgard, 1989). First forms of pollen of supposedly angiosperm origin are already recorded

from probably Valanginian to Hauterivian strata of Israel (Brenner, 1996), under a declining long-term trend in pCO2 (Willis and McElwain, 2002; Fig. 1.3). Even though they rapidly became the most important part of the continental flora with highest numbers of species, during the earliest Cretaceous they were still unimportant

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7

Introduction

Fig. 1.4. Microscopic images of a spore, pollen, and marine cyst (own photographs).

Classopollis

Fig. 1.5. Microscopic image of the pollen Classopollis (own photograph).

Fig. 1.6. Climatic belts of the Early Cretaceous, continents in grey (modified after Vakhrameev, 2010).

(e.g. Niklas et al., 1983). But of course it is of interest to reconstruct the environment that they were about to enter to better understand their history, too. A more important feature characterizing the earliest Cretaceous, at least that of the northern hemisphere, was the rapid diversification of pteridophytes in the mid-latitudes (Batten, 1984; Herngreen et al, 1996), maybe due to high sea-level and resulting provincialism (Diéguez et al., 2010). Plants may control the carbon-cycle to a certain degree and thereby influence the climate and are on the other hand strongly affected by climatic conditions. If plants shape the climate or are shaped by it, in either way they are a valuable tool for its reconstruction. Especially palynologic approaches are thereby of tremendous use.

1.4 The climate system of the Early CretaceousThe climate system of the Early Cretaceous was determined by some distinct factors, different from today. PCO2 levels during the Early Cretaceous varied between 560 and 1200 ppm (Haworth et al., 2005). A key driving factor for long term trends in pCO2 being the supercontinent cycle with crust production and volcanism (+CO2), peaking in the Cretaceous, and weathering of mountain belts (-CO2; Vaughan, 2007 and references therein). Mountain belts have in general been low during the Early Cretaceous, since continents were drifting apart from each other, not being an important contributor for long-term pCO2 drawdown (e.g. Skelton et al., 2003). The lack of major mountain belts of course has a direct effect on climate, as well, since it influences surface heating and winds. Plate tectonic constellations, a determinant factor for climatic settings, during the Early Cretaceous probably formed three large blocks with the southern continents (Gondwana and the Antarctic) being separated from the northern ones (North America/Laurasia; Fig. 1.7). As a consequence, a circumglobal oceanic connection was established north to the equator, with a wide eastern Tethys Ocean (e.g. Hay et al., 1999; Skelton et al., 2003; Sellwood and Valdes, 2007). The Cretaceous was predominantly ice-free (Takashima et al., 2006), but like mentioned in context with OAEs, was punctuated by cool intervals (Price, 1999). The poles were probably not covered by permanent ice caps but received winter snow and maybe at least seasonal sea ice was formed (Hay, 2002; Selwood and Valdes, 2007). Consequently, and probably additionally due to thermal expansion of water and high sea-floor spreading rates, sea-level during the Early Cretaceous was significantly higher than today (max. 40 %, Takashima et al., 2006; Selwood and Valdes, 2007). This led to archipelagic continental margins (e.g. southern Europe and southern Eurasia, Diéguez et al., 2010 and references therein) and the occurrence of epicontinental seas (Mutterlose and Kessels, 2000). A low equator to pole temperature gradient was established (e.g. Skelton et al., 2003; Hay, 2008). Seasonality is known to have been low (Diéguez et al., 2010). The zone of highest rainfall with precipitation exceeding evaporation was moving with the Inner Tropical Convergence Zone (ITCZ, Selwood and Valdes, 2007). An in general important factor controlling climate change, and thereby also of importance for Cretaceous climate, is the change in insolation driven by the earth’s orbital parameters obliquity, precession, and eccentricity, known as Milankovitch cycles. Major cycles therein occur on 19, 41, and 100 kyr and larger cycles on 400 kyr, 1.24, 2.35, and 4.6 Ma frequency, additionally affected by planetary constellations, important for long-term climate change (e.g. Vaughan, 2007 and references therein). On short-term scale the changing intensity in solar output is as well an important factor (e.g. Vaughan, 2007). A crucial factor concerning the interplay of vegetation, or terrestrial environments in general, and climate is moisture. Given the presence of a large ocean extended in east-west direction close to the equator and a parallel shoreline of southern North America and southern Eurasia during the Early Cretaceous, the

8

Chapter 1

establishment of monsoonal climates controlling moisture for a significant portion of the earth surface can be assumed. A monsoonal circulation is not, like thought earlier, a simple land-sea breeze on a larger scale (Wang, 2009). Its moisture transport is indeed accelerated by a land-ocean heat contrast, leading during summer to the flow of moisture from highs over oceans to lows over continents. The determinant factor for its latitudinal expression is, however, the annual track of the ITCZ and its poleward migration, and the expansion of the associated atmospheric circulation system of the Hadley cells (Wang, 2009; Roedel and Wagner, 2011). This determines the location of the trade winds, forming the monsoonal winds. The trade winds transport air masses towards the ITCZ, where they ascend and thereby climatologically separate both hemispheres (Bridgman and Oliver, 2006). Depending on the latitudinal position of the ITCZ, these winds transport moist air from the ocean towards the coastline or dry air from the continents towards the ocean (Fig. 1.7, e.g. Roedel and Wagner, 2011). During the Early Cretaceous, under the absence of expanded polar ice, no polar highs were established which in the modern world push the ITCZ towards the equator (Ziegler et al., 1987, Ueda et al., 2010). This probably led to a higher amplitude of poleward ITCZ-migration and less concentrated rainfall across the equator, with less established wet zones (c.f. Fig. 1.6) and areas further poleward being affected by this circulation system. A less restricted moisture distribution probably caused highly seasonal precipitation patterns. The fact that the southern Eurasian coastline has been located by ~10° closer to the equator than today (e.g. Ziegler et al., 1987), in combination with the higher amplitude of ITCZ migration, highlights the possibility that Paleo-Europe was, different from today, during the Cretaceous affected by this ITCZ-circulation. It has long been discussed what may steer the position of the ITCZ and its annual track. Now it is known that it is determined by the zone of ascending air-masses over the area of maximum surface temperature, annually circulating approximately under the 90° inclination of the sun, influenced by continent-ocean constellations (e.g. Bridgman and Oliver, 2006). Thereby, the expression of monsoonal climates on geological time scales is to a certain degree controlled by changes in insolation, controlled by the Milankovitch cycles. Today, local expressions of the monsoonal precipitation intensity and distribution are strongly affected by major mountain belts like the Himalaya with the Tibetan Plateau (Hay, 1996). But monsoonal circulations are also known from West Africa, southwestern North America, southern Mexico, and northern Australia (Bridgman and Oliver, 2006). Since major mountain belts are missing during the Early Cretaceous, the global monsoon pattern was probably more homogenously expressed.

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9

Introduction

Fig. 1.7. Cretaceous northern hemisphere (NH) summer monsoon affecting the southern margin of Laurasia and probably North America (here the ITCZ-track is speculative depending on the max. heating of the adjacent continents of Gondwana and North America), with the northernmost extension of the ITCZ over the heated continents (monsoonal reconstruction based on Wang (2009) and Roedel and Wagner (2011); plate tectonic reconstruction based on Ziegler et al. (1987), Smith et al. (1994), Hay (2002), Mutterlose et al. (2003), Rees McAllister et al. (2004), Blakey (2010), Vakhrameev (2010)), L=low pressure, H=high pressure.

A scenario for the Cretaceous atmospheric circulation proposed by Hay (2002) with a two-cell circulation system per hemisphere and Hadley cells descending as far polewards as ~70° is unlikely. Based on the distribution of desert belts Hasegawa et al. (2011) reconstructed the Cretaceous location of the subtropical high-pressure belt, located under the descending branches of the Hadley cells, for the Asian interior, not migrating further polewards than 45°N. Furthermore, such largely expanded Hadley cells would probably be instable at least over continents due to the distracting coriolis force. Additionally, the location of the descending branches of the Hadley cells is relatively persistent in relation to the ITCZ, not enlarging in addiction to adjacent poleward wind systems (Bridgman and Oliver, 2006). The determinant factor for a poleward expansion of Hadley cells was assumed to be the amount of atmospheric pCO2 and temperature (Hasegawa et al., 2011). Further research on this atmospheric circulation system is needed. The monsoonal system during the Cretaceous and its local expression and influence on environments is to date not yet fully understood. It was, however, probably important for terrestrial biomass production, vegetation structure, the occurrence of swamps, etc. and therewith regarding terrestrial carbon storage.

1.5 Aims and objectives of this studyThe present study is a multi-proxy investigation of environmental and climatic change accompanying the Valanginian CIE. The established records are on a high resolution and allow for a detailed investigation of the lead and lag order of environmental change accompanying this event. This allows for a differentiation of causes and symptoms. Potential initial triggers of this carbon cycle anomaly can thereby be investigated. Previous studies have already highlighted the potentially important role of the terrestrial realm during the Valanginian event, which differs in various aspects from younger Cretaceous events of carbon cycle perturbations (van de Schootbruge et al., 2000; Price and Mutterlose, 2004; Westermann et al., 2010). High-resolution records on terrestrial change are to date missing. This study aims at helping to close this gap. In order to do this two sites of the mid-latitudes have been investigated. One located in southeast France, in the Vocontian Basin, the other one in central Poland, in the Mid-Polish Trough. The French succession represents a composite section based on three outcrop sections (La Charce, Vergol, Morenas), the Polish section was sampled in form of a drill core, taken close to the village of Wąwał. The latter was located in the centre of the Carpathian seaway, connecting the northwestern Tethys and the Boreal realm. The French section was located in the centre of a marginal marine basin at the northwestern edge of the Tethys Ocean. Both sites provide marine sediments admixed with a certain amount of terrestrial input, allowing to study the marine depositional environment as well as terrestrial environmental changes on adjacent continents. This thesis will provide valuable information on the functioning modes of the Cretaceous ocean-atmosphere-biosphere system in close correlation to the carbon cycle during the Valanginian CIE. Key objectives of this study are the establishment of records on terrestrial environmental change accompanying the carbon cycle anomaly, based on palynologic approaches. Spore-pollen records as well as records on changes in the palynofacies are established for both presented sites. They are associated with climate dynamics. The already questioned establishment of a globally expressed OAE is tested for the Vocontian Basin by geochemical approaches. Chemostratigraphy as well as biostratigraphy are used for stratigraphic purpose and are partly newly established. A record for changes in pCO2 is established and correlated with environmental and climatic change. Geochemical approaches in form of biomarker investigations and compound specific carbon isotope measurements are used to support the data sets.

10

Chapter 1

Fundamental questions are:

i) Can the establishment of anoxia during the Valanginian CIE be affirmed or excluded?

ii) Did mid-latitudinal vegetation and terrestrial environments in general change during this

carbon cycle perturbation?

iii) If so, what does this imply regarding causes and consequences of the Valanginian carbon cycle perturbation?

iv) Is this CIE accompanied by changes in pCO2?

1.6 Outline of this manuscriptChapter 1 introduces into the theory of Cretaceous carbon cycle perturbations and demonstrates the importance of studying them regarding an understanding of the carbon-climate system. The use and possibilities of investigating vegetation changes in this respect are explained and the climate scenario of the Early Cretaceous, where the Valanginian, which is in focus of this thesis is placed in, is presented. In Chapter 2 the methods used in this study are briefly presented (see respective chapters for details). Chapter 3 tests the establishment of anoxia in the Vocontian Basin of southeast France during the Valanginian CIE by the first Valanginian biomarker study established for the Tethys, in connection with carbonate carbon chemostratigraphy. In combination with additional geochemical analysis the composition of the OM is investigated. Chapter 4 deals with vegetation and deduced climate dynamics based on palynologic approaches. Records of changing compositions in plant assemblages for the French and the Polish site are established. Based on botanical affinities of the identified taxa, habitats are reconstructed and moisture changes are deduced. The possibility of varying compositions of the vegetation during the Valanginian CIE is investigated. Changes in the vegatation are compared to potential changes in climate settings. A correlation and comparison between the Tethyan and Boreal realms is accomplished. In Chapter 5 sea-level and environmental change is reconstructed based on palynofacies and n-alkanes and compared between the two studied sites. The use of palynofacies changes for sea-level reconstructions for the Valanginian is tested based on comparisons with established records on sea-level change (Gréselle and Pittet, 2010; Mutterlose and Bornemann, 2000). Chapter 6 provides a high resolution record for the Vocontian Basin on changing trends in pCO2 based on the paleo- pCO2 proxy in the form of a ∆δ record, calculated from bulk material measurements on δ13Corg measurements of predominantly phytoplankton origin and δ13Ccarb. It is flanked by compound specific carbon isotope measurements on isoprenoids and a potentially land-plant derived n-alkane of the French section. The established record presents the first reliable high-resolution proxy record for Valanginian changes in pCO2. The role of changes in pCO2 within the comlex suite of environmental perturbations during the Valanginian CIE is estimated. In Chapter 7 a synthesis of the present study is assembled and the main questions will be answered. Additionally, future perspectives are given.

The project was supervised by Prof. U. Heimhofer (University of Hannover, Germany) and Prof. J. Mutterlose (Ruhr-University Bochum, Germany). Collaborational work has been done with Dr. C. Ostertag-Henning (BGR, Hannover, Germany) and Prof. Stefan Schouten (NIOZ, Texel, The Netherlands) for biomarker analysis. For palynology, a collaboration was accomplished with Prof. P. A. Hochuli (ETH Zürich, Switzerland). The field work in France was accompanied by Dr. B. Gréselle (Neftex Petroleum

11

Introduction

Consultants Ldt., UK). In Poland a core was sampled, drilled by the Geologic Institute of Warsaw, Poland. This was accompanied by Dr. I. Ploch (Geologic Institute of Warsaw, Poland), Dr. T. Adatte, and C. Morales and enabled by Prof. K. Föllmi (all University of Lausanne, Switzerland).

The present thesis and its chapters are discrete manuscripts, prepared for submission to international peer-reviewed journals or already published (Chapter 3). Repetitions are therefore unavoidable.

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Crane, P.R., Lidgard, S., 1989. Angiosperm diversification and paleolatitudinal gradients in Cretaceous floristic diversity. Science 246, 657-678.

Damsté, J.S.S., Köster, J., 1998. An euxinic southern North Atlantic Ocean during the Cenomanian/Turonian oceanic anoxic event. Earth and Planetary Science Letters 158, 165-173.

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Diéguez, C., Peyrot, D., Barrón, E., 2010. Floristic and vegetational changes in the Iberian Peninsula during Jurassic and Cretaceous. Review of Palaeobotany and Palynology 162, 325-340.

Erba, E., Channel, J.E.T., Claps, M., Jones, C.E., Larson, R.L., Opdyke, B., Premoli Silva, I., Rica, A., Salvini, G., Torrriceli, S., 1999. Integrated stratigraphy of the Cismon Apticore (southern Alps, Italy); a

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Introduction

“reference section” for the Barremian-Aptian interval at low latitudes. Journal of Foraminiferal Research 29, 371-391.

Erba, E. Bartolini, A., Larson, R.L., 2004. Valanginian Weissert oceanic anoxic event. Geology 32 (2), 149-152.

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Erbacher, J., Huber, B.T., Norris, R.D., Markey, M., 2001. Increased thermohaline stratification as a possible cause for an oceanic anoxic event in the Cretaceous period. Nature 409, 325-327.

Farquhar, G.D., Ehlringer, J.R., Hubick, K.T., 1989. Carbon Isotope Discrimination and Photosynthesis. Annual Review of Plant Physiology and Plant Molecular Biology 40, 503-537.

Feist-Burkhardt, S., Götz, A.E., 2002. Palynofazies und Sequenzstratigraphie (K1). Schriftenreihe der Deutschen Geologischen Gesellschaft 18, 57-72.

Föllmi, K., 2012. Early Cretaceous life, climate and anoxia. Cretaceous Research 35, 230-257.

Föllmi, K., Godet, A., Bodin, S., Linder, P., 2006. Interactions between environmental change and shallow water carbonate buildup along the northern Tethyan margin and their impact on the Early Cretaceous carbon isotope record. Paleoceanography 21, 1-6.

Forster, A., Schouten, S., Moriya, K., Wilson, P.A., Sinninghe Damsté, J.S., 2007. Tropical warming and intermittent cooling during the Cenomanian/Turonian oceanic anoxic event 2: sea surface temperature records from the equatorial Atlantic. Paleoceanography 22, doi:10.1029/2006PA001349.

Foster, P., Ramaswamy, V., Artaxo, P., Berntsen, T., Betts, R., Fahey, D.W., Haywood, J., Lean, J., Lowe, D.C., Myhre, G., Nganga, J., Prinn, R., Raga, G., Schulz, M., Van Dorland, R., 2007. Changes in Atmospheric Constituents and in Radiative Forcing. In: (Eds.) Solomon, S., Quin, D., Manning, M., Chen, Z., Marquis, M., Averyt, K.B., Tignor, M., Miller, H.L., Climate Change 2007: The Physical Science Basis. Contribution of Working Group I to the Fourth Assessment Report of the Intergovernmental Panel on Climate Change. Cambridge University Press, Cambridge, United Kingdom and New York, USA, 129-234.

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Gréselle, B., Pittet, B., Mattioli, E. Joachimski, M., Barbarin, N., Riquier, L., Reboulet, S., Pucéat, E., 2011. The Valanginian isotope event: A complex suite of palaeoenvironmental perturbations. Palaeogeography, Palaeoclimatology, Palaeoecology 306, 41-57.

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Hasegawa, H., Tada, R., Jiang, X., Suganuma, Y., Imsamut, S., Charusiri, P., Ichinnorov, N., Khand, Y., 2011. Drastic shrinking of the Hadley circulation during the mid-Cretaceous supergreenhouse. Climate of the Past Discussions 7, 119-151.

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Hay, W.W., DeConto, R., Wilson, K.M., Voigt, S., Schulz, M., Rossby Wold, A., Dullo, W.C., Ronov, A.B., Balukhovsky, A.N., Söding, E., 1999. Alternative global Cretaceous paleogeography. In: (Eds.) Barrera, E., Johnson, C.C., Evolution of the Cretaceous Ocean-Climate system. Geological Society of America Special Paper 332, 1-47.

Hennig, S., Weissert, H., Bulot, L., 1999. C-isotope stratigraphy, a calibration tool between Ammonite- and Magnetostratigraphy: The Valanginian-Hauterivian Transition. Geologica Carpathica 50 (1), 91-96.

Herngreen, G.F.W., Kedves, M., Rovina, L.V., Smirnova, S.B., 1996. Chapter 29, Vegetational Histora, 29C. Cretaceous Palynofloral Provinces: A review. In: (Eds.) Jansonius, J., McGregor, D.C.m Palynology; principles and applications. American Association of Stratigraphic Palynologists Foundation 3, 1157-1188.

Herrle, J.O., Kössler, P., Friedrich, O., Erlenkeuser, H., Hemleben, C., 2004. High-resolution carbon isotope records of the Aptian to Lower Albian from SE France and the Mazagan Plateau (DSDP Site 545): a stratigraphic tool for paleoceanographic and paleobiologic reconstruction. Earth and Planetary Science Letters 218, 148-161.

Jenkyns, H.C., 2003.Evidence for rapid climate change in the Mesozoic-Palaeogene greenhouse world. Philosophical Transactions of the Royal Society London A 361, 1885-1916.

Jenkyns, H.C., 2010. Geochemistry of Oceanic Anoxic Events. Geochemistry, Geophysics. Geosystems 11:Q03004.

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Kuypers, M.M.M., Schouten, S., Erba, E., Sinninghe Damsté, J.S., 2004. N2-fixing cyanobacteria supplied nutrient N for Cretaceous oceanic anoxic events. Geology 32, 853-856.

Leckie, R.M., Bralower, T.J., Cashman, R., 2002. Oceanic anoxic events and plankton evolution: Biotic response to tectonic forcing during the mid-Cretaceous. Paleoceanography 17 (3), 13-29.

Lini, A., Weissert, H., Erba, E., 1992. The Valanginian carbon isotope event: a first episode of greenhouse climate conditions during the Cretaceous. Terra Nova 4, 374-384.

McElwain, J.C., Wade-Murphy, J., Hesselbo, S., 2005. Changes in carbon dioxide during an oceanic anoxic

15

Introduction

event linked to intrusions into Gondwana coals. Nature 435, 479-482.

McElwain, J.C., Punyasena, S.W., 2007. Mass extinction events and the plant fossil record. Trends in Ecology and Evolution 22 (10), 548-557.

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17

Introduction

stratigraphy and platform drowning events: a possible link? Palaeogeography, Palaeoclimatology, Palaeoecology 137, 189-203.

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19

Introduction

20

21

2. Methods

2.1 Field WorkIn 2009 field work was done in southeast France to sample three sites based on an existing log (of Dr. B. Gréselle) which indicated the sites as covering Valanginian strata. The sites form a composite succession of ~175 m. They include three outcrop sections, La Charce, Vergol, and Morenas. Marly intervals have been sampled from marl-limestone alternations by using a rock hammer. To obtain fresh material the uppermost 15-20 cm have been removed before sampling. Furthermore, also in 2009, a drill core of ~18 m was sampled close to the village of Wąwał, central Poland, which was taken by the Geological Institute of Warsaw, Poland. Due to an adjacent clay pit that is no longer accessible but was well dated by ammonite biostratigraphy (Kutek et al., 1989) the core was assumed to cover Valanginian strata and thereby was of interest for this study. The sampled core material consists of clay to claystone material and was taken with a spattle. Geochemical bags were used for storage of the sample material to avoid contamination with hydrocarbon compounds of conventional plastic bags, problematic for biomarker analyses.

2.2 Carbon and oxygen isotope analysisFor the French site samples have been measured for stable carbonate carbon and oxygen isotope values (δ13C, δ18O). Marls have been powderized with an achat mortar or drilled with a micro drill to gain material for the measurements. The latter was done in case a sample was only prepared for isotope measurements and no further analyses were performed, since in this case only a small amount of powdered material was needed. This drilling was done in matrix material to ensure representativeness for the sample.Measurements were accomplished on ~0.6 mg of the material using a Gasbench II carbonate device (Thermo Fisher Scientific MAT) connected to an isotope ratio mass spectrometer (IR-MS; Thermo Fisher Scientific Delta S) at the Ruhr-University, Bochum, Germany, calibrated with international reference material. Calcite samples were reacted with phosphoric acid at 70°C for one hour before measurement. Relative abundances of isotopes (12C/13C, 16O/18O) are measured and given in parts per thousand (‰) based on correlation to international standards, multiplied by 1000.

2.3 Organic carbon isotope analysisThe isotopic composition of the organic carbon (δ13Corg) was measured for the French samples on an amount of ~0.6 mg powdered and decarbonatizized sample material. Therefore an elemental analyzer (EA; Thermo scientific Trace GC Ultra) coupled to an isotope ratio mass spectrometer (IR-MS; Finnigan Mat delta S via a Thermo scientific Finnigan GC Combustion III) available at the Ruhr-University Bochum, Germany was used, where the sample was flash combusted in a temporarily oxygen-enriched atmosphere, before MS-measurement of the oxidation products. Calibration is based on international reference material.

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Chapter 2

2.4 Carbonate carbon and total organic carbon contentTotal inorganic carbon (TIC) content was determined for the French samples by decarbonatization at the Ruhr-University, Bochum, Germany. Therefore, 2 mg of the powdered material was treated with 6 M HCL and neutralized with deionized water, repeated until a pH of 5 to 6 was reached. The weight difference before and after acid treatment and drying of the residue was assumed as the CaCO3 content in percent.To determine total organic carbon (TOC) content of the French samples, a LECO CS 200 was used, available at the BGR Hannover, Germany (only for some additional samples not measured by RockEval pyrolysis (see 2.5), where a determination of the TOC % is included). Therefore, 2 mg of the powdered sample material is treated with HCL and then combusted under low oxygen atmosphere. Carbon is burned to CO2 and thereby quantified by an infrared detector.

2.5 RockEval pyrolysisRockEval pyrolysis was used to determine type and thermal maturity of sediment samples of the French site. It is a rapid screening technique for characterizing organic matter with respect to biological sources, post-depositional degradation, and diagenetic overprint, which is based on thermal decomposition of organic matter under heat and in the absence of oxygen (Espitalié et al., 1977). An amount of ~100 mg of dried and powdered sample material was measured on a Rock-EvalTM6 at the BGR Hannover, Germany. Thereby TOC % is determined as well as a hydrogen index (HI), calculated from the amount of free hydrocarbons, and an oxygen index (OI), calculated from the amount of CO2 generated from the organic matter. The temperature of maximum hydrocarbon generation during determination of the HI is given as Tmax. Based on standard statistical evaluation of results in a van Krevelen diagram, information on type and maturity of the organic matter is gained (e.g. Lüniger and Schwark, 2002). If migrated bitumen is present can be estimated based on the production index (PI), which is calculated by a ratio of hydrocarbons volatilized versus those cracked from kerogen during heating.

2.6 Biomarker analysisFor analysis of biomarkers for both sites the organic solvent extractable bitumen was extracted from the sediment sample. This was done by an accelerated solvent extraction (ASE) technique at the BGR, Hannover Germany on ~15 to 28 g of the powdered sample using CH2CL2 (DCM=Dichloromethane). By adding activated copper to the extracted solvent, elemental sulfur was removed. Asphaltenes were separated from the extract by precipitation in petroleum ether. The aliphatic and aromatic hydrocarbons as well as the heterocompounds were received by fractionation of the remaining bitumen using a medium-pressure liquid chromatography (MPLC) with solvents of increasing polarity. Only the aliphatic fraction (apolar) was then analyzed by gas chromatography connected to a flame ionization detector (on an Agilent 6890 GC-FID). Compounds were quantified based on n-alkane standards. By this, identification and quantification of n-alkanes and the isoprenoids pristane and phytane was accomplished. For identification and quantification of further biomarkers a gas chromatography-mass spectrometry (GC-MS) was used (on an Agilent 6890-Finnigan MAT 95S), with full scan and metastable reaction monitoring modes (MRM). By this the transition of specific parent ions to daughter fragments allows to identify specific biomarkers like steranes and hopanes. For compound specific measurements of δ13C of pristane, phytane, and n-alkanes a C-isotope-ratio-monitoring gas chromatograph, connected to a mass spectrometer (GC-IRMS), at the NIOZ, Texel, the Netherlands was used.

23

Methods

2.7 PalynologyFor both sites analyses of palynology and palynofacies have been performed. Therefore, strew slides have been prepared at the “Geologischer Dienst Nordrhein-Westfalen”, Krefeld, Germany. An amount of ~30 g of the sample material has been mortared to a grain size of ~2mm. Carbonate has been removed by dissolution in heated HCl (hydrogenchloride). Silicates have been removed by dissolution in heated 70 % HF (hydrogenfluoride) in a sand-bath, between min. 2 to max. 25 hours. Here, time span was not relevant. Flouridic gels are formed by this process and removed by a second dissolution of the sample material in heated HCL. Samples are sieved under alternating vacuum with 10 μm mesh size. By this process larger particles of amorphous organic matter (AOM) are removed. This is considered as non-relevant for interpretation due to the fact that the occurrence of AOM will still be indicated by smaller particles of this fraction, counted for palynofacies analysis. Sample material is stored in glass vials. For conservation PVA (polyvenylalcohol) and distilled H2O (3 g to 100 ml) are used, dissolved under 200°C using a magnetic stirrer. A drop of formaldehyde prevents infection with fungi. After 12 hours a small amount is transferred to a glass slide by using a dropping bottle made of nalgene. These strew slides are prepared for permanence using synthetic resin (Elvacite 2044, of Tennants) dissolved in Xylol (18 g to 30 ml). Identification and counting of particles was performed on an OLYMPUS BX-51 transmitted light microscope, connected to a camera (Altra20 Soft Imaging System). For identification of palynomorphs and palynodebris relevant literature was consulted.

2.8 Calcareous nannofossil biostratigraphyCalcareous nannofossil preparation and identification used for biostratigraphy of the Polish site was done by Sebastian Pauly, Ruhr-University Bochum, Germany. Investigation of taxa was performed on smear slides based on standard procedures, by using an OLYMPUS BH-2 light microscope with cross-polarized light. Relevant literature was consulted for identification of calcareous nannofossils.

(For more information see respective chapters.)

ReferencesEspitalié, J., Laporte, J.L., Madec, M., Marquis, F., Leplat, P., Paulet, J., Botefeu, A., 1977. Méthode rapide de caractérisation des roches metres, de leur potential pétrolier et de leur degré d’évolution. Oil and Gas Science and Technology 32 (1), 23-42,

Kutek, J., Marcinowski, R., Wiedmann, J., 1989. The Wąwał Section, central Poland – an important link between Boreal and Tethyan Valanginian. Cretaceous of the Western Tethys (ed. Wiedmann, J.), Proceedings of the 3rd International Cretaceous Symposium, Tübingen 1987, pp. 717-754 (Schweizbart, Stuttgart).

Lüniger, G., Schwark, L., 2002. Characterization of sedimentary organic matter by bulk and molecular geochemical proxies: an example from Oligocene maar-type Lake Enspel, Germany. Sedimentary Geology 148, 275-288.

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3. No evidence for anoxia during the Valanginian carbon isotope event – An organic-geochemical study from the

Vocontian Basin, SE France*

AbstractThe Valanginian time interval (Early Cretaceous) is characterized by a positive carbon isotope excursion (CIE) which represents the first of several prominent Cretaceous δ13C anomalies. A combined chemostratigraphic and organic-geochemical approach has been chosen to investigate the composition and distribution of sedimentary organic matter (OM) deposited before the Valanginian CIE, during its onset and plateau-phase. This was done to test whether this CIE is accompanied by changes in marine primary production and/or OM preservation. Biostratigraphically well-calibrated deposits from two hemipelagic sections located in the Vocontian Basin of SE France are used as sedimentary archives. A newly established high-resolution δ13C record covering the composite succession shows a characteristic Valanginian pattern and enables a detailed correlation with existing carbon isotope curves from the northern Tethyan margin. The analyzed solvent extractable fraction of the sedimentary OM is mainly composed of a marine origin with an admixture of land plant material. Variations in specific biomarkers for cyanobacteria (2α-methyl-hopanes), dinoflagellates (dinosterane or 4-desmethyl-23,24-dimethyl steranes) and terrigenous plant-derived OM (odd numbered long-chain n-alkanes) as well as the sterane/hopane ratio, the C35 hopane index and the isoprenoids pristane and phytane were investigated. In contrast to the well-studied mid-Cretaceous Oceanic Anoxic Events (OAEs), neither significant OM enrichment nor prominent fluctuations in the selected biomarker abundances can be observed during the build-up phase of the Valanginian CIE. This points to relatively stable marine paleoenvironmental conditions with well-oxygenated bottom waters. Prior to the CIE, four cm-thick, finely laminated, dark layers (known as Barrande layers) with total organic carbon content reaching up to 4 % show an exception from the generally stable biomarker pattern. Sedimentological and biomarker evidence support deposition under less oxygenated conditions for the Barrande layers. However, their occurrence clearly predates the onset of the positive δ13Ccarb shift (by about 180 kyrs). Contrary to the subsequent mid-Cretaceous CIEs, the occurrence of widespread anoxia associated with the Valanginian CIE cannot be confirmed for the Vocontian Basin.

Keywordsoceanic anoxic events; carbon isotopes; organic matter; Cretaceous; Valanginian; biomarkers* This chapter is published as:Kujau, A., Heimhofer, U., Ostertag-Henning, C., Gréselle, B., Mutterlose, J., 2012. No evidence for anoxia during the Valanginian carbon isotope event - An organic-geochemical study from the Vocontian Basin, SE France. Global and Planetary Change 92-93, 92-104.

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3.1 IntroductionMajor perturbations of the Cretaceous global carbon-cycle are reflected in long- and short-term fluctuations of the δ13Ccarb and δ13Corg records, commonly associated with the deposition of organic-rich sediments in the world oceans. A major positive carbon isotope excursion (CIE) covering the Lower Valanginian to Lower Hauterivian time-interval is considered as the first of several high-amplitude shifts in the carbon isotope record after a time of relatively stable δ13C values prevailing during the Late Jurassic-earliest Cretaceous (Weissert and Erba, 2004; Föllmi et al., 2006). The Valanginian CIE displays a significant positive δ13C anomaly with amplitudes of ~1.5 to 2.5 ‰ for marine carbonates (Cotillon and Rio, 1984; Lini et al., 1992; Hennig et al., 1999; Wortmann and Weissert, 2000; Erba et al., 2004; Sprovieri et al., 2006; McArthur et al., 2007; Gréselle et al., 2011), ~2 to 3 ‰ for organic matter (OM) of marine origin (Cotillon and Rio, 1984; Lini et al., 1992; Wortmann and Weissert, 2000), and ~4 to 5 ‰ for terrestrial plant-derived OM (Gröcke et al., 2005, Nunn et al., 2010). This earliest Cretaceous CIE has been postulated to be associated with an Oceanic Anoxic Event (OAE), also known as the “Weissert anoxic event” (Erba et al., 2004). OAEs represent major perturbations of the ocean-atmosphere system and are characterized by increased accumulation of organic-rich deposits due to enhanced preservation and/or productivity on a supra-regional to global scale (Schlanger and Jenkyns, 1976; Jenkyns, 2010). For the Valanginian event increased levels of atmospheric CO2 and subsequent climate warming due to enhanced volcanic activity of the Paraná-Etendeka Large Igneous Province (Brazil-Angola) were proposed as a potential cause for the observed paleoenvironmental changes (Lini et al., 1992; Price and Mutterlose, 2004; Weissert and Erba, 2004). A thereby enhanced accelerated hydrologic cycle could have caused high nutrient inputs from continental sources, probably augmented by the introduction of biolimiting metals at spreading ridges, to foster high primary productivity in ocean waters and, in consequence, increased organic carbon burial (Weissert, 1989; Föllmi et al., 1994; Weissert et al., 1998; Weissert and Erba, 2004, Föllmi et al., 2006). In addition, excess CO2 levels in ocean surface waters have been used to explain a decline in nannoconid abundance in the western Tethyan realm and widespread carbonate platform demise, a phenomenon referred to as “biocalcification crisis” (Wortmann and Weissert, 2000; Erba and Tremolada, 2004; Weissert and Erba, 2004). On the other hand, the Late Valanginian has been identified as a phase of substantial cooling. Evidence for a change towards a cooler climate is based on geochemical analysis (Weissert and Lini, 1991; Podlaha et al., 1998; Pucéat et al., 2003; McArthur et al., 2007; Brassell, 2009; Price and Nunn, 2010), changes in plankton assemblages (Mutterlose and Kessels, 2000; Melinte and Mutterlose, 2001; Mutterlose et al., 2003), and indirect evidence for the establishment of high-latitude ice caps in the form of ice-rafted debris and glendonites (Kemper, 1987; Frakes and Francis, 1988; Price, 1999; Gréselle and Pittet, 2010, and references therein). However, a recent study assumes rather stable Valanginian sea-surface temperatures for the Atlantic Ocean based on biomarker studies (Littler et al., 2011).The underlying causes for the Valanginian CIE and associated paleoenvironmental changes are still a matter of debate (Erba et al., 2004; Gröcke et al., 2005; McArthur et al., 2007; Westermann et al., 2010; Gréselle et al., 2011). In the Valanginian, only few centimetre-thick OM-rich layers have been identified in the Tethyan realm (Reboulet et al., 2003; Erba et al., 2004; Westermann et al., 2010), which precede a low-amplitude negative shift that in turn initiates the onset of the positive CIE. However, a causal link between these few organic-rich layers and the Valanginian CIE has been called into question (e.g., Bornemann and Mutterlose, 2008; Westermann et al., 2010). In contrast to the subsequent Cretaceous CIEs (OAE1a, OAE1b, OAE2) only a single detailed geochemical study of the sedimentary OM deposited in the North Atlantic during the

No evidence for anoxia during the Valanginian carbon isotope event - An organic-geochemical study from the Vocontian Basin, SE France

27

Valanginian CIE is currently available (Brassell, 2009). To date, biomarker data from Tethyan localities are absent. This study investigates the composition and distribution of the sedimentary OM deposited prior to the Valanginian CIE and during its onset and plateau-phase. The chosen approach combines sedimentological and chemostratigraphic information with geochemical analysis of the bulk OM. The overall aim of this study is to provide information on changes in marine primary production, depositional conditions, and OM degradation during a time of major paleoenvironmental and climatic perturbations. For this study, hemipelagic sediments deposited in the Vocontian Basin (SE France) covering Lower Valanginian to Upper Valanginian strata (Busnardoites campylotoxus to Saynoceras verrucosum ammonite zones; Hoedemaeker et al., 2003; Reboulet et al., 2006) have been analyzed with regard to the preserved sedimentary OM.

3.2 Geological settingThe Vocontian Basin of SE France is known for its well exposed Cretaceous successions and represents a “classic” site for studying Early Cretaceous paleoceanography. During the Early Cretaceous, this basin was located at a paleolatitude of ~30°N in a marginal marine position of the Western Ligurian Tethys Ocean, opened to the east (Masse, 1993; Hay et al., 1999; Fig. 3.1A). The central part of the Vocontian Basin is characterized by hemipelagic deposits composed of open-marine autochthonous carbonates and marls, carbonate fine-fraction exported from three surrounding platforms (Reboulet et al., 2003; Gréselle and Pittet, 2010), and terrigenous material, mostly eroded from the Mid-European continents (Bréhéret, 1994; Fesneau et al., 2009).

A)

NEOTETHYS

Mid-European continents

Africa

VocontianBasinATLANTIC

Ligurian Ocean

continentsshallow marine seaopen marine sea

N

B)

6°

Marseille

Ain

Isère

Drôme

Aygues

Durance Verdon

46°

45°

44°

43°

Toulouse

Limoges

4°

MorenasVergol

Lyon

Platform marginVocontian Basinstudied sections

2°

Mediterranean Sea100 km

20°

30°

20°

30°

40°

Fig. 3.1. A) Reconstruction of the location of the Vocontian Basin within the Tethyan Ocean during the Early Cretaceous. B) Simplified map of southeast France with an excerpt of a geologic map showing the paleogeographic extension of the Vocontian Basin (modified after Gréselle et al., 2011). Black asterisks refer to the studied sections, for colour codes see legends within maps.

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For this study two sections have been sampled including the Vergol and Morenas sections (Fig. 3.1B). The resulting composite succession has a total thickness of ~63 m and covers the Lower Valanginian B. campylotoxus to Upper Valanginian S. verrucosum zones (Reboulet et al., 1992; 2003; Reboulet and Atrops, 1999, Gréselle and Pittet, 2010; Fig. 3.2A). In outcrop, the sampled deposits are rich in marine invertebrate fossils and consist of well exposed marl-limestone alternations that are stacked in bundles. These bundles are hierarchically organized in three orders of depositional cyclic sequences and their formation was suggested to be controlled by changes in orbital parameters (Gréselle and Pittet, 2010). The Vergol section (E 5°25’9’’, N 44°12’12’’ is located in the Drôme department and crops out between the villages Montbrun-les-Bains and Vergol. Logged and sampled strata (thickness of ~57 m) cover sedimentary rocks from the Lower Valanginian B. campylotoxus Zone to the top of the S. verrucosum Subzone, S. verrucosum Zone, Upper Valanginian. The lower part of the succession is dominated by calcareous marls whereas the upper part consists mainly of argillaceous sediments. Between ~25 and 30 m the Vergol section exhibits an interval characterized by intense slumping just below a bundle of four calcareous marker beds informally called “tétrade” (Gréselle, 2007; Fig. 3.2A), which mark the transition from the Lower to the Upper Valanginian. To compensate for the resulting hiatus, the slumped part at the Vergol section has been covered by samples collected at the nearby and well correlated Morenas section (Gréselle and Pittet, 2010; Gréselle et al., 2011). Between ~17.5 and 19.5 m, the Vergol section contains four cm-thick, finely laminated, non-bioturbated, dark layers, known as Barrande layers 1 to 4 (Reboulet, 2001; Reboulet et al., 2003, Fig. 3.2A). Similar conspicuous layers are e.g. also reported from the Angles section, located ~100 km northwest of Vergol (Reboulet et al., 2003) and in the form of two total organic carbon (TOC)-enriched marly layers in the Lombardian Basin (Breggia, North Italy, Bersezio et al., 2002), located ~600 km southeast of Vergol.The Morenas section (E 5°25’23’’, N 44°13’52’’) covers mainly argillaceous sediments (thickness of ~5 m) of the B. campylotoxus Zone and crops out ~1 km northwest of the city of Aulan close to the village of Morenas. It is located ~2 km north of the Vergol section.

Fig. 3.2. A) Ammonite biostratigraphy, cyclostratigraphy, thickness, lithology (see legend for details; Gréselle et al., 2011) and position of samples for carbon isotopes of the composite Vergol/Morenas section plotted against geochemical parameters. B1-4 indicate position of the Barrande layers, broken vertical line marks position of interval named “tétrade“, geochemical parameters include δ13Ccarb, CaCO3 %, TOC %, and hydrogen index (HI). Roman numbers to the left (I to III) correspond to the characteristic segments of the CIE. B) Cross plot of δ13C versus δ18O, R2 = 0.003.

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3.3 Methods and MaterialsA total of 97 samples have been collected regularly distributed throughout the entire succession, resulting in a sample spacing of ~0.5 m/sample (Vergol: 87 samples; Morenas: 10 samples). Samples have been taken exclusively from marly interbeds in order to avoid lithology-related diagenetic effects (Sprovieri et al., 2006). To minimize contamination by modern plant material and/or near-surface weathering processes, the uppermost 15-20 cm of sediment have been removed before sampling. Sample material has been dried at 50°C in a laboratory oven before crushing and homogenization in an agate mortar.

3.3.1 Stable carbon and oxygen isotopesMeasurements of stable carbon and oxygen isotopes of sedimentary carbonates have been carried out on powdered bulk sample material (~0.6 mg) on all 97 samples. A Thermo Fisher Scientific Gasbench II carbonate device connected to a Thermo Fisher Scientific Delta S isotope ratio mass spectrometer, available at the Ruhr-University Bochum, Germany, has been used for stable isotope analyses. The gas bench uses 100 % phosphoric acid at 70°C to release CO2 of the calcite from the sample material 1 h before the start of the measurement. Repeated analyses of certified carbonate standards (CO-1, CO-8, NBS-19) show an external reproducibility ± 0.1 ‰ for δ18O and ± 0.06 ‰ for δ13Ccarb. Values are expressed in conventional delta notation relative to the Vienna Pee Dee Formation belemnite (VPDB) international standard, in per mil (‰). 31 duplicate measurements show that the measured values are representative and indicate that the samples are quite homogenous (max. dev. δ18O ± 0.17 and δ13C ± 0.12).

3.3.2 Carbonate carbon, total organic carbon and RockEval pyrolysisTotal inorganic carbon (TIC) has been measured on 46 samples distributed regularly throughout the section. TIC % values have been determined via decarbonatization. An aliquot of 2 g of the powdered bulk material was treated with 6 M HCl and neutralized with deionized water to remove the carbonate. The weight difference before and after acid treatment and drying of the residue has been used to calculate the CaCO3 content in %. A total of 37 samples have been analyzed for their TOC content using a LECO CS 200 at the BGR Hannover, Germany. To determine type and thermal maturity of the sediment and to assess whether migrated bitumen is present, samples with TOC contents of ≥0.2 % have been analyzed using a Rock-EvalTM6 at the BGR Hannover, Germany. The parameters quantified during pyrolysis include S1, S2, S3, Tmax, PI, HI and OI (following Espitalié et al., 1977). The S1 and S2 peaks correspond to the maximum of hydrocarbon generation in mg/gsed, with S1 representing hydrocarbons volatilized from the powdered material and S2 representing hydrocarbons mainly formed during thermal cracking of the remaining kerogen. S3 corresponds to the amount of CO2 generated from the OM. Tmax (in °C) is the temperature of maximum rate of hydrocarbon generation during the pyrolysis. Production index (PI) is calculated from S1/(S1+S2). Hydrogen index (HI) is given in mg HC/g TOC and oxygen index (OI) in mg CO2/g TOC.

3.3.3 Biomarker analysisFor the samples selected for molecular organic-geochemical analyses an aliquot of 15-28 g powdered sample material was extracted with dichloromethane using a Dionex accelerated solvent extractor (ASE 200). Elemental sulfur was removed with activated copper. The asphaltenes were separated from the bulk extract by precipitation in petroleum ether. The remaining maltene and resins fraction was weighed and

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further fractionated into an aliphatic, an aromatic and a heterocompound fraction using medium-pressure liquid chromatography with solvents of increasing polarity. The aliphatic fraction was analyzed by gas chromatography (Agilent 6890 GC-FID) using a flame ionization detector signal for quantification. The separation of the compounds was achieved with a 60 m DB-1 column (ID 0.32 mm, film thickness 0.2 µm) using a temperature program of 50°(2min)-3°/min-320°(10min) and a constant carrier gas stream of helium of 1 ml/min. Quantification was achieved for all compounds using standards of n-alkanes and their response factors. For the identification and quantification of biomarkers gas chromatography-mass spectrometry (Agilent 6890-Finnigan MAT 95S) was employed using both full scan and metastable reaction monitoring modes (MRM). The GC conditions were identical to those in the above mentioned GC measurements. The scan range was m/z 50-550 with a cycle time of 0.7/s and a mass resolution of 1000. The MRM measurements monitored the transitions of specific parent ions to daughter fragments of the triterpane series (m/z 191 for hopanes, m/z 205 for methylhopanes, m/z 217 for steranes, and m/z 231 for methylsteranes). The composition of hopanoid and steroid biomarkers are reported as ratios of biomarkers.

3.4 Results3.4.1 Stable carbon isotope signatureThe carbon isotope (δ13Ccarb) values vary between 0.1 and 2.6 ‰ and show a positive CIE within the interval investigated (Fig. 3.2A). The carbon isotope record does not show an offset of δ13Ccarb values at the base or top of the different sections. Hence, a composite curve can be established. The carbon isotope record is interpreted to reflect a primary signal, which is supported by the characteristic trend of the CIE and by the lack of correlation between the δ18O and the δ13C values (Vergol and Morenas sections, R2=0.003, Jenkyns, 1996, Fig. 3.2B). In this study, the CIE is dissected into three segments including the pre-excursion interval (segment I, 0-26.1 m), the positive shift (segment II, 26.2-50.1 m) and the plateau-phase (segment III, 50.2-63 m).In the B. campylotoxus and Karakachiceras biassalense ammonite subzones (0 to 26.1 m, segment I) pre-excursion carbon isotope values oscillate between 0.5 and 1.0 ‰. Starting with an initial negative peak in the K. biassalense Subzone (most negative value: 0.1 ‰ at 26.1 m), the CIE is characterized by a two-step positive excursion with an overall shift of ~2.5 ‰ (segment II). A first increase (26.2 to 35.2 m) from 0.1 to 1.7 % covers the upper K. biassalense Subzone and is followed by an initial plateau (35.5 to 44.9 m) right at the Lower to Upper Valanginian boundary. The second increase (45.6 to 50.1 m) from 1.9 to 2.5 % occurs within the S. verrucosum Subzone and is followed by a second plateau-phase (50.2 m up section, segment III). The interval covered by the positive CIE is clearly dominated by more argillaceous sediments compared to the lower, carbonate-rich part of the record.

3.4.2 Carbonate carbon, organic carbon content and RockEval pyrolysisCaCO3 contents fluctuate between 41 and 78 % with an average value of 60 % (Fig. 3.2A, Table 3.1). Increased values are observed at the base of the composite succession (B. campylotoxus Zone), whereas the uppermost part shows values below 55 %. The studied deposits are comparatively poor in organic carbon, TOC contents range between 0.2 and 0.7 % (avg. 0.5 %). Elevated TOC contents ranging between 2.8 and 4.0 % are restricted to samples from the Barrande layers. RockEval pyrolysis exhibits HI values between 153 and 259 mg HC/g TOC and OI values ranging from 55 to 126 mg CO2/g TOC (Fig. 3.2A, Table 3.1). Again, the Barrande layers deviate from this general pattern with elevated HI values of 240 to 383 mg

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Table 3.1: Geochemical data shown according to depth level from Vergol and Morenas

Depth Sample CaCO3 TOC S1 S2 S3 Tmax PI HI OI δ13C (m) (%) (%) (mg/g) (mg/g) (mg/g) (°C) (S1/[S1+S2]) ([100*S2]/TOC) ([100*S3]/TOC) VPDB[%]

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Table 3.1 (continued)

Fig. 3.3. Van Krevelen diagram depicts samples as type III kerogen with medium maturity. Samples derived from the Barrande layers correspond to white dots, non-Barrande layer samples are represented by white dots.

HI (

mg

HC

/g T

OC

)

OI (mg CO2/g TOC)

0

800

700

600

500

400

300

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100

400300200100

Barrande layer

Type III

Type II

Type Iother samples

Depth Sample CaCO3 TOC S1 S2 S3 Tmax PI HI OI δ13C (m) (%) (%) (mg/g) (mg/g) (mg/g) (°C) (S1/[S1+S2]) ([100*S2]/TOC) ([100*S3]/TOC) VPDB[%]

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HC/g TOC and lower OI values of 19 to 39 mg CO2/g TOC. Tmax values vary between 430°C and 440°C (avg. 435°C, Table 3.1). Plotted in a van Krevelen diagram (Fig. 3.3), the OM of the analyzed samples can be assigned to a marine type II kerogen, close to type III, of low to moderate maturity (HI above ~ 170 mg HC/g TOC, Lüniger and Schwark, 2002). The low PI values (avg. 0.04) indicate that no migrated bitumen is present in the investigated samples (e.g. Schwark et al., 2009).

3.4.3 Biomarker analysisAll biomarkers of the aliphatic fraction (excluding n-alkanes) discussed in this contribution are listed in Table 3.2, an example of GC-MRM-MS traces for the hopanes and steranes is shown in Fig. 3.4. Only those biomarkers considered to be relevant for this study regarding marine primary production, depositional conditions and OM degradation are mentioned in the text. Minimum quantified hydrocarbon chain length was nC15, maximum chain length nC40 (Fig. 3.5). Samples are dominated by odd short-chain n-alkanes of 17 to 21 carbon atoms (Fig. 3.5, pre-Barrande). Furthermore, n-alkanes of odd chain lengths of nC25 to nC31

dominate over their even-numbered neighbours pointing towards a small contribution of terrigenous OM. The Barrande layers are exceptionally rich in isoprenoids and show high abundances of pristane compared to n-alkanes (Fig. 3.5, Barrande). All identified configurations of steranes occur with S and R isomers, with the 5α(H),14α(H),17α(H)R and 13β(H),17α(H)S isomers being the most abundant ones. The 17α(H),21β(H)S hopane isomers are the most abundant hopanes. The isomerisation of homohopanes at C-22 with %C22 values of 0.58 to 0.60 has reached equilibrium and points to a maturity of above 0.5 % vitrinite reflectance (Peters and Moldowan, 2005). But the isomerisation of regular C29 steranes at C-20 with %C20S values of 0.35 to 0.4 and at C-14 and C-17 with %ββ values of 0.3 to 0.34 have not reached equilibrium and depict a maturity range before the onset of or in the early oil window (Peters and Moldowan, 2005). This confirms the maturity assessed by

Table 3.2: List of investigated aliphatic biomarkers.

Peak Identification steranes Identification hopanes

1 5α,14α,17α-27-Norcholestane 20R 18α(H)-22,29,30-Trisnorhopane2 13β,17α Diacholestane 20S 18α-30-Norhopane3 13β,17α Diacholestane 20R 17α,21β(H)-Hopane4 13α,17β Diacholestane 20S 17α,21β(H)-Homohopane 22 S5 13α,17β Diacholestane 20R 17α,21β(H) Homohopane 22 R6 5α,14α,17α-Cholestane 20S 17β,21α (H) Homomoretane 22 S7 5α,14β,17β-Cholestane 20R 17β,21α (H) Homomoretane 22 R8 5α,14β,17β-Cholestane 20S 3β(H)-Methylhopane9 5α,14α,17α-Cholestane 20R 3β(H)-Methylhopane10 13β,17α-Diaergostane 20S 17α(H),21β(H) Bishomohopane 22 S11 13β,17α-Diaergostane 20S 17α(H),21β(H) Bishomohopane 22 R12 13β,17α-Diaergostane 20R 2α Methyl-homohopane S13 13β,17α-Diaergostane 20R 2α Methyl-homohopane R14 13α,17β-Diaergostane isomer 3β Methyl-homohopane S15 13α,17β-Diaergostane isomer 3β Methyl-homohopane R16 5α,14α,17α-Ergostane 20S 17α(H),21β(H) Pentakishomohopane 22S17 5α,14β,17β Ergostane 20R 17α(H),21β(H) Pentakishomohopane 22R 18 5α,14β,17β Ergostane 20S19 5α,14α,17α-Ergostane 20R Identification isoprenoids20 13β,17α-Diastigmastane 20S 21 13β,17α-Diastigmastane 20R Pristane (2,6,10,14-tetramethylpentadecane)22 13α,17β-Diastigmastane isomer Phytane (2,6,10,14-tetramethylhexadecane)23 13α,17β-Diastigmastane isomer24 5α,14α,17α-Stigmastane 20S25 5α,14β,17β-Stigmastane 20R26 5α,14β,17β-Stigmastane 20S27 5α,14α,17α-Stigmastane 20R28 5α,14α,17α-24-n-Propylcholestane 20S29 5α,14α,17α-24-n-Propylcholestane 20R30 4-methyl-24-Ethylcholestane isomer31 4-methyl-24-Ethylcholestane isomer32 4 α,23,24-Trimethylcholestane isomer (Dinosterane)33 4 α,23,24-Trimethylcholestane isomer (Dinosterane)34 4-methyl-24-Ethylcholestane isomer35 4 α,23,24-Trimethylcholestane isomer (Dinosterane)

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Fig. 3.4. MRM traces for (A) hopanes and (B) steranes. See Table 3.2 for identification of numbered peaks.

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Fig. 3.5. Total ion chromatograms of the aliphatic compounds of a pre-Barrande layer sample (left) compared to a Barrande layer sample (right).

Barrande layerpre-Barrande layerpristane

phytane

nC18nC17 pristane

phytane

nC18nC17

dete

ctor

resp

onse

[pA]

50

100

150

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250

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retention time [min]10 15 20 25 30 35 40 45 50 55 60

retention time [min]10 15 20 25 30 35 40 45 50 55 60

RockEval pyrolysis and rules out a contamination by recent vegetation or soil organic matter. Changes in biomarkers are described according to the three characteristic segments based on changes in the δ13C isotope record, which correspond to pre-CIE-, onset-, and plateau-phase (Fig. 3.2A, 3.6). The ratio of regular steranes to hopanes exhibits maximum values in the Barrande layers (above 1.0, segment I) with otherwise relatively low values that increase slightly in segment III. The ratio of dinosterane over regular steranes depicts low values in the Barrande layers (close to 0.2) and elevated values in the sediments deposited after these black shale layers (around 0.4, segment I, II). Highest values occur in segment III (above 0.4). The ratio of 2α-methyl- hopanes to homohopanes is overall low (between 0 and 1.0) with slightly elevated values at the base of the first segment and in the Barrande layers. Another maximum is evident in segment III (all around 2.0). The ratio of C35/(C31+C35) homohopanes shows overall low values (around 0.06) with a maximum around and within the Barrande layers (above 0.07, segment I). The ratio for pristane/phytane remains relatively constant throughout all segments (~ 3). Again, the Barrande layers differ from this general pattern and display highest values reaching up to 4.5.

Fig. 3.6. Ammonite biostratigraphy and thickness of the composite Vergol/Morenas section plotted against δ13Ccarb, steranes/hopanes, dinosterane/regular steranes, 2-methyl-hopanes/homohopanes, C35/(C31+C35),αβ-hopanes and pristane/phytane. Roman numbers to the left (I to III) correspond to the characteristic segments of the CIE.

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3.5 Discussion of results3.5.1 Comparison of the δ13Ccarb record to existing Tethyan curvesThe high-resolution δ13Ccarb record allows detailed chemostratigraphic correlation and comparison with existing δ13Ccarb records covering the Valanginian CIE event. The curve of this study does not exactly match the published δ13Ccarb data from the same locality (Gréselle et al., 2011; Fig. 3.7). Differences between both curves are probably due to a significantly lower sample spacing of the previous record and the use of mixed lithologies including limestone beds and marls. Different sources of the carbonate (terrestrial input, hemipelagic and platform derived material from different producers), and differing diagenetic alteration processes may further explain for the observed deviations. Varying contributions of carbonate from different sources to the marls exclusively analyzed in this study may also explain some of the small-scale fluctuations in the δ13Ccarb record presented here. Comparison and correlation with an existing δ13Ccarb record derived from the Angles section, Vocontian Basin, France (Duchamp-Alphonse et al., 2007), and with a δ13Ccarb record from the Capriolo section, Lombardian Basin, Italy (Channell et al., 1993), shows strong similarities as well as some distinct differences (Fig. 3.7). The Angles section is located within the same basin (about ~130 km south-eastward) and therefore the

Fig. 3.7. Correlation of carbon isotope trends from the Vergol/Morenas sections, Vocontian Basin, SE France (Gréselle et al., 2011 and this study) with the Angles section record, Vocontian Basin, SE France (Duchamp-Alphonse et al., 2007) and with a record from the Lombardian Basin, Italy (Channell et al., 1993).

excellent match between the two δ13Ccarb curves is not surprising. Despite differences in sample spacing, both records show the very same trends and shifts as well as similar absolute values. This allows for a detailed correlation of certain features of the two curves including a negative shift in the B. campylotoxus Zone and a maximum in the middle part of the S. verrucosum Subzone. Differences in small-scale structure of the two curves may to some extend be explained by local oceanographic phenomena (Royer et al., 2001 and references therein), local diagenetic effects, or terrestrial input that can mask the global carbon isotope signature. The occurrence of a short-lasting negative peak initiating the positive CIE is consistent with previously reported carbon isotope records (Lini et al., 1992; Erba et al., 2004; McArthur et al., 2007; Gréselle et al., 2011).

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Despite an overall similar shape, detailed comparison of the records from the Vocontian and the Lombardian basins (palaeodistance of ~ 600 km) reveals some discrepancies. The absolute values of the Capriolo record are in general 0.5 to 1.0 ‰ higher than those observed for the Vocontian Basin. Furthermore, the Capriolo record lacks the low-amplitude negative shift before the onset of the positive anomaly and shows a reduced scatter compared to the French records. The latter may be due to the fact that the Lombardian Basin represents a more stable pelagic system while the Vocontian Basin is hemipelagic and close to the continental margin. The difference in absolute values and the absence of small-scale variations hampers correlation of the two isotope curves on a high resolution. With regard to the biostratigraphic position of the onset of the CIE, the dataset from the Vocontian Basin resembles that from the Capriolo record. In all three records, the initiation of the CIE is located in the upper part of the B. campylotoxus Zone. This affirms the assumption of a temporal synchrony of major shifts in the global carbon isotope reservoir. However, despite the differences in absolute values, all the three records show a more or less similar overall trend for the Valanginian CIE, confirming the primary nature of the carbon isotope signal and its applicability as a correlation tool.

3.5.2 Specific sources of the OMThe organic-geochemical data derived from bulk kerogen and extractable biomarkers allows identification of different sources contributing to the sedimentary OM of the Vergol section. Particular focus is on variable contributions of marine algal and microbial groups and on terrestrial plant-derived OM input.RockEval pyrolysis data point to a type II kerogen of predominantly autochthonous marine origin probably admixed with a certain terrestrial OM fraction (of type III kerogen, Fig. 3.3). A similar composition has been reported from sections in the Vocontian Basin, the Ligurian Ocean, and Shatsky Rise in the Pacific Ocean (Bersezio et al., 2002; Westermann et al. 2010). A dominantly terrestrial OM (type III kerogen) has only been observed in the Carpathian seaway (Wąwał, central Poland) for the time interval of interest (Westermann et al. 2010). This general picture of predominantly marine OM contribution is confirmed by the dominance of short-chain odd-numbered n-alkanes, which similarly indicate high input from aquatic primary producers (Blumer et al., 1971, Poynter and Eglington, 1990; Meyers, 1997; Jeng and Huh, 2006). More specific information about variations in abundance of algal and microbial organism groups can be obtained from biomarkers of the aliphatic fraction including terpenoids of the hopane and sterane series. The ratio of steranes/hopanes gives a hint to the relative importance of eukaryote algae versus prokaryote-derived, OM contribution (Peters and Moldowan, 2005). The identified 2α-methyl-hopanes are presumably derived from cyanobacteria (Farrimond et al., 2004; Dumitrescu and Brassel, 2005). Variations in the abundance of cyanobacteria are generally interpreted to reflect changes in nutrient availability, with an increase in abundances pointing to nitrate-limited conditions within the photic zone (Dumitrescu and Brassel, 2005). The 4-methyl steranes, especially dinosterane (4-desmethyl-23,24-dimethyl sterane, here identified in the form of three isomers) are rather specific for dinoflagellates, but can also be produced by diatoms (Rampen et al., 2009, and references therein). Since diatoms did not diversify significantly until the Late Cretaceous (Round et al., 1990) dinosterane can be assigned to dinoflagellate input in this case. To assess the contribution and composition of the terrestrial OM fraction, the samples were analyzed for potential land plant biomarkers found elsewhere like retene and cadalene (van Aarssen et al., 1999; Hautevelle et al., 2006), ip-iHMN (van Aarssen et al., 1999), oleanane, ursane and lupane (Otto et al., 2005; Simoneit et al., 2003), dammar resin (van Aarssen et al., 1990), labdane and abietane (Bray and Anderson, 2008; Pereira et al., 2009), and others. No source-specific terrestrial biomarkers were, however,

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identified within the aliphatic and aromatic fractions of the analyzed bitumen. Nevertheless, the distribution of non-branched hydrocarbons (n-alkanes) allows for some statements regarding compositional variations in the terrestrial OM fraction. The observed odd-over-even predominance of long-chain (nC25-nC35) n-alkanes in the section points to a certain contribution of terrestrial plant remains to the OM, since vascular land plant cuticular waxes typically show this distribution pattern (Eglington and Hamilton, 1967; Gearing et al., 1976; Farrington, 1980; Cranwell, 1982). In summary, bulk and molecular data point to a predominantly marine origin of the bulk of the sedimentary OM, derived from various phytoplankton sources including dinoflagellate cysts as well as cyanobacteria. The lack of specific land plant biomarkers and low concentrations of long-chained n-alkanes support the conclusion that land plant input was generally small.

3.5.3 Depositional conditions during the CIE in the Vocontian BasinGiven the fact that the Vocontian Basin was a marginal marine basin, with only a narrow opening towards the Tethys, and characterized by a water depth of only about 300 m (Wilpshaar and Leereveld, 1994; Wilpshaar et al., 1997), as well as substantial detrital input (e.g. Bréhéret 1997; Fesneau et al., 2009), conditions would be considered favourable for the establishment of anoxia and therewith for enhanced OM preservation. During the Lower Aptian and Albian, numerous black shales have formed in the Vocontian Basin (e.g. Bréhéret 1997), which highlights the potential of this basin to act as a site for black shale formation. Overall low TOC values (avg. 0.65 %) of the Vergol/Morenas sediments, however, point to rather poor OM preservation during the Valanginian CIE in the Vocontian Basin, probably reflecting deposition under a generally well-oxygenated water column. Similarly low to moderate TOC contents (avg. 0.24 %) have been reported by Westermann et al. (2010) for the Angles section of the Vocontian Basin. Alternatively, low OM accumulation rates due to limited primary production or low terrestrial influx of nutrients may account for the observed pattern. However, there is evidence for enhanced terrestrial influx into the Vocontian Basin during the initiation of the positive δ13C shift. This is for example based on enhanced concentrations of the major elements iron (Fe) and manganese (Mn), which have been interpreted to indicate elevated continental weathering under humid climatic conditions (Kuhn et al., 2005). In addition, this enrichment was also suggested to be an indication for dys- to anoxic conditions by the same authors. The latter interpretation is however called into question (e.g. Westermann et al., 2010 and references therein). Enrichment in Fe and Mn in marine sediments can also occur under oxic conditions during times of high terrestrial input (Westermann et al., 2010). Comparatively high TOC values (up to 4.00 %) occurring within the Barrande layers of segment I, indicate short-termed episodes of anoxia (Westermann et al. (2010). Additional information on bottom water oxygenation during OM deposition can also be obtained from the C35 homohopane index. The ratio of C35/(C31+C35) homohopanes shows an abrupt increase in the upper part of segment I with highest values occurring within the Barrande layers. In segments II to III, a gradual decline can be observed. Enhanced values have been used to identify anoxic conditions, but this interpretation was called into question (ten Haven et al., 1988). In the Vocontian Basin, the observed increase of the C35 homohopane index within the Barrande interval is in good accordance with sedimentological evidence for a severely reduced oxygenation of the bottom waters. Supplementary information on the paleo-redox regime during OM deposition is provided by the ratio of the acyclic isoprenoids pristane and phytane. The early diagenetic break-down of phytoplanktonic chlorophyll

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results in the generation of pristane (under aerobic conditions) or phytane (under anaerobic conditions) depending on the redox conditions prevailing in the water column (Didyk et al., 1978; Peters and Moldowan, 2005). For the studied samples, the pristane/phytane (Pr/Ph) ratio shows relatively consistent values above 2, indicating deposition under oxic conditions in all three segments. Some outliers of the Pr/Ph ratio (>4) occurring in the Barrande layers of segment I are probably due to differences in depositional conditions and may reflect a specific preservation pathways for these isoprenoids. The significance of the Pr/Ph ratio has, however, been questioned due to variable sources for both compounds including methanogenic bacteria (ten Haven et al., 1985; 1988, Koopmans et al., 1999; Frimmel et al., 2004). Isorenieratane and other diagenetic products of isorenieratene have not been identified in the aromatic fraction of any of the analyzed samples. These compounds are considered to be indicative for a stratified water column and the establishment of euxinic waters reaching the photic zone (Summons and Powell, 1986; Grice et al., 1996; Schwark and Frimmel, 2004; Kenig et al., 2004; Heimhofer et al., 2008). Isorenieratene derivatives are ubiquitous compounds in many mid-Cretaceous black shales deposited during OAEs and have been reported from the Cenomanian-Turonian OAE2 (Sinninghe-Damste and Köster, 1998; Kuypers et al., 2002) as well as from the Coniacian-Santonian OAE3 (Wagner et al., 2004). The absence of these specific compounds in the Valanginian deposits of the Vocontian Basin clearly indicates that the environment of deposition did not reach euxinic conditions.

3.5.4 A close-up of the Barrande layersThe four approximately one cm-thick Barrande layers (at 17.5, 17.6, 17.8 and 18.8 m, Lower Valanginian) are of interest due to their undisturbed lamination, organic carbon enrichment and unusual geochemical signature. This fine-scale lamination pattern can be interpreted to reflect unfavourable conditions for bottom-dwelling detritivores. The Barrande layers occur stratigraphically ~6 m below the onset of the positive CIE and therefore represent a potential organic carbon sink for the proposed Weissert OAE sensu Erba et al. (2004). In a previous study of a short segment of the Vergol section, Westermann et al. (2010) suggested deposition under anoxic conditions for the finely laminated Barrande layers based on their consistent enrichment in redox-sensitive elements. This is consistent with slightly elevated C35/(C31+C35) αβ-hopane values in the Barrande layers perhaps pointing towards more reducing conditions in the water column. Euxinic conditions were probably not reached during Barrande layer formation, as indicated by the absence of isoreniratane or isoreniratene derivates. The establishment of anoxia may explain for the slightly different results obtained from maturity parameters for the Barrande layers compared to the rest of the Vergol section, since anoxic conditions and high OM accumulation rates may preserve labile biolipids (Didyk et al., 1978). The Barrande layers are characterized by relatively low CaCO3 contents (avg. 47.0  %) and enriched TOC contents (avg. 3.3 %), whereas their δ13Ccarb signature does not significantly deviate from the background values. In terms of hydrocarbon composition, all four layers show a similar pattern, but differ markedly from the rest of the succession. The laminated layers show elevated amounts of steranes over hopanes indicating a higher contribution of eukaryotic algae in comparison to prokaryotic organisms. Within the prokaryotic community, the importance of cyanobacteria is elevated. This observation is in line with environmental conditions postulated for other mid-Cretaceous black shale events, e.g. OAE1a and OAE2, where cyanobacteria occur in high abundances and have been invoked as indicator of altered nitrogen and iron cycling during anoxic episodes (Kuypers et al., 2004; Sepúlveda et al., 2009). In the Barrande layers the contribution of dinoflagellates to the eukaryotic community is reduced according to the

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ratio of dinosterane to regular steranes, pointing to unfavourable conditions for this group of organisms.The observed biomarker pattern favours a scenario with increased eukaryotic algal and cyanobacterial contribution and limited dinoflagellate productivity for the four organic-rich Barrande layers. Gréselle et al. (2011), in contrast, propose a strongly reduced productivity for carbonate producers for the Barrande layers based on a decrease in the nannofossil flux. Gréselle and Pittet (2010), and Gréselle et al. (2011) suggest a sea-level lowstand based on sequence stratigraphy, which may have favoured water column stratification and the establishment of dys- to anoxic conditions in the deep parts of the basin. Similarly, a link between enhanced primary productivity and the formation of Barrande layers 1-4 has been called into question by Reboulet et al. (2003). Data composed by Reboulet et al., (2003) on ammonites and calcareous nannofossils for the Barrande layers do point to a combination of seawater stratification and oxygen depletion at the sea bottom for the formation of the first three Barrande layers (1-3) while layer four deviates from this pattern and shows enhanced trophic levels. In this study, no distinct difference between the four layers can be observed with respect to the biomarker patterns.

3.5.5 Implications for paleoenvironmental changes in the course of the CIEIn order to assess the interplay between carbon cycle and changes in biota and paleoenvironment during the Valanginian, the stratigraphic evolution of the CIE is compared with the organic-geochemical record obtained from the Vergol/Morenas section and existing information from the Vocontian Basin and other localities.

Segment I: In the Vergol/Morenas record, evidence for short-termed dys- to anoxic conditions is restricted to the deposition of the Barrande layers. In general, biomarker and TOC data of segment I point to stable marine conditions, a well-oxygenated water column and elevated carbonate production. For the northwestern Tethys Ocean, this pre-excursion phase is known to be accompanied by relatively low sea-level, moderate sea-surface temperatures and medium fertility of surface ocean waters, strong carbonate platform growth, and reduced siliciclastic input (Weissert et al., 1998; Duchamp-Alphonse et al., 2007; McArthur et al., 2007; Gréselle et al., 2011; Föllmi et al., 2006). A gradual increase in clay influx and the beginning decline in nannoconids point to subsequent paleoenvironmental changes following segment I (Erba and Tremolada, 2004; Barbarin et al., 2012).

Segment II: During the build-up phase of the CIE, starting right after the negative δ13C peak, source-specific biomarkers show only minor fluctuations, indicating rather stable marine oxic conditions for the Vocontian Basin. A decreasing trend in the C35 homohopane index may point to enhanced mixing and reventilation of bottom waters in the aftermath of the Barrande layer episode. The increase in the dinosterane/regular sterane index is interpreted to reflect a proliferation of this group of algae. However, the consistently low TOC values and the absence of a clear biomarker signature for anoxia do not support a scenario of enhanced marine OM sequestration during this initial phase of the CIE. This pattern of comparatively stable paleoenvironmental conditions is in contrast to existing studies from the northwestern Tethys, where the onset of the CIE has been attributed to coincide with severe changes for the marine biota. The occurrence of high-fertility nannofossil taxa, combined with a continuous nannoconid decline (Erba and Tremolada, 2004; Duchamp-Alphonse et al., 2007; Barbarin et al., 2012), drastically reduced carbonate platform production associated with a shift from autotrophic to heterotrophic producers (Weissert et al., 1998; Föllmi et al., 2006; Gréselle et al., 2011), as well as high phosphorous accumulation rates (van de Schootbrugge et al., 2003), and accumulation of Mn and Fe (Kuhn et al., 2005) were used to

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indicate enhanced productivity in surface waters, probably due to increased terrestrial nutrient influx into the northwestern Tethys Ocean. Such a scenario is supported by enhanced kaolinite contents, observed in the Vocontian Basin for the very same interval (Fesneau, 2008). Paleotemperature estimates, based on Mg/Ca and δ18O data from the Vocontian Basin, point to high sea-surface temperatures during the build-up phase of the CIE (McArthur et al., 2007). According to Gréselle et al. (2011), the widespread decrease in carbonate contents observed in the Vocontian Basin was caused by a gradual decline in pelagic carbonate production and corresponds to the “biocalcification crisis” sensu Weissert and Erba (2004). This “biocalcification crisis” was postulated to reflect enhanced atmospheric CO2, temperature, and changes in surface water chemistry (Weissert and Erba, 2004). During the build-up phase of the CIE (segment II), changes in marine paleoenvironmental conditions were not limited to the Tethys Ocean. High-fertility taxa were also reported from the Pacific (Erba et al., 2004) and the Western Atlantic oceans (Bornemann and Mutterlose, 2008). Enhanced phosphorous, Mn and Fe accumulations were also reported from the Carpathian seaway (Krobichi and Wierzbowski, 1996; Kuhn et al., 2005).

Segment III: In the Vergol/Morenas record, the early plateau-phase of the CIE was shown to be accompanied by a lithologic change towards a more clay-rich sedimentation, clearly visible in the absence of calcareous interbeds and a further drop in the carbonate content. In contrast, TOC contents remain low and provide no indication for enhanced OM accumulation rates. The early plateau-phase is accompanied by a moderate increase in contributions of eukaryotic algal markers (steranes/hopanes), cyanobacteria, and dinosterane. This may point to enhanced surface water productivity, probably in concert with terrestrial nutrient input, due to a change towards more humid climatic conditions. This is supported by a continuous increase in kaolinite content, observed in the Vocontian Basin (Duchamp-Alphonse et al., 2007; Fesneau, 2008).During the plateau-phase of the CIE (segment III), probably initiating during the build-up phase of segment II, a period of substantial global cooling was proposed. This was based on nannofossil and belemnite evidence, as well as geochemical observations from the northwestern Tethys Ocean (Melinte and Mutterlose, 2001; Janssen and Clément, 2002; Pucéat et al., 2003; McArthur et al., 2007), the Boreal realm (Mutterlose and Kessels, 2000; Mutterlose et al., 2003; Price and Nunn, 2010), and the Western Pacific Ocean (here, the data interpretation is unclear, pointing to cooling or enhanced productivity, Brassell, 2009). Indirect evidence for the establishment of high-latitude ice caps is based on ice-rafted debris and glendonites (Kemper, 1987; Frakes and Francis, 1988; Price, 1999; Tarduno et al., 2002).

The Valanginian “Weissert event” was proposed to represent an OAE, causing a positive large-scale CIE due to enhanced marine carbon sequestration (Erba et al., 2004). The occurrence of widespread anoxia and a concomitant enhanced accumulation of OM in the marine realm, however, can not be confirmed for the Vocontian Basin. Given its marginal marine position and its shallow water depth as well as its record of mid-Cretaceous black shale formation (e.g. Bréhéret, 1997), the absence of widespread and continuous anoxic conditions in the Vocontian Basin during the Valanginian CIE is rather surprising. Short-lasting episodes of reducing conditions may have been prevalent during the time of Barrande layer deposition, which predate the CIE by about 180 kyrs (Gréselle et al., 2011). The Barrande layers itself can, however, apparently not explain the prominent positive δ13Ccarb shift since they only encompass a few centimetres in thickness. They rather reflect a change in bottom water redox conditions that took place before the initiation of the CIE. The absence of Valanginian black shales in the Vocontian Basin accompanying the positive CIE indicates that

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a widespread supra-regional or even global OAE did not occur. Evidence for sub- to anoxic conditions in Valanginian marine bottom waters correlating with the CIE exist for restricted areas in the Atlantic, Pacific, and Southern Ocean (e.g Summerhayes and Masran, 1983; Stein et al., 1986; Bersezio et al., 2002; Bralower et al., 2002; Reboulet et al., 2003; Westermann et al., 2010) but do not justify the assumption that oceanic anoxia and therewith changes in the marine realm have been the major and/or initial trigger for this globally recorded Valanginian CIE. Consequently, one should refer to this positive isotope anomaly as the “Weissert CIE”, rather than the “Weissert OAE”.In the studied succession, no severe variations in relative abundances of biomarkers of different producer groups can be observed during the CIE, especially not during the build-up phase, when changes in marine redox conditions were proposed to occur (e.g. Sprovieri et al., 2006). Fluctuations in productivity and the occurrence and abundance of marine biota, reported by other studies, in this study reflected in slightly enhanced amounts of biomarkers specific for dinoflagellates, are probably best explained by variations in nutrient availability via terrestrial input, as well as by changes in temperature and water column stratification (e.g. Bralower et al., 2002; Pestchevitskaya, 2008). The fact that source-specific biomarkers of this study show only minor fluctuations beyond the Barrande layers is therewith probably due to the fact that they serve as an indication for changes in redox conditions, which accordingly remained relatively stable in the Vocontian Basin. A change towards more humid conditions may have caused increased nutrient influx into oceans (e.g. Kuhn et al., 2005), but did not result in the establishment of bottom water anoxia, enhanced OM preservation and intensified carbon sequestration in the marine realm. Hence, the continental realm steps into the centre of attention for explaining the Valanginian carbon cycle perturbation (e.g. Weissert et al., 1998; van de Schootbrugge et al., 2000; Westermann et al., 2010). Since Littler et al. (2011) questioned the occurrence of a global Valanginian cold-snap during the positive carbon isotope shift, changes affecting continental environments in the form of moisture changes have to be considered as more important. The CIE may have been caused by carbon storage on continents, connected to this enhanced moisture availability. Changes in moisture levels associated with at least regionally more humid climates would have fostered the formation of marshes and swamps. This, and subsequent coal deposition would have drawn down the amount of 12C in the active carbon cycle (e.g. Westermann et al., 2010 and references therein). Indeed, there is evidence for enhanced Valanginian coastal carbon storage in the form of terrestrial coal deposits (Budyko et al., 1987; Ziegler et al., 1987; McCabe and Parrish, 1992; Rees McAllister et al., 2004). A further factor may have been storage of 13C depleted Corg under high sea-level by the formation of epicontinental seas coinciding with the initiation of the positive excursion (Gréselle et al., 2011). Further research on continental deposits and humidity proxies will help to enlighten this aspect.

3.6 ConclusionsThe geochemical results based on analyses of marine OM of the Vocontian Basin do not provide evidence for the occurrence of a Valanginian OAE associated with the positive CIE. There is no indication for anoxic or euxinic conditions in this basin during the onset and plateau-phase of the Valanginian CIE. TOC values are comparatively low and do not point to enhanced preservation of OM. Productivity may have been temporarily enhanced, caused by enhanced terrestrial nutrient influx, but did not cause dys- to anoxic conditions in the Vocontian Basin. The OM in the analyzed sedimentary rocks is predominantly of autochthonous origin with a high contribution from phytoplanktonic sources. Biomarker evidence points to less oxygenated bottom waters and enhanced productivity during the formation

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of the Barrande layers. Significant paleoenvironmental changes, associated with the formation of these cm-thick organic-rich layers, are probably related to changes in nutrient input and sea-level. As already stated by Westermann et al. (2010), carbon burial associated with the formation of these layers can most probably not account for the subsequent CIE. These findings may point to the restriction of anoxia to other oceanic basins or to paleoenvironmental changes on continents rather than in marine settings as major triggers for this positive Valanginian CIE.

AcknowledgmentsGreat thanks are due to Georg Scheeder, Monika Weiß and Annegret Tietjen (BGR, Hannover, Germany) for their help with biomarker analysis, and Rock-Eval and TOC measurements. Furthermore, thanks are due to Beate Gehnen and Ulrike Schulte (Ruhr-University Bochum, Germany) for their instructions in stable isotope measurements. Financial support from the DFG project HE4467/2-1 is gratefully acknowledged.

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Mutterlose, J., Brumsack, H., Flögel, S., Hay, W., Klein, C., Langrock, U., Lipinski, M., Ricken, W., Söding, E., Stein, R., Swientek., O., 2003. The Greenland-Norwegian Seaway: A key area for understanding Late Jurassic to Early Cretaceous palaeoenvironments. Paleoceanography 18: PA 000625.

Nunn, E.V., Price, G.D., Gröcke, D.R., Baraboshkin, E.Y., Leng, M.J., Hart, M.B., 2010. The Valanginian positive carbon isotope event in Arctic Russia: Evidence from terrestrial and marine isotope records and implications for global carbon cycling. Cretaceous Research 31, 577-592

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Pereira, R., de Souza Carvalho, I., Simoneit, B.R.T., de Almeida Azevedo, D., 2009. Molecular composition and chemosystematic aspects of Cretaceous amber from the Amazonas, Araripe and Recôncavo basins, Brazil. Organic Geochemistry 40, 863-875.

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Wortmann, U., Weissert, H., 2000. Tying platform drowning to perturbations of the global carbon cycle with a δ13Corg-curve from the Valanginian of DSDP Site 416. Terra Nova 12, 289-294.

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4. Reconstructing Valanginian (Early Cretaceous) mid-latitude vegetation and climate dynamics based on spore-pollen assemblages*

AbstractChanges in terrestrial vegetation patterns during the Valanginian (Early Cretaceous) and their link to major climatic and environmental alterations are poorly studied. In this study, the spatial and temporal changes in plant community structure are reconstructed based on spore-pollen records from two mid-latitude sites located in the Mid-Polish Trough (MPT, central Poland), and the Vocontian Basin (VB, southeast France). Stratigraphic control is provided by δ13Ccarb chemostratigraphy and calcareous nannofossil biostratigraphy. Reconstruction of hinterland vegetation is based on palynological investigations of 83 samples from hemipelagic (VB) and marginal marine (MPT) sediments rich in terrestrial palynomorphs. A total of 45 sporomorph taxa have been identified on genera level (30 spores, 15 pollen). Vegetation around the MPT was dominated by araucarian/cupressacean conifers while that enframing the VB was dominated by drought-resistant cheirolepidiacean conifers. At both sites the understorey and/or vegetation of open areas was dominated by pteridophytes. An early Valanginian gradual trend towards humid conditions at the MPT, well expressed by a distinct increase in the spore-pollen ratio, culminates in a short-termed spore-maximum, stratigraphically located at the early/late Valanginian boundary. It is characterized by low conifer abundances and high abundances of the fern spore taxa Cyathidites, Leiotriletes and Gleicheniidites, accompanied by enhanced abundances of the pteridosperm pollen Vitreisporites pallidus, which parent plants are assumed to be indicative of swamp habitats. The spore-maximum is coeval to a similar peak observed in the VB, characterized by essentially the same taxa. Here, the spore-maximum is preceded by a protracted phase of arid conditions, characterized by low spore abundances and exceptionally high numbers of the cheirolepediacean conifer pollen Classopollis. An occurrence of a cooling episode in the Late Valanginian may be inferred from an interval marked by high bisaccate and low spore abundances, most probably indicating comparatively cold and arid conditions. Changes in moisture, identified as the key climatic factor determining trends and turnovers in vegetation, were probably controlled by a monsoonal circulation. The supra-regional humid phase, expressed by the coeval spore maxima, was probably induced by an intensified monsoonal climate. The temporal influence of a northern hemisphere arid belt at the VB, under the influence of subtropical high-pressure belt, may have caused the temporal drying not affecting the MPT site, located further north.

KeywordsEarly Cretaceous, Valanginian palynology, Valanginian vegetation, δ13Corg chemostratigraphy, calcarous nannofossil biostratigraphy, paleoclimate* This chapter is based on a cooperation with:Hochuli, P.A. (Zurich, Switzerland), Pauly, S. (Bochum, Germany), Morales, C., Adatte, T., Föllmi, K. (all Lausanne, Switzerland), Ploch, I. (Warsaw, Poland).

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4.1 IntroductionThe Valanginian (Early Cretaceous) was characterized by several alterations of the ocean-atmosphere system. These primarily include prominent perturbations of the carbon cycle that were globally recorded by a positive carbon isotope excursion (CIE; e.g. Cotillon and Rio, 1984; Lini et al., 1992; Hennig et al., 1999; Wortmann and Weissert, 2000; Weissert and Erba, 2004; Gröcke et al., 2005; Föllmi et al., 2006; Nunn et al., 2010). This CIE was accompanied by fluctuations in the atmospheric pCO2 concentration, probably connected to volcanic activity (e.g. Lini et al., 1992; Price and Mutterlose, 2004; Weissert and Erba, 2004). Furthermore, changes in terrigenous input to ocean margins (Weissert, 1989; Föllmi et al., 1994; Weissert et al., 1998; Weissert and Erba, 2004; Föllmi et al., 2006) and a pronounced phase of climatic cooling were proposed based on geochemical analyses (Weissert and Lini, 1991; Podlaha et al., 1998; Pucéat et al., 2003; McArthur et al., 2007; Brassell, 2009; Price and Nunn, 2010) as well as changes in calcareous nannofossil assemblages (Mutterlose and Kessels, 2000; Melinte and Mutterlose, 2001; Mutterlose et al., 2003). Indirect evidence for the establishment of high-latitude ice caps in the form of ice-rafted debris and glendonites (Kemper, 1987; Frakes and Francis, 1988; Price, 1999; Gréselle and Pittet, 2010, and references therein) has been discussed. So far, most studies have concentrated on changes in the marine biosphere and on turnovers in numerous marine organisms (e.g. Mutterlose and Kessels, 2000; Melinte and Mutterlose, 2001; Mutterlose et al., 2003), including a “biocalcification crisis” in pelagic settings associated with a widespread river-influenced carbonate platform demise (e.g. Weissert and Erba, 2004) and a nannoconid decline (e.g. Barbarin et al. 2012). In contrast, continental environments have received less attention. Little is known about the interdependence of terrestrial plant communities and the extreme environmental and climatic perturbations associated with the Valanginian CIE. Other than for stratigraphic purposes information on Valanginian vegetation changes is scarce.Here, we present two palynological records covering Valanginian strata. The first one is based on marginal marine sediments from the Mid-Polish Trough (MPT) of central Poland. It was located within the Carpathian Seaway that connected the Boreal Realm and the Tethys. The second record is derived from hemipelagic deposits cropping out in the Vocontian Basin (VB) of southeast France, which was located further south and part of the northwestern Tethys. From a paleophytogeographic perspective, both sites were situated within the humid, warm-temperate European region, rich in ferns (Ziegler et al., 1987). In order to assess changes in vegetation structure and associated paleoenvironmental parameters, encountered sporomorphs were determined taxonomically and their stratigraphic ranges are presented for both sites. To compare the palynological records from the two localities, stratigraphic correlation has been carried out based on calcarous nannofossil biostratigraphy and trends in carbon isotopes (δ13Ccarb chemostratigraphy). Variations in the abundances of the spore-pollen allowed the definition of palynological assemblage zones for both sites. Based on botanical affinities to modern plants, ecological preferences of the fossil sporomorph taxa are deduced, providing information on changes in physical environmental parameters (e.g. moisture). Due to their paleogeographic position under the influence of the subtropical high-pressure belt, the vegetation from both sites serves as a sensitive recorder of changes in climatic patterns. The absence of large polar ice-caps in combination with the plate-tectonic configuration during the Valanginian is favourable for the establishment of a monsoonal climate on the southern margin of paleo-Europe. This potential monsoonal influence, expressed in vegetation changes, as well as the effect of the proposed Valanginian “cold snap” have been investigated.

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Reconstructing Valanginian (Early Cretaceous) mid-latitude vegetation and climate dynamics based on spore-pollen assemblages

4.2 Geological setting and stratigraphy of studied sections4.2.1 The Mid-Polish Trough (MPT) recordMaterial from a core (PIG1) which was drilled close to the village of Wąwał (E 19°15’0’’, N 52°25’0”) by the Polish Institute of Geology, Warsaw, was studied (Fig. 4.1A and 4.2). Wąwał is located about 4 km southeast to the city of Tomaszów Mazowiecki, and about 115 km southwest to Warsaw. The highly condensed succession has a thickness of 18 m. Its lithology is composed of clay and claystone with layers of more sandy intervals, shells and sideritic, calcareous and phosphatic nodules. The basal part (80 cm) consists of limestone, presumably of Berriasian age (Kutek et al., 1989). The core material lacks ammonites of biostratigraphic significance; the age model is based on calcareous nannofossil biostratigraphy and δ13Corg-δ13Ccarb chemostratigraphy. Additional stratigraphic control is supplied by lithostratigraphic correlation of the core with a near-by clay pit well dated by ammonites (Kutek et al., 1989; Kaim, 2001). Accordingly, the PIG1 core from Wąwał covers deposits of Berriasian/early Valanginian to late Valanginian age (Fig. 4.2). During the Cretaceous, the Carpathian Seaway connected the Tethys and the Boreal Realm via the Lower Saxony Basin and the North Sea Basin (Kutek et al., 1989; Mutterlose, 1992; Kaim, 2001). The study site is located in the MPT in the centre of the Carpathian Seaway at a paleolatitude of ~35-40°N (e.g. Blakey, 2010).

Fig. 4.1. Location of sites, A) Polish site (Wąwał) within the Mid-Polish Trough (MPT; modified after Daldez, 2003), B) French site (La Charce, Vergol, Morenas) within the Vocontian Basin (VB; modified after Gréselle, 2007). Black asterisks refer to the studied sections, for other signatures and colours see legends within maps.

4.2.2 The Vocontian Basin (VB) recordThe material studied in the VB is derived from three outcrops (La Charce, Vergol, Morenas) located in the Drome department (Fig. 4.1B and 4.2). The Vergol section (E 5°25’9’’, N 44°12’12”) is located between the villages Montbrun-les-Bains and Vergol. Logged and sampled strata (thickness of ~57 m) cover sedimentary rocks from the lower Valanginian (Busnardoites campylotoxus Zone) to the upper Valanginian (Saynoceras verrucosum Subzone). To compensate for a hiatus of slumping, the interval between 25-30 m at the Vergol section has been covered by samples collected at the nearby, lithostratigraphically well correlated Morenas section (E 5°25’23’’, N 44°13’52’’). The La Charce section (E 5°26’23’’, N 44°28’13’’) is well accessible on the hill slopes west of the village of La Charce. Logged and sampled strata (thickness of ~113 m) cover the stratigraphic interval from the lower part of the Upper Valanginian (Karakaschiceras pronecostatum Subzone) to the Lower Hauterivian (Acanthodiscus radiatus Zone; Gréselle and Pittet, 2010, and references therein). During the Early Cretaceous the VB was located at a paleolatitude of ~30°N in a marginal marine position

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Fig. 4.2. Biostratigraphic (calcarous nannofossils) and chemostratigraphic (δ13Ccarb) correlation of the VB (left, Kujau et al., 2012 and this study) and MPT (right, Morales et al., in prep.). Furthermore, thickness, lithology (for legend see Figs. 4.3, 4.4), composition of VB composite succession (Vergol, Morenas, La Charce), and position of samples (only MPT, for VB see Fig. 4.4) are shown. Biostratigraphic correlation in based on Tethyan (NK, NC after Gréselle, 2011) and Boreal (BC, this study) nannofossil zones, following Bown (1998). FO=first occurrence, LO=last occurrence. Position of palynologic samples is marked as elongated horiontal lines (only MPT), short lines mark additional samples for δ13C. Orientation of correlation is following the lower to upper Valanginian boundary (light dashed line).

of the Western Ligurian Tethys Ocean, open to the east. To the north the VB was separated from the Boreal Realm by the archipelago of mid-European continents (Ziegler, 1982; Hay et al., 1999; Masse, 1993; Blakey, 2010). The central part of the VB is characterized by hemipelagic deposits composed of autochthonous carbonates, allochthonous carbonate fine fraction exported from three surrounding platforms (Fig. 4.1A; Reboulet et al., 2003; Gréselle and Pittet, 2010), and terrigenous material, mostly eroded from the Massif Central area (Bréhéret, 1994; Fesneau et al., 2009). The studied deposits are rich in marine invertebrate fossils and consist of well exposed orbitally controlled marl-limestone alternations, stacked in bundles (Gréselle and Pittet, 2010). The composite section has a well-established age model based on ammonites, calcareous nannofossils, and cyclostratigraphy; it covers sediments from the lower Valanginian to lower Hauterivian (Gréselle et al., 2011 and references therein). Age control is also provided by δ13Ccarb chemostratigraphy that has been compared to existing, well dated Tethyan records (Kujau et al., 2012). The composite succession has a total thickness of ~175 m and covers the interval of the lower Valanginian B. campylotoxus to lower Hauterivian A. radiatus ammonite zones (Gréselle and Pittet, 2010, Fig. 4.2).

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Reconstructing Valanginian (Early Cretaceous) mid-latitude vegetation and climate dynamics based on spore-pollen assemblages

4.3 Material and Methods4.3.1 Terrestrial palynologyA total of 83 palynological samples (46 from MPT and 37 from the VB) was studied quantitatively and qualitatively using transmitted light microscopy on an OLYMPUS BX-51. For samples from the MPT locality contamination by modern plant material can be excluded due to sampling of fresh drill core material. In order to minimize contamination of the VB based samples, the uppermost 15-20 cm of sediment were removed before sampling. Preparation of strew slides was done at the Geological Survey NRW, Germany, following standard procedures (Traverse, 2007). For plotting the results the software Tilia® was used. For each sample spore-pollen counts up to a total of 300 miospores (+/- 5) were done. This was done on one to two slides depending on the richness in miospores for each sample.

4.3.2 Nannofossil biostratigraphyFor the investigation of calcareous nannofossils from the MPT a total of 135 of samples have been studied, the preparation of standard smear slides followed Bown (1998). Nannofossil information for the VB, using the Tethyan (NC) zonation scheme, was adopted from Gréselle et al. (2011). Identification of calcareous nannofossils for the MPT was carried out using an OLYMPUS BH-2 light microscope with cross-polarized light at a magnification of × 1250. At least two traverses of each smear slide were studied. Identification of taxa follows the taxonomic concepts of Perch-Nielsen (1985) and Bown et al. (1998). For this section we used the BC (Boreal calcareous nannofossil) zonation scheme of Bown et al. (1998).

4.3.3 Stable carbon isotopesThe stable carbon isotope record (δ13Ccarb) for the MPT is adopted from Morales et al. (in prep.), performed on the same core (PIG 1). Measurements of δ13Ccarb for the VB have been carried out on powdered bulk sample material (~0.6 mg) on 295 samples (the first 97 are already presented in Kujau et al., 2012). A Thermo Fisher Scientific Gasbench II carbonate device connected to a Thermo Fisher Scientific Delta S isotope ratio mass spectrometer, available at the Ruhr-University Bochum, Germany, has been used for stable isotope analyses. The gas bench uses 100 % phosphoric acid at 70°C to release CO2 of the calcite from the sample material one hour prior to the start of the measurement. Repeated analyses of certified carbonate standards (CO- 1, CO-8, NBS-19) show an external reproducibility ± 0.1 ‰ for δ18O and ± 0.06 ‰ for δ13Ccarb. Values are expressed in conventional delta notation relative to the VPDB international standard, in per mil (‰).Duplicate measurements (31) show that the measured values are representative and indicate that the samples are homogenous (max. dev. δ18O ± 0.17 and δ13C ± 0.12).

4.4 Results4.4.1 Terrestrial palynologyA total of 45 sporomorph taxa, have been identified on the generic level (30 spores and 15 pollen) throughout the investigated samples. Botanical affinities and ecologic preferences of the parent plant are provided in Tables 4.1 and 4.2. Palynological results from the two areas show strong similarities regarding diversity of the taxa. Only four spores found in the MPT section (Aratrisporites, Staplinisporites, Clavatisporites, Todisporites) are absent from the VB. Cyathidites represents the most abundant spore at both localities. All identified pollen grains occur at both sites. The most abundant pollen for the VB sections is Classopollis, for the MPT section it is Callialasporites. Angiosperm pollen are absent from all samples.

58

Chapter 4

Table 4.1. Recorded sporomorph taxa from the composite section of the Vocontian Basin and the Wawal core and their probable botanical affinities. For describing author see Table 4.2.

Taxa Botanical affinity

Pollen Araucariacites1 Araucariaceae1

Callialasporites2 Araucariaceae1,2

Inaperturopollenites3(inaperturate) Cupressaceae (formerly Taxodiaceae, e.g. Birkenmajer et al., 2010)

/Araucarian1/3

Perinopollenites4 (inaperturate) Cupressaceae (formerly Taxodiaceae) /Araucarian1/3

Cerebropollenites5 Cupressaceae (formerly Taxodiaceae)4

Cycadopites6(monosulcate) Cycadales, Bennettitales, Ginkgoales3

Eucommiidites7(monosulcate) Cycadales1, 5

Ephedripites8(monosulcate) Ephedrales/Ephedraceae6

Classopollis9(=Corollina,Circulina) Cheirolepideaceae7, 5

Exesipollenites10 Bennettitales8, 9

Alisporitesgroup11(bisaccate) Conifers and Pteridospermae/ Corystospermales (Pteridosperms)2, 10/1, 5

Vitreisporitespallidus15(bisaccate) Caytoniales (Pteridosperms)1Parvisaccites12(bisaccate) Podocarpaceae11

Podocarpidites14(bisaccate) Podocarpaceae12

Pinuspollenites13(bisaccate) Coniferopsida13

Spores Cyathidites16(Deltoidospora) Dicksoniaceae, Cyatheaceae, Dipteridaceae, Filicopsida, Matoniaceae3, 4, 5, 14

Trilobosporites19 Dicksoniaceae15, 3

Gleicheniidites20 Gleicheniaceae16

Neoraisrtrickia21 Lycopodiales3

Clavatisporites22 Filicopsida17

Biretisporites23 Filicopsida17

Leiotriletes24(Deltoidospora) Filicopsida17

Osmundacidites25 Filicopsida or Osmundaceae (Osmunda-type)7, 3

Gemmatriletes26 Filicopsida17

Verrucosisporites Filicopsida17

Klukisporites27 (Filicopsida or) Schizaeaceae18

Todisporites28 Osmundaceae (Todites-type)7, 3

Retriletes29 Lycopodiaceae (Lycopodium type) 16, 3

Camarozonosporites30 Lycopodiaceae19

Lycopodiumsporites31 Lycopsida: Lycopodiaceae (Lycopodium) 17

Leptolepidites32 Lycopodiales19, 3

Lycopodiacidites33 Lycopodiales19

Staplinisporites34 Bryophyta and Lycopodiaceae16, 2, 3

Aratrisporites35 Lycopsida17

Densoisporites36 Lycopodiaceae or Pleuromeiaceae (Selaginella)/Pleuromeiaceae20, 21,/22, 3

Cicatricosisporites37 Schizaeaceae, Anemia/Mohria type23, 24, 3

Ischyosporites38 Schizaeaceae, Lygodium-type23, 24, 3

Stereisporites39 Bryophyta19

Aequitriradites40 Bryophyta17

Foraminisporis41 Bryophyta, namely hepatics/Anthocerotaceae25/19

Concavissimisporites42(Deltoidospora) Dickinsoniaceae, Cyatheaceae26, 3

Concavisporites43 Matoniaceae14

Foveosporites44 Lycopodiales/Selaginellales27/28

Echinatisporis45 Lycopods/Selaginella19

Microreticulatisporis ?

Zygnemataceae Ovoidites(Scafati et al., 2009; freshwater algae)Prasinophyceae Leiosphaeridia (Scafati et al., 2009; marine to freshwater algae)Acritarch Leiofusa(van de Schootbrugge et al., 2007; “disaster species”)

Selected references for botanical affinities: 1Van Konijnenburg-Van Cittert, 1971; 2Boulter and Windle, 1993; 3Abbink et al., 2004; 4van der Burgh and Van Konijnenburg-Van Cittert, 1989; 5Balme, 1995; 6Azéma and Boltenhagen, 1974; 7Van Konijnenburg-Van Cittert, 1987; 8Harris, 1969, 91974; 10Traverse, 2007;

11Schrank, 2010; 12Raine et al., 2006; 13Larsson, 2009 (Kemp, 1970); 14Van Konijnenburg-Van Cittert, 1993; 15Venkatachala et al., 1969; 16Potonié, 1967; 17Raine et al., 2008; 18Couper, 1958; 19Filatoff, 1975; 20Knox, 1950; 21Lundblad, 1950; 22Raine et al., 1988; 23Van Konijnenburg-Van Cittert, 1981, 241991; 25Dettmann, 1963;

26Potonié, 1970; 27Döring, 1965; 28Courtinat, 1989.

Colours of thin walled miospores for both sites range from light yellowish to orange (low thermal alteration index = 2+; e.g. Zobaa, 2006), with lighter colours for the MPT. Preservation is consistent throughout the samples of each site, no sign of carbonization is observed. Reworking and contamination with miospores of younger or older strata can thereby be excluded. Due to pre- or post-depositional stress preservation of the spore and pollen grains varies from excellently preserved to highly fragmented specimens. A high degree of fragmentation hampers the determination especially between inaperturate palynomorphs like Araucariacites, Inaperturopollenites and simple thin-walled leiospheres (e.g. Filatoff, 1975; Vajda, 2001; Schrank, 2010). Furthermore, pollen of Araucariacites and Callialasporites are known to show high polymorphism, and between the two genera transitional forms seem to exist (Balme, 1995; Schrank, 2010). In some cases identification of bisaccate pollen and (mainly deltoid) spores was impossible due to poor preservation, fragmentation, or a cover of amorphous organic matter or pyrite, or due to their orientation in the slide. Consequently, these forms were labelled and grouped as undifferentiated bisaccates and undifferentiated

59

Reconstructing Valanginian (Early Cretaceous) mid-latitude vegetation and climate dynamics based on spore-pollen assemblages

Tabl

e 4.

2. S

pore

s and

pol

len

grou

ped

acco

rdin

g to

fam

ily/g

ener

a le

vel w

ith e

xem

plar

y re

fere

nce

and

auth

or o

f firs

t des

crip

tion,

pla

te a

nd fi

gure

, and

affi

liatio

n to

cla

ssis

with

shor

t des

crip

tion

of e

colo

gica

l pre

fere

nce

of p

aren

t pla

nt fo

r im

porta

nt ta

xa, i

f kno

wn.

Taxa

A

utho

r and

refe

renc

e

Plat

e an

d Fi

gure

C

lass

is

Ec

olog

ical

pre

fere

nce

of p

aren

t pla

nts

bisa

ccat

e po

llen:

Vitreisporitespallidus

(R

eiss

inge

r) N

illso

n, 1

958,

in B

alte

s, 19

67

Pl

ate

I, Fi

gure

11

Pter

idos

perm

s/G

inkg

os(B

alm

e, 1

995)

D

elta

ic e

nviro

nmen

ts, fl

ood

plai

n to

bac

ksw

amps

, ind

icat

ive

of

hu

mid

env

ironm

ents

(Abb

ink

et a

l., 2

004

and

refe

renc

es th

erei

n).

Alisporites

Cou

per,

1958

, in

Hoc

huli

and

Kel

ts, 1

980

Pl

ate

I, Fi

gure

1, 2

Pt

erid

ospe

rms/

Gin

kgos

(or c

onife

rs)

(B

atte

n, 1

975;

Bal

me,

199

5; S

chra

nk,

2010

)Parvisaccites

C

oupe

r, 19

58, i

n B

alm

e, 1

995

Plat

e I,

Figu

re 3

, 4

Con

ifers

(Bal

me,

199

5)

Mou

ntai

n ha

bita

ts, c

oole

r con

ditio

ns, s

omet

imes

asc

ribed

as t

ypic

al

flora

l ele

men

ts o

f bor

eal r

egio

ns w

ith b

road

tem

pera

ture

rang

e

but p

refe

renc

e fo

r rel

ativ

ely

hum

id c

ondi

tions

(Sch

rank

,

2010

; Hoc

huli

and

Kel

ts, 1

980)

.Podocarpidites

(C

ooks

on) C

oupe

r, 19

53, i

n B

alm

e, 1

995

Pl

ate

I, Fi

gure

5

Con

ifers

(Bal

me,

199

5)

Pref

eren

ces l

ike Parvisaccites.

Pinuspollenites

Po

toni

é, 1

985

in B

alm

e, 1

995

Plat

e I,

Figu

re 6

C

onife

rs

Ty

pica

l flor

al e

lem

ent o

f bor

eal r

egio

ns, g

row

ing

unde

r rel

ativ

ely

hu

mid

con

ditio

ns w

ith a

bro

ad ra

nge

of te

mpe

ratu

re to

lera

nce

(H

ochu

li an

d K

elts

, 198

0).

non-

sacc

ate

polle

n:Araucariacites

(C

ooks

on) C

oupe

r, 19

53, i

n B

alm

e, 1

995

Pl

ate

I, Fi

gure

7, 8

C

onife

rs (B

alm

e, 1

995)

W

ide

rang

ing,

rain

fore

sts t

o co

ol te

mpe

rate

fore

sts,

rece

nt

Araucaria

mos

tly o

ccur

ring

in m

oist

tem

pera

te e

nviro

nmen

ts o

n

tropi

cal m

ount

ain

flank

s. H

igh

abun

danc

es o

f Araucaria

indi

cativ

e

of re

lativ

ely

arid

con

ditio

ns, l

ow a

bund

ance

s con

nect

ed w

ith h

igh

di

vers

ity in

dica

tive

of h

umid

con

ditio

ns (S

chra

nk, 2

010)

.Callialasporites

Su

kh D

ev, 1

961,

in B

alm

e, 1

995

Plat

e I,

Figu

re 9

, 10

Con

ifers

/ (B

alm

e, 1

995)

Pref

eren

ces l

ike Araucariacites.

Inaperturopollenites

B

alm

e, 1

957,

in C

onw

ay, 1

996

Plat

e I,

Figu

re 1

2, 1

3 C

onife

rs (B

alm

e, 1

995)

Metasequoia

type

, ind

icat

ive

of m

ore

hum

id c

ondi

tions

(Birk

enm

ajer

et a

l., 2

010)

.Perinopollenites

C

oupe

r, 19

58, i

n B

alm

e, 1

995

Plat

e I,

Figu

re 1

4 C

onife

rs (B

alm

e, 1

995;

Vaj

da, 2

001)

Prob

ably

war

m a

nd a

rid c

oast

al e

nviro

nmen

ts (B

alm

e, 1

995;

Vaj

da,

2001

). Cerebropollenites

(C

oupe

r) N

ilsso

n, 1

958,

in H

erng

reen

, 197

1

Plat

e I,

Figu

re 1

5,16

C

onife

rs (V

ajda

, 200

1)

Supp

osed

to b

e re

late

d to

rece

nt tr

ee Tsuga

(Vaj

da, 2

001

and

refe

renc

es th

erei

n), t

here

with

may

hav

e gr

own

in h

umid

and

coo

l

ha

bita

ts.

Classopollis(=Corrollina,

Mal

javk

ina,

194

9, in

Bal

me,

199

5

Pl

ate

I, Fi

gure

17,

18

Con

ifers

(Bal

me,

199

5)

Cos

mop

olita

n, h

ighl

y di

vers

e fa

mily

, ada

pted

to w

ide

rang

e of

Circulina;

Yi e

t al.,

199

3)

ha

bita

ts. M

ost c

omm

only

ass

igne

d to

sem

i-arid

to a

rid lo

wla

nd

shor

elin

e ha

bita

ts, a

rid to

seas

onal

ly a

rid c

ondi

tions

, and

pro

babl

y

assi

gned

to h

igh

tem

pera

ture

s. So

me

may

hav

e be

en a

dapt

ed to

co

asta

l sal

ine

soil

cond

ition

s, pr

obab

ly a

ssig

ned

to h

igh

tem

pera

ture

s (Va

khra

mee

v, 1

981,

Doy

le e

t al.,

198

2; B

atte

n, 1

984;

Sc

hran

k, 2

010)

.Exesipollenites

B

alm

e, 1

957,

in B

alm

e, 1

995

Plat

e I,

Figu

re 2

2 C

onife

rs (o

r cyc

ads)

(Bal

me,

199

5)

A

rid c

ondi

tions

(Sch

rank

, 201

0).

Cycadopites

W

ilson

and

Web

ster

, 194

6, in

Bal

me,

199

5

Plat

e I,

Figu

re 1

9, 2

0 C

ycad

s (B

alm

e, 1

995;

Abb

ink

et a

l., 2

004)

A

rid, w

arm

con

ditio

ns, m

oder

n flo

ra m

ainl

y tro

pica

l to

subt

ropi

cal,

prob

ably

not

in w

etla

nds (

Yi,

1993

;Bal

me,

199

5; A

bbin

k et

al.,

20

04; C

oiffa

rd e

t al.,

200

7; V

akhr

amee

v, 2

010)

. Eucommiidites

(E

rdtm

an) P

oton

ié, 1

958,

in B

alm

e, 1

995

Pl

ate

I, Fi

gure

23

Cyc

ads (

or g

neta

les)

(Bal

me,

199

5; B

atte

n

-

and

Dut

ta, 1

997)

Ephedripites

B

olkh

oviti

na e

x Po

toni

é, 1

958,

in B

alm

e, 1

995

Pl

ate

I, Fi

gure

21

Gne

tale

s (Sc

hran

k, 2

010)

May

indi

cate

war

m se

mi-a

rid to

arid

con

ditio

ns, a

lso

wid

er ra

nge.

Lo

w p

erce

ntag

es a

lso

due

to p

aren

t pla

nts b

eing

wea

k po

llen

prod

ucer

s, po

llen

is n

ot e

asily

tran

spor

ted.

Hig

h pe

rcen

tage

s

po

int t

o th

e oc

curr

ence

of p

aren

t pla

nts c

lose

to si

te o

f dep

ositi

on

(Sch

rank

, 201

0). M

oder

n pl

ants

ada

pted

to e

xtre

me

tem

pera

ture

co

nditi

ons (

hot a

nd c

old;

Yi e

t al.,

199

3).

spor

es:

Cyathidites

Cou

per,

1953

, in

Bal

me,

199

5

Pl

ate

I, Fi

gure

31

Fern

s (Sc

hran

k, 2

010)

M

oder

n re

lativ

es p

refe

rent

ially

gro

w u

nder

tro

pica

l to

sub

tropi

cal

(=D

elto

idos

pora

?)

cond

ition

s (Y

i et a

l., 1

993)

.Leiotriletes

(Loo

se) P

oton

ié a

nd K

rem

p, 1

954,

in B

alm

e, 1

995)

Pl

ate

I, Fi

gure

24,

25

Fern

s (B

alm

e, 1

995;

Sch

rank

, 201

0)

Sh

orel

ine

clos

e ve

geta

tion

and

herb

aceo

us sw

amps

(Wild

e, 1

989)

.(=

Del

toid

ospo

ra?)

Concavisporites

(C

oupe

r, 19

53) K

rutz

sch,

195

9 in

Rai

ne e

t al.,

200

8)

Plat

e I,

Figu

re 3

5, 3

6 Fe

rns (

Rai

ne e

t al.,

200

8)

-

(=D

elto

idos

pora

?)Concavissimisporites

(D

elco

urt a

nd S

prum

ont)

Del

cour

t, D

ettm

ann

Pl

ate

I, Fi

gure

38

Fern

s (Sc

hran

k, 2

010)

-

and

Hug

hes,

1963

, in

Her

ngre

en, 1

971

Trilobosporites

C

oupe

r, 19

58, i

n H

erng

reen

, 197

1

Pl

ate

I, Fi

gure

29

Fern

s (Sc

hran

k, 2

010)

-

Cicatricosisporites

Po

toni

é an

d G

elle

tich,

193

3, in

Bal

me,

199

5

Plat

e I,

Figu

re 2

6, 2

7 Fe

rns (

Bal

me,

199

5)

May

poi

nt to

hum

id c

ondi

tions

, gro

wn

in m

arsh

es, r

iver

eco

syst

ems

or u

nder

stor

ey in

fore

sts.

May

als

o be

ada

pted

to m

ore

arid

cond

ition

s o

f hea

thla

nds l

ike

rece

nt sc

hiza

eace

ous f

ern

Ane

mia

(S

chra

nk a

nd M

ahm

oud,

199

8; S

chra

nk, 2

010)

.

60

Chapter 4

Tabl

e 4.

2. (continuation)

Spo

res a

nd p

olle

n gr

oupe

d ac

cord

ing

to fa

mily

/gen

era

leve

l with

exe

mpl

ary

refe

renc

e an

d au

thor

of fi

rst d

escr

iptio

n, p

late

and

figu

re, a

nd a

ffilia

tion

to c

lass

is w

ith sh

ort d

escr

iptio

n of

eco

logi

cal p

refe

renc

e of

par

ent p

lant

for i

mpo

rtant

taxa

, if k

now

n.

Taxa

A

utho

r and

refe

renc

e

Plat

e an

d Fi

gure

C

lass

is

Ec

olog

ical

pre

fere

nce

of p

aren

t pla

nts

Klukisporites

C

oupe

r, 19

58, i

n H

erng

reen

, 197

1

Pl

ate

II, F

igur

e 41

Fe

rns (

Schr

ank.

, 201

0)

-Ischyosporites

B

alm

e, 1

957,

in B

alm

e, 1

995

Plat

e II

, Fig

ure

33, 3

4 Fe

rns (

Schr

ank,

201

0)

-Osmundacidites

C

oupe

r, 19

53, i

n B

alm

e, 1

995

Plat

e II

, Fig

ure

47

Fern

s (Sc

hran

k, 2

010)

-

Clavatisporites

K

edve

s and

Sim

oncs

ics,

1965

, in

Schr

ank

and

Pl

ate

II, F

igur

e 44

Fe

rns (

Rai

ne e

t al.,

200

8)

-

Mah

mou

d, 1

998

Gleicheniidites

R

oss,

1949

, in

Bal

me,

199

5

Pl

ate

II, F

igur

e 39

Fe

rns (

Bal

me,

199

5)

Opp

ortu

nist

ic, p

ione

erin

g pl

ants

, als

o gr

owin

g in

uns

tabl

e

(incl

udin

g Clavifera)

(B

olkh

oviti

na, 1

966,

in W

aksm

undz

ka, 1

981)

envi

ronm

ents

, rud

eral

stra

tege

rs (H

erng

reen

, 197

0; C

oiffa

rd, 2

007)

.Todisporites

C

oupe

r, 19

58 in

Sch

rank

et a

l., 2

012

Plat

e II

, Fig

ure

42

Fern

s (Sc

hran

k, 2

010)

M

ay p

oint

to h

umid

con

ditio

ns g

row

n in

mar

shes

, riv

er e

cosy

stem

s

or

as u

nder

stor

ey in

fore

sts.

May

als

o be

ada

pted

to m

ore

arid

cond

ition

s (Sc

hran

k., 2

010)

.Gemmatriletes

(B

renn

er) S

ingh

, 197

1, in

Zob

aa 2

006

Plat

e II

, Fig

ure

30

Fern

s (R

aine

et a

l., 2

008)

-Biretisporites

D

elco

urt a

nd S

prum

ont,

1955

, in

Plat

e II

, Fig

ure

43

Fern

s (R

aine

et a

l., 2

008)

-Ta

ugou

rdea

u-La

ntz,

198

8 Verrucosisporites

Vo

lkhe

imer

197

2, in

Rai

ne e

t al.,

200

8

Plat

e II

, Fig

ure

40

Fern

s (R

aine

et a

l., 2

008)

-Neoraisrtrickia

(C

ooks

on) P

oton

ié, 1

956,

in Z

obaa

200

6

Plat

e II

, Fig

ure

32

Lyco

pods

(Abb

ink

et a

l., 2

004)

-

Retriletes

(Coo

kson

) Dör

ing

et a

l., 1

963,

in C

onw

ay, 1

996

Pl

ate

II, F

igur

e 45

, 46

Lyco

pods

(Abb

ink

et a

l., 2

004)

-

Camarozonosporites

(H

edlu

nd, 1

966)

Nor

ris, 1

967,

in G

ranz

ow, 2

000

Pl

ate

II, F

igur

e 48

Ly

copo

ds (A

bbin

k et

al.,

200

4)

-Aratrisporites

Le

schi

k, 1

956,

in B

alm

e, 1

995

Plat

e II

, Fig

ure

28

Lyco

pods

(Bal

me,

199

5)

-

Densoisporites

W

eyla

nd a

nd K

riege

r, 19

53, i

n B

alm

e, 1

995

Pl

ate

II, F

igur

e 50

Ly

copo

ds (B

alm

e, 1

995)

Tida

lly-in

fluen

ced

ecos

yste

m (S

chra

nk, 2

010)

, doe

s not

tole

rate

ar

idity

(Hed

lund

, 196

6).

Foveosporites

B

alm

e, 1

957,

in B

alm

e, 1

995

Plat

e II

, Fig

ure

51

Lyco

pods

(Abb

ink

et a

l., 2

004)

-

Lycopodiacidites

C

oupe

r 195

3, in

Rai

ne e

t al.,

200

8

Pl

ate

II, F

igur

e 52

Ly

copo

ds (R

aine

et a

l., 2

008)

-

Echinatisporis

K

rutz

sch

1963

, in

Schr

ank

et a

l, 20

10

Pl

ate

II, F

igur

e 53

Ly

copo

ds (A

bbin

k et

al.,

200

4)

-Leptolepidites

N

orris

, 196

9, in

Zob

aa, 2

006

Plat

e II

, Fig

ure

58

Lyco

pods

(or f

erns

) (A

bbin

k et

al.,

-

2004

; Sch

rank

, 201

0)

-

Lycopodiumsporites

Si

ngh,

196

4, in

Tau

gour

deau

-Lan

tz, 1

988

Pl

ate

II, F

igur

e 54

Ly

copo

ds

(Wak

smun

dzka

, 198

1)

Fa

vour

hum

id c

ondi

tions

but

can

tole

rate

long

per

iods

of d

roug

ht

(Sch

rank

., 20

10)

Staplinisporites

(B

alm

e) P

ococ

k, 1

962,

in T

augo

urde

au-L

antz

, 198

8 Pl

ate

II, F

igur

e 56

B

ryop

hyte

s (or

lyco

pods

)

-

(Abb

ink

et a

l., 2

004)

Foraminisporis

D

örin

g, 1

973,

in S

chra

nk, 2

010

Plat

e II

, Fig

ure

55

Bry

ophy

tes (

Schr

ank,

201

0)

-Stereisporites

(W

ilson

and

Web

ster

) Det

tman

n 19

63, i

n Va

jda,

200

1 Pl

ate

II, F

igur

e 57

(S

phag

num

) bry

ophy

tes

Hab

itats

like

pea

t bog

s and

mire

s.

(Wak

smun

dzka

, 198

1)Aequitriradites

(C

ooks

on a

nd D

ettm

an) D

ettm

an 1

961,

in

Pl

ate

II, F

igur

e 59

Ly

copo

ds (m

arch

antio

phyt

a, li

verw

orts

) N

ear r

iver

site

s or a

rid p

lace

s, pr

obab

ly x

erop

hytic

(Sch

rank

and

Her

ngre

en, 1

971

(S

chra

nk a

nd M

ahm

oud,

199

8)

M

ahm

oud,

199

8).

Microreticulatisporis

K

nox,

195

0 in

Pér

ez L

oina

ze a

nd C

ésar

i, 20

04

Pl

ate

II, F

igur

e 60

?

-

61

Reconstructing Valanginian (Early Cretaceous) mid-latitude vegetation and climate dynamics based on spore-pollen assemblages

spores. A total of four samples (MOR5, VER24, VER26, VER29) has been excluded from spore-pollen analysis due to high amounts of amorphous organic matter, which did not allow for miospore identification. Stratigraphic occorrence and range as well as diversity of spore-pollen taxa are given in Figs. 4.3 and 4.4. The abundances in percentage of miospores and variations therein are shown for the MPT (Fig. 4.5) and the VB (Fig. 4.6) sites. The arrangement in assemblage zones is based on changes in these abundances. Stratigraphic ranges are shown in Figs. 4.3 and 4.4, the main features of each zone are given below. For the MPT site four assemblage zones are proposed (MPT 1-4), for the VB site five zones can be distinguished (VB 1-5).

4.4.1.2 Palynological assemblages and abundances for the Mid-Polish Trough (MPT) recordThe spore-pollen assemblage identified for the MPT is predominantly composed of gymnosperm pollen (ø 52.1 %) with Cerebropollenites (Pl. 4.I, Fig. 15, 16), Callialasporites (Pl. 4.I, Fig. 9, 10), Araucariacites (Pl. 4.I, Fig. 7, 8), Perinopollis (Pl. 4.I, Fig. 14) and Inaperturopollenites (Pl. 4.I, Fig. 12, 13) being the most important non-saccate forms. On average Classopollis (Pl. 4.I, Fig. 17, 18) accounts for only 6.3 % with the exeption of two samples where these pollen reach up to more than 73.0 %. Exesipollenites is of minor importance and ranges between 0.0 and 1.6 %. Less common non-conifer forms include Eucommiidites (Pl. 4.I, Fig. 23), Cycadopites (Pl. 4.I, Fig. 19, 20) and Ephedripites (Pl. 4.I, Fig. 21), which account for less than ø ~2.2 %. Most common bisaccates are Vitreisporites pallidus and Alisporites with ø 1.5 and 0.8 %, respectively. Spores represent an important component of the miospore assemblages and account for ø 45.4 %. Dominant taxa include the fern spores Cyathidites (Pl. 4.I, Fig. 31), Gleicheniidites (Pl. 4.I, Fig. 39), and Leiotriletes (Pl. 4.I, Fig. 24, 25). Abundant non-fern spores are the lycopods-derived Echinatisporis (Pl. 4.II, Fig. 53), Neoraistrickia (Pl. 4.I, Fig. 32), and Leptolepidites (Pl. 4.II, Fig. 58), with none of them exceeding 0.5 %. Aratrisporites (Pl. 4.I, Fig. 28) is restricted to two samples only, where it shows relatively high abundances up to 7.0 %. Spores of bryophyte origin are rare components of the assemblage, including Stereisporites (Pl. 4.II, Fig. 5.7), and Lycopodiumsporites amongst other forms.

Zone MPT1 – Upper Berriasian to lower Valanginian (0 to 3.9 m).This zone shows an increase in spore abundances up to 54.8 %, with Cyathidites being prominent (ø 10.6 %) and showing its peak occurrence (max. 21.5 %), decreasing values for bisaccates (avg. 17.4 %), and relatively stable conditions for non-saccate pollen (ø 40.5 %). Araucariacites shows a spike (max. 13.3 %) in a generally increasing trend (ø 4.5 %). Cycadopites shows peak values (max. 7.0 %). Following a peak in the middle of the zone diversity is decreasing.

Zone MPT 2 – Lower Valanginian to upper Valanginian (3.9 to 11.9 m).This zone is characterized by relatively stable abundances of non-saccate pollen (ø 38.0 %). Perinopollenites is decreasing (ø 7.0 %) whereas Callialasporites shows a steady increase (ø 6.7 %). Ephedripites is comparatively abundant in this zone (ø 0.8 %). Bisaccates show a decreasing trend (ø 13.6 %). Spores are gradually increasing (ø 45.8 %) with only Cyathidites showing a decreasing trend (ø 8.3 %). Monolete spores show slightly increased abundances (ø 0.4 %). Diversity is medium (20-25 taxa).

Zone MPT 3 – Upper Valanginian (11.9 to 13.9 m).The zone is characterized by peak of spore abundances (max. 61.3 %), mainly due to high abundances of

62

Chapter 4

Sam

ples

Pol

len

non-

sacc

ate

Pol

len

bisa

ccat

e

Alisporites

Classopollis

Araucariacites

Calliallasporites

Cerebropollenites

Spo

res

Perinopollenites

Inaperturopollenites

Pinuspollenites

Vitreisporites

Parvisaccites

Podocarpidites

8 6 24

10

12

14

18

16

Lith

olog

y

Thickness (m)

PIG

1.5

PIG

1.13

PIG

1.21

PIG

1.28

PIG

1.35

PIG

1.45

PIG

1.56

PIG

1.63

PIG

1.71

PIG

1.78

PIG

1.86

PIG

1.91

PIG

1.98

PIG

1.10

6

PIG

1.11

4

PIG

1.12

2P

IG1.

131

PIG

1.14

0

PIG

1.15

0

PIG

1.15

9

PIG

1.16

6

PIG

1.17

4

PIG

1.18

2

clay

sand

limes

tone

shel

ls

nodu

les

mar

ly c

lay

MP

T1

MP

T2

MP

T3

MP

T4Assemblage zones

Div

ersi

ty Number of taxa 2520

UPPERBerrias. pp.

UPPER VALANGINIAN pp.Sub-stages LOWER VALANGINIAN

Eucommiidites

Cycadopites

Ephedripites

Exesipollenites

Cicatricosisporites

Leptolepidites

Cyathidites

Gleicheniidites

Leiotriletes

Osmundacidites

Neoraistrickia

Concavisporites

Stereisporites

Echinatisporis

Trilobosporites

Foveosporites

Biretisporites

Camarozonosporites

RetriletesDensoisporites

Concavissimisporites

Microreticulatisporis

Todisporites

Ischyosporites

Staplinisporites

Clavatisporites

Foraminisporis

Lycopodiacidites

Aratrisporites

Lycopodiumsporites

Klukisporites

GemmatriletesVerrucosisporites

Aequitriradites

Fig.

4.3

. Abu

ndan

ces (

for l

egen

d se

e Fi

g. 4

.4) a

nd st

ratig

raph

ic o

ccur

renc

es o

f spo

rom

orph

taxa

of t

he M

PT, a

rran

ged

by st

ratig

raph

ic fi

rst o

ccur

renc

es. S

ub-s

tage

s, th

ickn

ess,

litho

logy

(see

le

gend

for d

etai

ls), p

ositi

on o

f sam

ples

(elo

ngat

ed li

nes:

paly

nolo

gica

l sam

ples

; sho

rt li

nes δ

13C

), di

vers

ity b

ased

on

num

ber o

f tax

a, a

nd a

ssem

blag

e zo

nes.

63

Reconstructing Valanginian (Early Cretaceous) mid-latitude vegetation and climate dynamics based on spore-pollen assemblages

limes

tone

blac

k sh

ale

mar

l

calc

arou

s m

arl

40

50

60

30

10

20

70

90

10

0

11

0

12

0

13

0

14

0

15

0

16

0

17

0

80

stac

king

pat

tern

Thickness (m)

Sub-stages

Ammonitezonation

Cyclostra--tigraphy

B. camp.K. biassalenseS. fuhriN. platycos.

B. campylotoxusUPPER VALANGINIAN

S. verrucosumK. pronecostatumS. verrucosumN. peregrinusC. furcillata

C. furcillata N. peregrinusO. nicklesiT. callidisc.

H 1 V 12 V 11 V 10 V 9 V 8 V 7 V 6 V 5

2 2 2 2 2 2 2 2 21 1 1 1 1 1 1 1 13 3 3 3 3 3 3 3 34 4 4 4 4 4 4 4 4

LOWER VALANGINIAN p.p.LOWER HAUTERIVIAN p.p.A. radiatus

Sam

ples V

ER

26

VE

R29

MO

R5

VE

R41

VE

R56

VE

R72

VE

R86

LC10

LC38

LC66

LC81

LC98

LC13

2

LC14

8

LC16

6

LC17

9

LC19

8

Pol

len

non-

sacc

ate

Pol

len

bisa

ccat

e

Classopollis

Araucariacites

Calliallasporites

Cerebropollenites

Spo

res

Cicatricosisporites

Inaperturopollenites

Ephedripites

Pinuspollenites

Retriletes

Klukisporites

Densoisporites

Lycopodiumsporites

Foraminisporis

Gemmatriletes

VE

R2

VE

R11

VE

R22

VE

R36

VE

R40

VE

R50

<66-

30>6

031

-60

not c

ount

able

, hig

h A

OM

not c

ount

able

, hig

h A

OM

not c

ount

able

, hig

h A

OM

VB

1

VB

2

VB

3

VB

4

VB

5

Div

ersi

ty Number of taxa 25

Assemblage zones

20

Eucommiidites

Cycadopites

Exesipollenites

Perinopollenites

Vitreisporites

Alisporites

Podocarpidites

Parvisaccites

Leptolepidites

Stereisporites

Neoraistrickia

Ischyosporites

Cyathidites

Echinatisporis

Gleicheniidites

Leiotriletes

Osmundacidites

Concavisporites

Biretisporites

Lycopodiacidites

Concavissimisporites

Verrucosisporites

Foveosporites

Trilobosporites

Camarozonosporites

Aequitriradites

Microreticulatisporis

Fig. 4.4. Abundances (see legend) and stratigraphic occurrences of sporomorph taxa of the VB, arranged by stratigraphic first occurrences. Sub-stages, ammonite zonation, cyclostratigraphy (adopted from Gréselle et al., 2011), thickness, lithology (see legend for details), position of samples (elongated lines: palynological samples; short lines: δ13C), diversity based on number of taxa, and assemblage zones.

64

Chapter 4

Con

ifers

Cyc

ads,

Gin

kgos

Gne

tale

sP

olle

n no

n-sa

ccat

eC

onife

rsP

olle

n bi

sacc

ate

Gin

kgos

,P

terid

ospe

rms

Cerebropo

llenit

es

Callialla

sporite

sArau

caria

cites

Classo

pollis

Exesip

ollen

ites

Perinop

ollenit

es Inapert

uropo

llenite

s

Eucommiid

ites

Cycad

opite

sEphe

dripit

es

Pollen s

ummed

bisacc

ate un

diff.Pinus

polle

nites

Parvisa

ccite

s

Podoca

rpidit

es

Alispo

rites

Vitreisp

orites

pallidu

s

bisacc

ate su

mmed

1010

1010

105

1020

1020

Thickness (m) 218 16 14 12 10 8 6 4

Sub-stages UPPER VALANGINIAN pp. LOWER VALANGINIANUPPERBerrias.

pp.

2010

550

55

55

530

505

55

5Pra

sino

phyt

esP

terid

ophy

tes

Spo

res

Lyco

pods

Prasinop

hyce

ae sum

med

Spore un

diff. Trilo

bosp

orites

Cicatric

osisp

orites

Cyathidi

tes

Gleich

eniidi

tes

Leiot

riletes Osm

unda

cidite

s

Klukispori

tes

Ischyo

spori

tes

Conca

vispo

rites

Biretis

porites

Conca

vissim

isporite

s

Verruco

sispo

rites

Gemmatrilete

s

Clavati

spori

tes

Todisp

orites

Lepto

lepidite

s

Echina

tispori

s

Retrile

tesDen

soisp

orites

Neoraist

rickia

Foveos

porites

Lyco

podia

cidite

s

Aequitri

radite

s

Camaro

zono

sporite

s

Aratris

porites

strig

osus

Staplin

ispori

tes

1010

2010

1010

Lyco

podiu

mspori

tes

525

55

55

55

55

55

55

55

55

55

5

MP

T1

MP

T2

MP

T3

MP

T4

Bry

ophy

tes

Spore m

onolete

Microretic

ulatis

poris

Foramini

spori

s

Spores s

ummed

Stereis

porite

s

55

55

5025

Conif.,

Gink

g.,

Pterido

sp.

Assemblagezones

2575

Fig.

4.5

. Spo

re-p

olle

n-as

sem

blag

es M

PT. S

ub-s

tage

s, th

ickn

ess,

abun

danc

es o

f mio

spor

es in

%, a

nd a

ssem

blag

e zo

nes.

65

Reconstructing Valanginian (Early Cretaceous) mid-latitude vegetation and climate dynamics based on spore-pollen assemblages

Pra

sino

phyt

es

Thickness (m)

Sub-stagesAmmonite-zonation

L. H. pp. UPPER VALANGINIAN L. VAL. pp.

A. radiatus C. furcillata N. peregrinus S. verrucosum B. campylotoxus

1040 30 208090 70 60 50100

110

120

160

150

140

130

170

Con

ifers

Gin

kgos

,C

ycad

sG

neta

les

Pol

len

non-

sacc

ate

Con

ifers

Conif.,

Gink

g.,

Pterido

sp.

Pol

len

bisa

ccat

eG

inkg

os,

Pte

ridos

perm

s

Cerebropo

llenit

es

Callialla

sporite

s Arauca

riacit

esClas

sopoll

is

Exesip

ollen

ites

Perinop

ollenit

es

Inapert

uropo

llenite

sEuco

mmiidite

s

Cycad

opite

s

Ephedri

pites

Pollen s

ummed

bisacc

ate un

diff. Pinuspo

llenit

es

Parvisa

ccite

s

Podoca

rpidit

es

Alispo

rites

Vitreisp

orites

pallidu

s

bisacc

ate su

mmed

1010

2025

505

1020

55

205

1020

105

550

105

55

5

Prasinop

hyce

ae sum

med

5

Pte

ridop

hyte

sS

pore

sLy

copo

ds

Spore un

diff.

Trilobo

spori

tes

Cicatric

osisp

orites

Cyathidi

tes Gleich

eniidi

tesLe

iotrile

tes Osmun

dacid

ites

Klukispori

tes

Ischyo

spori

tes

Conca

vispo

rites

Biretis

porites

Conca

vissim

isporite

s

Verruco

sispo

rites

Gemmatrilete

s

Lepto

lepidite

s

Echina

tispori

s

1020

510

1010

55

55

55

55

55

10

Retrile

tesDen

soisp

orites

Neoraist

rickia

Foveos

porites

Lyco

podia

cidite

s

Aequitri

radite

s

Camaro

zono

sporite

s

55

55

55

5

Lyco

podiu

mspori

tes

5

VB1

VB3

VB4

VB5

VB2

Bry

ophy

tes

Spore m

onolete

Microretic

ulatis

poris

Foramini

spori

s

Spores s

ummed

50

Stereis

porite

s

55

55

25

Assemblagezones

Fig.

4.6

. Sp

ore-

polle

n-as

sem

blag

es V

B. S

ub-s

tage

s, am

mon

ite z

ones

(fro

m G

rése

lle e

t al

., 20

11)

thic

knes

s, ab

unda

nces

of

mio

spor

es in

%, a

nd a

ssem

blag

e zo

nes.

66

Chapter 4

Cyathidites (max. 13.7 %), Gleicheniitides (7.4 %) and Leiotriletes (14.1 %). In contrast, non-saccate (max. 32.1 %) and bisaccates pollen (max. 14.7 %) show low abundances. Classopollis is very rare (ø 1.4 %). Vitreisporites pallidus shows an exception from this pattern and reaches peak values (max. 6.3 %). Diversity is decreasing.

Zone MPT 4– Upper Valanginian (13.9 to 17.8 m). This uppermost zone of the Wąwał section zone is characterized by two pronounced peaks in bisaccate pollen (max. 31.1 %). Vitreisporites pallidus shows maximum values (max. 4.3 %). Araucariacites shows an increasing trend (ø 7.4 %) while percentages of non-saccate pollen remain relatively stable (ø 36.9 %). Perinopollenites peaks towards the end of the zone (max. 18.7 %). Spores abundances are quite stable (ø 44.0 %) after a minimum at the beginning of the zone (23.8 %), with the exception of declining numbers of Cyathidites (ø 7.0 %) and Leiotriletes (avg. 3.6 %). Diversity is medium with the uppermost sample being of especially high (>25 taxa) diversity.

4.4.1.3 Palynological assemblages and abundances for the Vocontian Basin (VB) recordIn the VB, the identified spore-pollen assemblages are mainly composed of gymnosperm pollen (ø 60.2 %) with Classopollis (Pl. 4.I, Fig. 17, 18), Araucariacites (Pl. 4.I, Fig. 7, 8), Inaperturopollenites (Pl. 4.I, Fig. 12, 13), and Callialasporites (Pl. 4.I, Fig. 9, 10) being the most important non-saccate forms. Cerebropollenites (Pl. 4.I, Fig. 15, 16) is of minor importance with ø 3.8 %, but reaches maxima of up to 11.7 %. Less common non-conifer forms include Eucommiidites (Pl. 4.I, Fig. 23), Cycadopites (Pl. 4.I, Fig. 19, 20) and Ephedripites (Pl. 4.I, Fig. 21), which account for less than ø ~1.6 %. Most common bisaccates are Vitreisporites pallidus and Alisporites with ø 1.3 and 0.6 %, respectively. Spores occur with ø 38.6 %. The fern spores Cyathidites (Pl. 4.I, Fig. 31), followed by Gleicheniidites (Pl.4.I, Fig. 39) and Leiotriletes (Pl. 4.I, Fig. 24, 25) are dominant spores. Abundant non-fern spores are the lycopods Echinatisporis (Pl. 4.II, Fig. 53), and Leptolepidites (Pl. 4.II, Fig. 58), with numbers up to ø 2.8 %. An regularly recorded bryophyte is Stereisporites (Pl. 4.II, Fig. 57; ø 0.2 %).

Zone VB1 –Lower Valanginian (0 to 19 m).This zone is characterized by relatively high abundances of spores, exhibiting a decreasing trend (ø 43.5 %), with the dominant taxa being Cyathidites (max. 14.7 %), Gleicheniidites (max. 9.3 %) and Leiotriletes (max. 7.0 %). Bisaccates as well show a decreasing trend (ø 9.7 %). Classopollis (ø 12.3 %), Inaperturopollenites (ø 8.6 %), and Araucariacites (ø 9.4 %) are dominant non-saccate pollen. Diversity is medium.

Zone VB2 –Lower Valanginian to upper Valanginian (19 to 47 m).Non-saccate pollen peak in the middle part of this zone (max. 74.3 %), whereas spores (ø 26.8 %) and bisaccates (ø 9.4 %) are sharply decreasing in abundances. Classopollis (max. 45.0 %) and Araucariacites (max. 19.0 %) represent the dominant non-saccate pollen. Callialasporites shows high abundances in the uppermost part of the zone (max. 15.0 %). Spores are dominated by Leiotriletes (max. 6.7 %) and Cyathidites (max. 10.0 %). Diversity is low (<20 taxa).

Zone VB3 – Upper Valanginian (47 to 65 m).This zone is characterized by a pronounced increase in spore abundances (max. 56.0 %). Leiotriletes

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(max. 11.7 %), Gleicheniidites (max. 15.3 %), and Concavisporites (max. 2.7 %) show maximum abundances. Non-saccate pollen (ø 35.0 %) and bisaccates (ø 13.4 %) show only low abundances. Classopollis (max. 11.0 %) and Inaperturopollenites (max. 12.7 %) show their lowest abundances. Diversity is fluctuating from high to low.

Zone VB4 – Upper Valanginian (65 to 145 m).This zone is characterized by an increasing trend in the abundance of spores (ø 40.1 %) following low values at the very base of this zone (min. 31.3 %). Cyathidites is dominating (max. 15.3 %) and Echinatisporis (max. 7.7 %) reaches maximum values in the upper part of the zone. Bisaccates show an opposing trend, after peak values at the beginning of the zone (max. 27.3 %) abundances are decreasing (min. 3.7 %). The pteridosperms/ginkgos Vitreisporites pallidus and Alisporites show maximum abundances (max. 4.7 %, and max. 4.7 %, respectively). Non-saccate pollen are relatively stable (ø 44.0 %), with Classopollis being the dominant taxon (max. 28.0 %). Diversity is declining towards the uppermost part of the zone.

Zone VB5 –Upper Valanginian to Lower Hauterivian (145 to 176 m).This zone is characterized by a second maximum in spore abundances (max. 63.6 %) at the very base, followed by a decreasing trend, while non-saccate pollen (ø 47.7 %) show an opposing pattern. Bisaccates remain relatively stable (ø 7.5 %). Classopollis and Inaperturopollenites are the dominant non-saccate pollen taxa (max. 25.8 %, and max. 20.9 %, respectively) with the latter showing maximum values for the whole section. Diversity is fluctuating.

4.4.2 Nannofossil biostratigraphyThe preservation of the calcareous nannofossil assemblages of the MPT varies throughout the studied section. Samples PIG 1.1 to 1.9 are barren of calcareous nannofossils, samples PIG 1.10 and 1.11 yield only a few preserved specimens. The calcareous nannofossils of interval PIG 1.12 to 1.175 are well preserved, except for nine samples, PIG 1.12, PIG 1.16, PIG 1.19, PIG 1.25, PIG 1.114, PIG 1.116, PIG 1.118, PIG 1.174, PIG 1.175, which show a moderate preservation. Samples PIG 1.178 to 1.188 are barren. Seven nannofossil events including first occurrences (FOs) and last occurrences (LOs) were observed throughout the upper Berriasian to Valanginian interval. These allowed the recognition of four Boreal calcareous nannofossil zones (BC2–BC5). The occurrence of Sollasites arcuatus in samples PIG 1.012 to 1.021 imply a late Berriasian age (BC2) for this interval. Based on the absence of both S. arcuatus and Triquerhabdulus shetlandensis in samples PIG 1.22 to 1.25 this interval has been assigned to nannofossil subzone BC3a (Berriasian/Valanginian boundary). The base of the following BC3b subzone (lowermost Valanginian) is marked by the FO of T. shetlandensis (sample PIG 1.26). The top of this zone is defined by the FO of Micrantholithus speetonensis (sample PIG 1.44). The total range of M. speetonensis (samples PIG 1.44 to 1.112), a taxon which is rare throughout the studied section, spans the lower Valanginian (BC4 Zone). This zone can further be subdivided into the subzones BC4a and BC4b by using the FO of Eiffellithus windii (sample PIG 1.96). The BC5 Zone (samples PIG 1.113 to 1.175) spans the entire upper Valanginian and can be subdivided by the FO of Eiffellithus striatus (sample PIG 1.142). The base of this zone is defined by the LO of M. speetonensis, the top by the LO of T. shetlandensis. The upper boundary of this zone, however, is arbitrary since the overlying interval is barren of calcareous nannofossils.

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4.4.3 Stable carbon isotopesThe δ13Ccarb signals for the VB vary between 0.1 and 2.6 ‰ and show a positive CIE of about 2 ‰ (Fig. 4.2). The record shows no offset of δ13Ccarb values between the different sections, enabling the establishment of a composite curve. In the NK3A nannofossil zone values are comparatively stable at the beginning (0 to 26 m, ~ 0.7 ‰), with a sharp decrease to lowest values (~ 0.1 ‰ at 26 m). After that a first increase to more positive values up to ~1.7 ‰ (26 to 32 m) can be observed. In the lowest NK3B zone (32 to 45 m) values are stagnant around 1.5 ‰, with a rapid shift to more positive values (to ~2.5 ‰ at 45 to 50 m). Values remain high with two phases of maximum values of up to ~2.7 ‰ in the middle NK3B and around the transition of zone NK3B to NC4. Subsequently in zone NC4 a decreasing trend establishes (until 150 m in the upper NC4 zone), with the uppermost part showing stable values (around 1.1 ‰). The lower part of this record (Vergol and Morenas sections) was shown to reveal a comparable pattern to other Tethyan δ13Ccarb records (Kujau et al., 2012), which affirms the stratigraphic framework.

4.5 Discussion4.5.1 Correlation of sitesThe two studied sites can be correlated following Bown (1998; Fig. 4.2) by using the calcareous nannofossil zonation schemes of the Tethys (NC) and of the Boreal Realm (BC). The NC zonation for the well dated VB has been adopted from Gréselle et al. (2011; see also Reboulet et al., 2003), accomplished independently for the same sampling sites. The Polish site covers sediments from the Berriasian/lower Valanginian to upper Valanginian, evidenced by the Boreal nannofossil zones BC2 to BC5. For the VB a range between the lower Valanginian to Lower Hauterivian is documented by the Tethyan nannofossil zones NK3A to NC4. Additional time constraint is provided by chemostratigraphic correlation of the two sites; both δ13Ccarb records reveal comparable trends. The comparatively low δ13C values established prior to the lower/upper Valanginian boundary (MPT ~1 ‰, VB ~0.5 ‰), are followed by a sharp increase over the lower/upper Valanginian boundary (MPT up to ~3 ‰, VB up to ~2.5 ‰). Subsequently, values remain high (interrupted by one sample of especially low values at the MPT at 13.3 m) before they turn into a decreasing trend (MPT at 17.0 m, VB at ~120 m). The MPT record (Morales et al., in prep.) is, however, characterized by high fluctuations hampering a high-resolution chemostratigraphic correlation. Following the Boreal-Tethys correlation of Bown (1998), the overlapping interval for the MPT and the VB spans sediments from the lower Valanginian to upper Valanginian, upper NK3A to NK3B and upper BC4a to BC5, respectively.

4.5.2 Paleoenvironmental significanceEarly Cretaceous spore-pollen taxa are known to be long-ranging and thereby no suitable biostratigraphic tool on shorter time scales (Taugourdeau-Lantz, 1988). They can, however, be used for reconstructing paleoenvironmental conditions. Based on the spore-pollen assemblages, an image of the ancient plant communities can be inferred. Most spore-pollen can be assigned to parent plants or major plant groups (botanical affinity). Variations in spore-pollen assemblages are thereby reflecting changes in the vegetation of the hinterland (e.g. Traverse, 2007). By comparing the spore-pollen assemblages of two different sites, supra-regional trends in environmental and climatic changes can be differentiated from those of only local expression. The composition and abundances of spores and pollen in the sedimentary record is, however, not only reflecting the vegetation of a particular area and i.e. environmental and climatic conditions. It is also influenced by a number of factors that affect the transport of miospores, such as wind direction and strength,

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as well as by sedimentation processes in connection with changes in sea-level, marine currents, and river drainage (Hochuli and Kelts, 1980; Schrank, 2010 and references therein). Due to the specific morphology of some taxa, like bisaccates and Classopollis, that favours long-distance atmospheric and hydrologic transport (Tyson, 1995; Birkenmajer et al., 2010; Schrank, 2010), these taxa may be selectively accumulated in marine environments. Another factor influencing abundances of spore-pollen is the number of miospores produced by the assigned parent plant. Pinus trees for example are strong pollen producers (Tyson, 1995; Birkenmajer et al., 2010 and references therein). Despite these limitations overall environmental and climatic trends as well as comparisons between different sites can be inferred from palynological records.

4.5.3 Paleophytogeographic frameworkFor the middle to Late Cretaceous a wealth of studies dealing with continental palynology exist (e.g. Brenner, 1976; Hochuli, 1981; Batten, 1984; Chumakov et al., 1995; Zhichen and Fei, 1997), whereas publications with a focus on the earliest Cretaceous are comparatively rare. Two major floral provinces have been proposed for the northern hemisphere of the Early Cretaceous by Ziegler et al. (1987, modified after Vakhrameev, 1978), with a boundary at around 50°N, namely a cool temperate northern Siberian-Canadian province, dominated by gingkos and gingko-like Czekanowskiales, and a more southern warm temperate (Indo-)European province, rich in pteridophytes, including Weichselia, and pteridosperms like Tempesakya. During the Early Cretaceous, unlike today, maximum plant productivity and diversity was concentrated at the mid- and high latitudes (c.f. Ziegler et al., 1987; Rees Mc Allister et al., 2004; Vakhrameev, 2010). The continental areas around the VB and the MPT were part of the warm temperate European province, without seasonal freezing. It was reported that the composition of the vegetation of the northern hemisphere changed markedly close to the Late Jurassic-Early Cretaceous boundary, with the rapid appearance of many new groups of pteridophytes in the mid-latitudes (Batten, 1984; Herngreen et al., 1996). This change can probably be assigned to the ongoing separation of land masses and fluctuations in sea-level that caused the formation of epicontinental seas and archipelagic coastal habitats, accelerating endemic evolution (e.g. Diéguez et al., 2010). A general increase in pteridophyte abundances compared to the Jurassic was interpreted as an increase in humidity (Vajda, 2001). Pteridophytes are a dominant and diverse floral element at both investigated sites. Observed pteridophytes are of dickinsoniacean, cyatheaceaen, dipteridacean, filicopsidan or matoniacean as well as gleicheniacean origin. Most abundant forms are Cyathidites, Gleicheniidites and Leiotriletes. Gleicheniaceae and Schizeaceae were reported to have spread during the Early Cretaceous, being typical floral elements of Europe, with Gleicheniaceae being the dominant ferns (Herngreen et al., 1996; Vakhrameev 2010). High abundances of Deltoidospora (e.g. Vajda, 2001, probably synonymous for Cyathidites and Leiotriletes), Clavifera (in this study assigned to Gleicheniidites), and Gleicheniidites, were also reported from Sweden (Vajda, 2001). Also in accordance to the Swedish study are low abundances of Schizeaceae like Cicatricosisporites and Klukisporites. Reported common Jurassic forms that were still abundant in the Early Cretaceous and were found in this study are Gleicheniidites, Ischyosporites/Klukisporites, and Concavissimisporites/Trilobosporites. Apart from similarities in the composition of pteridophyte assemblages, the MPT and VB sites reveal distinct differences. The paleoflora surrounding the MPT is dominated by conifers of araucarian and cupressacean origin, whereas cheirolepidiacean conifers are of minor importance. The paleoflora in the hinterland of the VB is as well characterized by conifers like Araucarians and Cupressaceans, but also by drought-resistant Cheirolepidiaceans. Herngreen et al. (1996) as well reported the abundance of gymnosperm pollen like

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Araucariacites (Araucariaceae)/Inaperturopollenites and Calliallasporites for the European realm. The contribution of Cheirolepediaceae like Brachyophyllum, Pagyophyllum, Frenelopsis and Pseudofrenelopsis, reflected in high amounts of Classopollis (up to 70 % during arid intervals) was reported in addition to be significant for European assemblages (Herngreen et al., 1996; Vakhrameev, 2010). Their rarity at the MPT may be related to the latitudinal position of this site further north. In contrast to Ziegler et al. (1987), but in accordance to Herngreen et al. (1996) ephedroid and monosulcate pollen are scarce at the MPT and VB. Ginkgoaceae, were reported to be of low diversity, only reflected by the genus Ginkgo (Vakhrameev, 2010). In this study Cycadopites of potential cycadalean, bennettitalean or ginkgoalean origin show only temporarily moderate abundances. The discrepancies may be due to specific local characters of the analyzed vegetation. Ginkgoales, Czekanowskiales, and cycadophytes are supposed to have dramatically declined from the Jurassic to the Early Cretaceous (MacLeod and Hills, 1991).Moderate abundance was observed for Pinaceae and Podocarpaceae (Batten, 1984; Vakhrameev, 2010). Herngreen et al. (1996), however, mentioned considerable amounts of bisaccates like Podocarpidites, Vitreisporites, and Alisporites. The latter two taxa can probably both be assigned to the pteridosperms, which were reported as important and diverse floral elements of the European realm (Ziegler et al., 1987; Herngreen et al., 1996). This is in accordance to the present study, where at both sites Vitreisporites and Alisporites show high abundances. A more detailed analysis of the bisaccates is, however, hampered by the high amount of non-identifiable pollen at both sites.

4.5.4 Temporal evolution of habitats and environmental moisture levelsA spore-pollen ratio for each site is calculated (Fig. 4.7), with high values indicating high spore-abundances. Spore-producing plants (pteridophytes, lycopods, bryophytes) are in general considered to be abundant under relatively humid conditions, optimal for their growth and reproduction. High spore abundances are thereby interpreted to reflect increased humidity (Visscher and van der Zwan, 1981; Fowell et al., 1994; Abbink, 1998; Vajda, 2001; Bornemann et al., 2005; Hochuli and Vigran, 2010; Schrank, 2010). On the other hand, gymnosperms are assumed to dominate over spore-producers under comparatively dry conditions (Fowell et al., 1994). Therewith, the spore-pollen ratio is used here to infer trends in environmental moisture levels. Further indication of moisture trends stems from variations in the abundance of the drought-resistant conifer pollen Classopollis of cheirolepidiacean origin, with high values pointing to arid conditions (Fig. 4.7). Besides changes in the spore-pollen ratio and Classopollis abundances, successive changes in plant community structure, habitat, and moisture availability are described based on botanical affinities of significant taxa (Table 4.2, c.f.; Fig. 4.5, 4.6).

4.5.4.1 The MPT siteFor the MPT abundances of Classopollis are overall low (max. 13.3 %), with temporarily slightly enhanced values pointing to drier conditions. Two samples of the lower MPT1 (PIG1.8) and lower MPT2 (PIG1.56) assemblage zones are excluded from interpretation due to their extremely high numbers of Classopollis (73.0 and 55 %, respectively, see Fig. 4.5). They may represent short-lived extremely dry conditions but more probably do not reflect changes in the vegetation structure but in sedimentological processes artificially enhancing the Classopollis abundances in these samples (see also 4.5.2). This interpretation is supported by the fact that these samples represent intervals rich in carbonate and shells, deviating from the rest of the section (Fig. 4.3).

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Fig. 4.7. Comparison of spore-pollen ratios and Classopollis abundances between the VB (left) and MPT (right). Sub-stages, ammonite zones, assemblages zones, Classopollis (%), spore-pollen ratio, main trends as bold grey lines. Orientation is following the lower to upper Valanginian boundary (broken line).

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Spore abundance is increasing in the MPT1 assemblage zone, and culminates in a first spore maximum at the MPT1 to MPT2 assemblage zone boundary. This peak is characterized by high abundances of the fern spores Cyathidites and Gleicheniidites and a reduction in the gymnosperm pollen of conifer origin, namely Perinopollenites, pointing to a reduction in Cupressaceae/Taxodiaceae. High numbers of the conifer pollen of probably araucarian and taxodiacean type (Inaperturopollentites, Calliallisporites, Cerebropollenites) in assemblage zones MPT1 and MPT2 point to a swampy shoreline environment (Markevich et al., 2009). In assemblage zone MPT2, a gradual trend towards more humid conditions is observed by increasing numbers of spores. In the upper half of assemblage zone MPT2, the spore-pollen ratio shows strong fluctuations. Lycopods are slightly more common up to the end of assemblage zone MPT2, as reflected in high numbers of Neoraistrickia, pointing towards humid lowland, river or tidally-influenced habitats (Abbink et al., 2004). The gymnosperm pollen Callialasporites and Inaperturopollenites become dominant tree pollen. Persistently higher numbers of the fern spore Osmundacidites are a further indication for a humid environment of floodplains and swamps (Coiffard et al., 2007). Towards the end of MPT2 Cyathidites is declining while non-identifiable spores are increasing. The gradual change in plant communities might be partly due to a reduction of coastal plains by increasing sea-level. This is confirmed by a decrease in the conifer pollen Perinopollenites, of probably cupressacean or taxodiacean origin, whose parent plants were probably adapted

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to coastal environments, which would be expected to be reduced under rising sea-level. Maximum humidity is established relatively abruptly in assemblage zone MPT3, evidenced by a pronounced spore maximum characterized by high numbers of the fern spores Cyathidites, Gleicheniidites and Leiotriletes, e.g. of Dickinsoniacean, Gleicheniacean, and Osmundacean origin as well as by undifferentiated spores. The generally abundant gymnosperm pollen of cupressacean/taxodiacean origin, Cerebropollenites and Perinopollenites, are decreasing. Classopollis abundances consequently show their lowest values. Bryophytes are increasing towards assemblage zone MPT3, confirming a trend towards higher humidity (Schrank, 2010). High abundances of the bisaccate pteridosperm pollen Vitreisporites pallidus, peaking in assemblage zone MPT3, accompanying the spore-maximum, provide evidence for the establishment of swamps and a dominance of deltaic environments and backswamps. The peak in the fern spore Leiotriletes points to abundant herbaceous swamps. Trees represent rare floral elements in assemblage zone MPT3. Assemblage zone MPT4 is characterized by high bisaccate abundances as well as by high amounts of the araucarian pollen Calliallasporites and Araucariacites, indicating relatively moist, temperate conditions. The spore-pollen ratio is decreasing to pre-maximum abundances.

4.5.4.2 The VB siteFor the VB, Classopollis shows overall high abundances, temporarily extremely enhanced (max. 45  %). Assemblage zone VB1 is characterized by a trend towards arid conditions, indicated by a decreasing spore-pollen ratio and increasing Classopollis abundances (Fig. 4.6, 4.7). Abundances of other conifer pollen like Calliallasporites, Cerebropollenites, and Inaperturopollenites are decreasing. This may reflect a disturbed environment, with a reduced competitiveness of trees under increasingly dry conditions. A loss in diversity (Fig. 4.4) may also be caused by this aridification event that culminates in assemblage zone VB2. This comparatively dry interval is followed by an abrupt shift in the spore-pollen ratio to maximum values in assemblage zone VB3, indicating relatively humid conditions. This maximum in the spore-pollen ratio is characterized by high abundances of the fern spores Cyathidites, Gleicheniidites, and Leiotriletes of e.g. dickinsoniacean, gleicheniacean, and osmundacean origin as well as by spores of unclear taxonomy. At the same time, the coniferous gymnosperm pollen Classopollis shows low abundances, which indicates a reduction in xerophytic vegetation. A peak in Leiotriletes in assemblage zone VB3 points to abundant herbaceous swamps, with a rare arborescent vegetation. High abundances of the pteridosperm pollen Vitreisporites points to the establishment of deltaic environments and backswamps at the VB3/VB4 assemblage zone boundary. Climatic conditions in assemblage zone VB4 are interpreted as relatively cold and rather dry, based on a combination of high amounts of bisaccates and a low diversity in fern spores (Herngreen et al., 1996). Increased abundances of Exesipollenites and Cycadopites as well point to drier conditions. A reestablishment of forested areas may be inferred from an increase in Araucariacites pollen towards the end of the assemblage zone. A gradual shift towards humid conditions culminates in another spore maximum at the assemblage zone boundary between VB4 and VB5. This second spore maximum shows again high numbers of the fern spores Cyathidites, Gleicheniidites, and Leiotriletes as well as of undefined spores. Bisaccate pollen show low abundances. In assemblage zone VB5, a trend towards more arid conditions is following the second maximum in the spore-pollen ratio, accompanied by the establishment of a tree-dominated vegetation, with abundant Araucariacites, Perinopollenites, and Cerebropollenites. Unlike the first spore maximum in assemblage zone VB3, this second spore-maximum is not preceded by elevated numbers of Classopollis.

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In the VB, abundances in Classopollis are significantly higher (max. 45.0 %) compared to the MPT (max. 13.3 %). Furthermore, in the VB, the abundances of Classopollis fluctuate on a high amplitude, while at the MPT they are comparatively stable. This may be due to an in general higher level of humidity around the MPT, in accordance with abundant swamp habitats, whereas around the VB swamp habitats were only temporarily abundant. The most characteristic feature observed at both sites is represented by a coeval maximum in the spore-pollen ratio (assemblage zones VB3/MPT3), interpreted to reflect supra-regionally exceptionally humid conditions in the early late Valanginian. The composition of spore taxa during this coeval spore maximum is shown to be similar at both sites. Subsequently, climatic conditions became again drier at both localities. At the MPT site, the phase of increased humidity is preceded by a gradual trend towards humid conditions. In contrast, palynological evidence at the VB provides evidence for a protracted phase of relatively arid conditions preceding the humidity peak. At both sites comparable vegetation structures reflect especially humid conditions after the early/late Valanginian boundary. Preceding conditions are, however, dissimilar, indicating different environmental evolutions.

4.5.5 Paleoclimatic implicationsDuring the Valanginian the coastlines of Paleo-Europe showed an archipelagic arrangement (Diéguez et al., 2010); it was located by 10° closer to the equator, adjoining the Tethys Ocean to the south, while the North Atlantic was not yet well established (e.g. Ziegler et al, 1987). Accordingly, the major moisture source for the southern part of Paleo-Europe was the Tethys, probably controlled by a monsoonal climate. During northern hemisphere summer, southern monsoonal trade winds were likely to be established over the northwestern Tethys (Fig. 4.8; c.f. Herrle et al., 2003; Wang, 2009; Giorgino et al., 2012). On geologic time scales variations in the monsoonal circulation are significantly controlled by changes in orbital parameters, known as Milankovitch cycles, which control the geographic distribution of insolation (c.f. Mayer and Appel, 1999; Crowley, 2002; Wang, 2009). The gradual increase in humidity in assemblage zone MPT2 (Fig. 4.7), culminating in extremely humid conditions expressed at both sites in assemblage zones MPT3/VB3, may represent an increase in the intensity of a monsoonal precipitation, and/or an extension of the annual humid period. This may explain for the observed restructuring of the vegetation towards a pteridophyte dominance. Eccentricity cycles are known to be responsible for longer-term variations in monsoonal intensity, with phases of high intensity occurring

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Fig. 4.8. Paleogeographic map of the Valanginian (supposed land in grey, based on Ziegler et al., 1987; Smith et al., 1994; Hay, 2002; Mutterlose et al., 2003; Rees McAllister et al., 2004; Blakey, 2010; Vakhrameev, 2010). Asterisks mark study sites (upper one MPT, lower one VB), arrows mark trade winds. Potential northernmost ITCZ trajectory of Valanginian northern hemisphere summer (broken line) defining the area of influence of a northern hemisphere paleo-monsoon (monsoonal reconstruction follows Wang, 2009, and Roedel and Wagner, 2011). H=high presure, L=low pressure.

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during maxima in eccentricity (e.g. Wang, 2009). The 2.4 Ma eccentricity cycle was shown to be in a maximum during the initial phase of the positive carbon isotope excursion of the Valanginian (Sprovieri et al., 2006). This phase coincides with the interval of the coeval spore maxima. Consequently, monsoonal precipitation can be expected to have been in a maximum. This scenario of enhanced precipitation may have been intensified by a general northward shift of the Inner Tropical Convergence Zone (ITCZ), defining the center of northern and southern trade winds (e.g. Roedel and Wagner, 2011). This may have prolonged the annual humid period for both sites. This shift in the ITCZ can be caused by specific constellations in the orbital parameters obliquity and precision, controlling the affinity of the ITCZ track towards the northern or southern hemisphere, explaining for short-term variations and insolation differences between the hemispheres (Häckel, 1999; Wang et al., 2003; Lauer and Bendix, 2006). Furthermore, factors like solar maxima or positive precipitation-vegetation-feedbacks may have additionally influenced humidity (e.g. Ganopolski et al., 1998; Vaughan, 2007). The preceding and subsequent phases of less pronounced maxima in the spore-pollen ratios (assemblage zone boundaries MPT1/2 and VB4/5, respectively) coincide with a preceding and a subsequent eccentricity maximum (Sprovieri et al., 2006). The fact that these spore maxima are less pronounced compared to that coevally occurring in assemblage zones MPT3 and VB3 may be explained by the lack of further intensifying factors. The abrupt drying indicated for the VB in assemblage zone VB2, preceding the supra-regional pronounced humid phase of the coeval spore-maximum, is characterized by high abundances of Classopollis and low spore abundances. It may reflect a northward shift or expansion of the northern hemisphere arid belt, probably due to a northward shift/expansion of the subtropical high-pressure belt (c.f. Hasegawa et al., 2010). This extension of the arid belt may have reduced moisture availability in the hinterland of the VB. The fact that this dry phase is not reflected in the continental areas surrounding the Carpathian seaway may be due to its paleogeographic location at a higher latitude. Further evidence for a meandering Valanginian arid belt is provided by aeolian deposits found on the Asian continent up to ~44°N (Hasegawa et al., 2011). Herngreen et al. (1996) as well describe an arid belt affecting the northern Tethys coastline throughout the Berriasian to Hauterivian, based on abundances of Classopollis. The subsequent phase of supra-regional humid conditions probably reflects a retread of the arid belt. A potential cold-snap interval (e.g. Kemper, 1987; Weissert and Lini, 1991; Podlaha et al., 1998; Pucéat et al., 2003; McArthur et al., 2007; Brassell, 2009) may have decelerated the hydrologic cycle reflected by arid conditions in assemblage zone VB4. Potential changes in Valanginian temperature, however, have probably been less relevant for environmental changes and seasonality than changes in moisture levels, controlled by monsoonal climates.

4.6 ConclusionsBased on spore-pollen abundances of two mid-latitudinal sites of Paleo-Europe in the MPT and the VB, changes in Valanginian vegetation patterns were analysed and used to deduce varying moisture levels. Both sites are similarly characterized by high pteridophyte abundances represented by Cyathidites, Leiotriletes and Gleicheniidites. Dominant pollen are different for both sites, at the MPT araucarians/cupressaceans dominate, while at the VB cheirolepediaceans are most abundant, which are extremely rare at the MPT. Bisaccates are of moderate abundance at both sites. Moisture availability was probably the determining factor for changes in vegetation patterns, controlled by a monsoonal circulation. A supra-regional interval of most pronounced humid conditions, indicated by a coeval spore-maximum, is established for the interval following the early/

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late Valanginian boundary, probably under strong monsoonal influence. Prior to the early/late Valanginian boundary a temporary northward shift or expansion of an arid belt is expressed by local differences between the sites causing a temporary drying at the VB, characterized by increased abundances of the drought-resistant conifer pollen Classopollis and reduced spore abundances. This dry physe is not reflected by the vegetation around the MPT. A proposed Valanginian “cold snap” is perhaps evidenced by an interval of more arid conditions in the late Valanginian, characterized by enhanced bisaccate pollen and relatively reduced spore abundances.

AcknowledgementsSpecial thanks are due to Benjamin Gréselle (Neftex Petroleum Consultants Ldt., Oxfordshire, UK) for his assistance during field work in France. Financial support from the DFG project HE4467/2-1 and the Swiss National Science Foundation (project 200020_126455) are gratefully acknowledged.

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Plate 4.INames of taxa with sample and coordinates of the photo. Latin numbers according to pictures on plates, indicating one (two) example(s) of taxa. For references see Table 4.2.

Pollen:1 Alisporites PIG1.56, 133.3/13.12 Alisporites PIG1.56, 146.9/19.93 Parvisaccites PIG1.56, 146.1/11.34 Parvisaccites PIG1.56, 134.0/16.55 Podocarpidites PIG1.56, 140.1/13.46 Pinuspollenites PIG1.122, 140.7/18.17 Araucariacites PIG1.21, 130.8/10.48 Araucariacites PIG1.122, 141.9/6.79 Callialasporites PIG1.98, 142.9/19.510 Callialasporites PIG1.13, 131.4/7.511 Vitreisporites PIG1.35, 135.1/16.312 Inaperturopollenites PIG1.98, 134.9/13.313 Inaperturopollenites PIG1.28, 141.7/18.714 Perinopollenites PIG1.45, 134.2/7.915 Cerebropollenites PIG1.56, 129.4/7.316 Cerebropollenites PIG1.56, 149.1/14.217 Classopollis PIG1.56, 146.2/14.218 Classopollis PIG1.122, 146.7/16.719 Cycadopites PIG1.56, 132.9/17.820 Cycadopites PIG1.35, 144.2/4.221 Ephedrepites PIG1.71, 134.5/12.022 Exesipollenites PIG1.71, 135.3/18.723 Eucommiidites PIG1.13, 126.7/12.7

Spores:24 Leiotriletes PIG1.71, 126.4/8.725 Leiotriletes PIG1.21, 144.3/10.426 Cicatricosisporites PIG1.5, 142.1/11.927 Cicatricosisporites PIG1.13, 136.3/13.828 Aratrisporites PIG1.56, 137.8/10.029 Trilobosporites PIG1.63, 132.8/7.330 Gemmatriletes PIG1.166, 152.8/6.231 Cyathidites VER78, 144.5/17.432 Neoraistrickia PIG1.174, 144.6/19.133 Ischyosporites PIG1.45, 138.5/7.934 Ischyosporites PIG1.45, 132.4/7.735 Concavisporites PIG1.91, 134.9/4.536 Concavisporites PIG1.45, 143.9/19.937 Trilobosporites PIG1.35, 140.9/11.538 Concavissimisporites PIG1.114, 138.5/13.339 Gleicheniidites PIG1.91, 128.9/15.540 Verrucosisporites PIG1.49, 149.1/9.6

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Plate 4.I

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Plate 4.IINames of taxa with sample and coordinates of the photo. Latin numbers according to pictures on plates, indicating one (two) example(s) of taxa. For references see Table 4.2.

Spores:41 Klukisporites PIG1.35, 145.2/6.042 Todisporites PIG1.150, 129.3/9.743 Birtetisporites PIG1.91, 150.7/15.644 Clavatisporites PIG1.35, 148.6/7.645 Retriletes PIG1.166, 148.9/7.146 Retriletes PIG1.56, 128.5/5.747 Osmundaciidites PIG1.166, 150.0/14.748 Camarozonosporites PIG1.13, 130.0/14.549 Gleicheniidites PIG1.45, 139.5/11.650 Densoisporites PIG1.21, 146.8/8.751 Foveosporites PIG1.159, 140.0/15.452 Lycopodiacidites PIG1.91, 143.5/4.353 Echinatisporis PIG1.13, 148.2/12.654 Lycopodiumsporites PIG1.106, 136.3/8.655 Foraminisporis PIG1.45, 139.2/9.756 Staplinisporites PIG1.159, 140.0/15.457 Stereisporites PIG1.71, 150.1/10.558 Leptolepidites PIG1.59, 14.1/133.459 Aequitriradites PIG1.45, 138.4/16.560 Microreticulatisporis PIG1.45, 147.5/17.4

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Plate 4.II

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5. Fluctuations in sea-level and terrestrial input at the NW Tethys and the Carpathian seaway during the Valanginian

(Early Cretaceous) –

Evidence from palynofacies and n-alkanes

AbstractThe Valanginian (Early Cretaceous) was a time of major perturbations of the global carbon cycle, accompanied by severe climatic and environmental changes. Causes, consequences, and interlinkages are an ongoing matter of research. Severe carbonate platform drowning is known to have initiated around the early/late Valanginian boundary, initial triggers are still questionable. The record of Valanginian sea-level fluctuations, probably related to platform drowning, is not yet fully understood, the sequence of major trends is still debated. In this study, palynofacies and biomarkers in the form of n-alkane ratios were investigated for two mid-latitudinal sites of Valanginian age. One site is located in the northwestern Tethys (Vocontian Basin, southeast France), the other one in the Carpathian seaway (Mid-Polish Trough, central Poland). The chosen approach provides a deeper insight into changes in the organic matter of the investigated sites. Sea-level fluctuations were reconstructed based on the varying composition of marine and terrestrial palynomorphs. Furthermore, variations in terrestrial input and deduced paleoenvironmental conditions have been analyzed based on changes within the terrestrial palynofacies fraction and n-alkane ratios. High sea-level was established around the early/late Valanginian boundary, followed by low sea-level during the late Valanginian, only turning into an increasing trend in the earliest Hauterivian. During the first phase of high sea-level, a change within land plants towards higher chain lengths may be a sign for an increase in environmental moisture availability. Carbonate platform drowning initiated during a time of abrupt rise in sea-level, which was probably an important initial trigger for platform demise. This demise was probably elongated by the late Valanginian low-stand which may have caused platform exposure. The role of an eutrophication for the platform drowning cannot be fully determined by this study.

KeywordsPalynofacies, n-alkanes, sea-level, Early Cretaceous, Valanginian, Vocontian Basin, Carpathian seaway

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5.1 IntroductionThe Valanginian (Early Cretaceous) was characterized by several pronounced climatic and environmental perturbations. These include a severe perturbation of the carbon cycle, globally recorded by a positive excursion of about 2.5 ‰ in δ13Ccarb and δ13Corg records (e.g. Cotillon and Rio, 1984; Lini et al., 1992; Hennig et al., 1999; Wortmann and Weissert, 2000; Weissert and Erba, 2004; Gröcke et al., 2005; Föllmi et al., 2006; Nunn et al., 2010), a cooling interval (Weissert and Lini, 1991; Podlaha et al., 1998; Pucéat et al., 2003; McArthur et al., 2007; Brassell, 2009; Price and Nunn, 2010), maybe accompanied by the formation of high-latitudinal ice-caps (Kemper, 1987; Frakes and Francis, 1988; Price, 1999; McArthur et al., 2007), a “biocalcification crisis” e.g. manifested in a decline in nannoconid abundances and carbonate platform drowning in the Tethyan realm (Wortmann and Weissert, 2000; Erba and Tremolada, 2004; Weissert and Erba, 2004), various turnovers for marine organisms (Mutterlose and Kessels, 2000; Melinte and Mutterlose, 2001; Mutterlose et al., 2003), and changes in pCO2 (e.g. Lini et al., 1992; Price and Mutterlose, 2004; Weissert and Erba, 2004). Major triggers and initial causes are still a matter of debate. The Valanginian was also shown to be a time of sea-level change, whereas it remains questionable whether the major trend was an increasing or decreasing one (Gréselle and Pittet, 2010). Major hiatuses and the occurrence of condensed levels are reported for Valanginian strata e.g. from England, the North Sea, Canada, Mexico, Poland, Germany, and the Volga Basin and have by some authors been interpreted to reflect a major transgression connected to platform drowning (e.g. Föllmi et al., 1994; Bosellini and Morsilli, 1997; Föllmi et al., 2006). Others have interpreted these features contrastingly as being the result of a sea-level fall and the progradation of sedimentary systems (e.g. Vilas et al., 2003; Spalletti et al., 2001). The potential existence of glacio-eustaic control of Valanginian small-scale sea-level fluctuations with amplitudes of more than 50 m has been discussed (Gréselle and Pittet, 2010), in accordance with low rates of sea-floor spreading which thereby were unimportant for eustatic sea-level change (Mutterlose and Kessels, 2000).The reconstruction of sea-level fluctuations is a crucial factor for paleoenvironmental and climatic reconstructions (e.g. Haq et al., 1987). Sea-level changes can have manifold impacts on adjacent continents, e.g. by causing variations in moisture availability due to the enlargement of the area of evaporation or by the flooding or expansion of coastal habitats. Furthermore, sea-level strongly influences marine sedimentation, ocean chemistry, and marine productivity (e.g. Haq et al., 1987). For this study the stratigraphic distribution of palynofacies and n-alkane ratios of two mid-latitudinal sites spanning Valanginian strata is studied with the aim of inferring relative changes in sea-level as well as in terrestrial input at the depositional sites. One site is located in the Vocontian Bain (VB, southeast France), the other one in the Mid-Polish Trough (MPT, central Poland). The site in the MPT was a tipping point during the Valanginian, being located in the Carpathian seaway that connected the Tethys with the Boreal realm (e.g. Kutek et al., 1989). Due to its shallow depth the trespassing of this seaway was only temporal, depending on sea-level, but constantly open during the Valanginian (Melinte and Mutterlose, 2001). The VB site represents a site characterized by relatively high sedimentation rates, while the MPT shows condensed sedimentation for the studied interval (c.f. Kujau et al., in prep.).Palynofacies analysis is defined as the study of the organic facies by transmitted light microscopy (Tyson, 1995). Factors influencing palynofacies include amongst others: terrestrial input, salinity, oxidation, and productivity (Feist-Burkhardt and Götz, 2002). Changes in palynofacies can be interpreted in terms of sea-level fluctuations based on variations in the composition and abundances of marine and terrestrial organic-walled particles (Götz et al., 2009). Biomarkers in the form of n-alkanes can be assigned to marine and

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terrestrial sources and thereby provide information about the origin of the sedimentary organic matter. The dominant sources for n-alkanes are algae, bacteria, and terrigenous plants (Bianchi and Canuel, 2011).Major issues of this study concern the questions if palynofacies investigations serve as a useful tool for Valanginian sea-level reconstructions and if supra-regional patterns in sea-level change affecting the entire Boreal-Tethys realm can be identified. The proportion in the terrestrial palynofacies fraction, in combination with changes in n-alkane ratios, will be used to interpret changes in terrestrial input and thereby in the continental environments. It will, furthermore, be tested if either sea-level fluctuations or eutrophication can be identified as potential triggers for the Valanginian carbonate platform drowning observed in the Tethys.

5.2 Geologic setting and stratigraphy of studied sections 5.2.1 The Mid-Polish Trough (MPT) recordMaterial from a core (PIG1) which was drilled close to the village of Wąwał (E 19°15’0’’, N 52°25’0”) by the Polish Institute of Geology, Warsaw, was studied (Fig. 5.1). Wąwał is located ~4 km southeast to the city of Tomaszów Mazowiecki, and ~115 km southwest to Warsaw. The highly condensed succession has a thickness of ~18 m. Its lithology is composed of clay and claystone with subtle layers of more sandy intervals, shells and sideritic, calcareous and phosphatic nodules. The basal part (~80 cm) consists of limestone, presumably of Berriasian age (Kutek et al., 1989). The core material lacks ammonites of biostratigraphic significance; the age model is based on calcareous nannofossil biostratigraphy and stable carbon isotope chemostratigraphy (Kujau et al., in prep.). Additional stratigraphic control is supplied by lithostratigraphic correlation of the core with a near-by clay pit well dated by ammonites (Kutek et al., 1989; Kaim, 2001). Accordingly, the PIG1 core from Wąwał covers deposits of Berriasian/early Valanginian to late Valanginian age (Fig. 5.2). During the Early Cretaceous, the Carpathian Seaway connected the Tethys and the Boreal Realm via the Lower Saxony Basin and the North Sea Basin (Kutek et al., 1989; Mutterlose, 1992; Kaim, 2001). Its paleo-depth was about 200 m (littoral zone; Rees, 2005). The study site is located in the MPT in the centre of the Carpathian Seaway at a paleolatitude of ~35-40°N (e.g. Blakey, 2010).

Fig. 5.1. Location of sites, A) Polish site (Wąwał) within the Mid-Polish Trough (MPT; modified after Daldez, 2003), B) French site (La Charce, Vergol, Morenas) within the Vocontian Basin (VB; modified after Gréselle, 2007). Black asterisks refer to the studied sections, for other signatures and colour codes see legends within maps.

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5.2.2 The Vocontian Basin (VB) recordThe material studied in the VB is derived from three outcrops (La Charce, Vergol, Morenas) located in the Drome department (Fig. 5.1, 5.2). The Vergol section (E 5°25’9’’, N 44°12’12”) is located between the villages Montbrun-les-Bains and Vergol. Logged and sampled strata (thickness of ~57 m) cover sedimentary rocks from the lower Valanginian (Busnatdoites campylotoxus Zone) to the upper Valanginian (Saynoceras verrucosum Subzone; Gréselle and Pittet, 2010, and references therein). To compensate for a hiatus of slumping, the interval between 25-30 m at the Vergol section has been covered by samples collected at the nearby, lithostratigraphically well correlated Morenas section (E 5°25’23’’, N 44°13’52’’). The La Charce section (E 5°26’23’’, N 44°28’13’’) is well accessible on the hill slopes west of the village of La Charce. Logged and sampled strata (thickness of ~113 m) cover the stratigraphic interval from the lower part of the Upper Valanginian (Karakaschiceras pronecostatum Subzone) to the Lower Hauterivian (Acanthodiscus radiatus Zone; Bulot, 1992; Reboulet and Atrops, 1999). During the Early Cretaceous the VB was located at a paleolatitude of ~30°N in a marginal marine position of the Western Ligurian Tethys Ocean, open to the east. To the north the VB was separated from the Boreal Realm by the archipelago of mid-European continents (Ziegler et al., 1982; Hay et al., 1999; Masse, 1993). Its paleo-depth was about 300 m (Wilpshaar and Leereveld, 1994; Wilpshaar et al., 1997). The central part of the VB is characterized by hemipelagic deposits composed of autochthonous carbonates, allochthonous carbonate fine fraction exported from three surrounding platforms (Fig. 5.1; Reboulet et al., 2003; Gréselle and Pittet, 2010), and terrigenous material, mostly eroded from the Massif Central area (Bréhéret, 1994; Fesneau et al., 2009). The studied deposits are rich in marine invertebrate fossils and consist of well exposed orbitally controlled marl-limestone alternations, stacked in bundles (Gréselle and Pittet, 2010). The composite section has a well-established age model based on ammonites, calcareous nannofossils, and cyclostratigraphy; it covers sediments from the lower Valanginian to lower Hauterivian (Gréselle et al., 2011 and references therein). Age control is also provided by δ13Ccarb chemostratigraphy that has been compared to existing, well dated Tethyan records (Kujau et al., 2012). The composite succession has a total thickness of ~175 m and covers the interval of the lower Valanginian B. campylotoxus to lower Hauterivian A. radiatus ammonite zones (Reboulet et al., 1992; 2003; Reboulet and Atrops, 1999, Gréselle and Pittet, 2010, Fig. 5.2). The MPT and VB records were correlated based on bio- and chemostratigraphy (calcareous nannofossil and stable carbon isotope records (see Kujau et al., in prep.). In the figures presented in this study the records are arranged to each other according to this correlation.

5.3 Methods5.3.1 PalynofaciesA total of 86 samples (40 VB, 46 MPT) was studied quantitatively and qualitatively for its palynofacies by using transmitted light microscopy. At the VB samples were only taken from marly intervals of the marl-limestone alternations. For the MPT clay to claystone was sampled. Preparation of strew slides was done at the Geological Service in Krefeld, Germany, following standard procedures (Traverse, 2007). For plotting the results the software Tilia® was used. For each sample palynofacies counts up to a minimum of 400 particles were done. Sedimentary organic particles were categorized into a marine and a terrestrial fraction. The marine fraction comprises organic walled dinoflagellate cysts (Plate 5.1, Fig. 1, 2, 3), acritarchs (Plate 5.1, Fig. 4, 5), and foraminifera linings (Plate 5.1, Fig. 6). The terrestrial fraction comprises translucent (Plate 5.1, Fig. 7, 8, 9)

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and opaque (Plate 5.1, Fig. 10, 11) phytoclasts, true charcoal (Plate 5.1, Fig. 12), cuticle fragments (Plate 5.1, Fig. 17, 18), and pollen (Plate 5.1, Fig. 13, 14) and spores (Plate 5.1, Fig. 15, 16; e.g. Pittet and Gorin, 1997; Feist-Burkhardt and Götz, 2002). Dinoflagellate cysts and acritarchs are mainly marine cysts and algal bodies, foraminifers are marine protists (Traverse, 2007). Phytoclasts comprise various remains of land plant origin, here separated according to their translucence. Opaque phytoclasts are black or almost black particles that represent wood that lacks any cellular details. Translucent phytoclasts are brown and of either cortex, stem, and root tissue or, in case of more structured particles, additionally of leaf tissue (e.g. Bombaridere and Gorin, 2000; Müller et al., 2006). True charcoal is defined as carbon-rich residue formed by pyrolysis under low oxygen availability, namely carbonized wood or coal (e.g. Enache and Cumming, 2006, and references therein). Cuticles represent the outer layer of epidermal cells of leaves, with a structure revealing single cell outlines (e.g. Boulter and Riddick, 1986). Spores and pollen are the medium of reproduction of non-seed producing land plants and seed-plants, respectively (Traverse, 2007). AOM is in general of a predominantly phytoplankton origin but can as well be formed by degradation of many other palynodebris such as humic particles (e.g. Boulter and Riddick, 1986; Gorin and Steffen, 1991; Feist-Burkhardt and Götz, 2002). For some particles identification was hampered due to fragmentation, orientation in the slide, or masking by pyrite or AOM. These are assigned to the “not identifiable” fraction. Their amount was of max. 33.7 % at the MPT and max. 22.6 % at the VB. Endomycorrhizial fungi were excluded from analysis, since differentiation between recent and ancient ones is problematic (Schrank, 2010). Particles that were counted but not assigned to the marine or terrestrial fraction are amorphous organic matter (AOM; Plate 5.1, Fig. 20, 21) and non-cuticular membranes (Plate 5.1, Fig. 22). AOM is degraded phytoplankton-, bacterial-, or higher plant-resin-derived (Tyson, 1995). Due to their filigrane structure non-cuticular membranes may be of a dominantly marine origin, derived from algae, egg cases, or cyst particles. Land-plant tissue may, however, have been a further potential source (c.f. Batten and Stead, 2005). Due to their non-particulate character and problematic potential origin from various sources they are excluded from interpretation. The composition of phytoclasts including true charcoals is compared based on their amounts in a sample in percent (Fig. 5.3). This is done due to the fact that an enhanced amount of opaque phytoclasts and charcoals compared to translucent phytoclasts may point to an enhanced distance to the shore due a higher potential of preservation of these particles (e.g. Bombardiere and Gorin, 2000; Feist-Burkhardt and Götz, 2002; Fig. 5.3).A marine/terrestrial ratio is formed, based on the amounts of the terrestrial (true charcoals, translucent and opaque phytoclasts, cuticles, spores and pollen) compared to the marine (foram linings, dinoflagellate cysts, acritarchs) palynofacies fraction, with high values indicating a high amount of the marine fraction and vice versa. Here, a high amount of marine particles indicates an enhanced distance to the shore and therewith a higher sea-level (e.g. Bombardiere and Gorin, 2000; Feist-Burkhardt and Götz, 2002; Fig. 5.3).

5.3.2 N-alkanesA total of 36 samples of the VB and 22 of the MPT have been analysed with regard to their n-alkane content. Analyses have been done at the BGR (Federal Institute for Geosciences and Natural Resources), Hannover, Germany. An aliquot of 15-28 g powdered sample material was extracted with dichloromethane using a Dionex accelerated solvent extractor (ASE 200). Elemental sulphur was removed with activated copper. The asphaltenes were separated from the bulk extract by precipitation in petroleum ether. The remaining maltenes and resins fraction was weighed and further fractionated into an aliphatic, an aromatic

Fluctuations in sea-level and terrestrial input at the NW Tethys and the Carpathian seaway during the Valanginian (Early Cretaceous) - Evidence from palynofacies and n-alkanes

95

Fig.

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Chapter 5

96

and a heterocompound fraction using medium-pressure liquid chromatography with solvents of increasing polarity. The aliphatic fraction was then analysed by gas chromatography (Agilent 6890 GC-FID) using a flame ionization detector signal for quantification. The separation of the compounds was achieved with a 60 m DB-1 column (ID 0.32 mm, film thickness 0.2 μm) using a temperature program of 50° (2min)-3°/min-320°(10min) and a constant carrier gas stream of helium of 1 ml/min. N-alkanes can be used as indicators for organic matter sources for marine sediments and changes therein. Their sources are autochthonous in form of algae and bacteria as well as allochthonous in form of terrestrial plant input (Bianchi and Canuel, 2011). Even numbered n-alkanes of nC12 to nC24 are of dominantly bacterial origin. High amounts of odd-numbered nC27, nC29 and nC31 point to terrestrial/cuticular sources. A dominance of nC23-nC25 points to a mainly mixed phototrophic source. High amounts of nC15, nC17 and nC19 point to a dominantly marine algae and phytoplankton source.A ratio in the form of the carbon preference index (CPI) allows to more easily identify the varying dominances of n-alkanes in a sample (e.g. Jeng, 2006; Zech et al., 2008; Bianchi and Canuel, 2011). The CPI ratio uses nine adjacent long-chain n-alkanes (nC24-32), relates their abundances within a sample, and thereby identifies the preference for odd or even n-alkanes (e.g. Handley et al., 2011). A high fraction of land-plant derived n-alkanes has a characteristic odd over even predominance with a CPI above 1. A CPI of 1 would point to no apparent carbon number predominance and to high amounts of marine derived n-alkanes (e.g. Bendle et al., 2007; Jauro et al., 2007). Changes in the nC31/(nC29+nC31) ratio, only including n-alkanes of mainly land-plant origin, point to changes in the dominances of these n-alkanes and thereby in chain-length of these compounds within land-plants which is supposed to change with moisture availability (Schefuss et al., 2003). Due to the fact that for the La Charce section the n-alkanes nC30, 31 are masked by hopanes, steranes, and other abundant compounds with the same weight the ratios can only be calculated for the Vergol and Morenas sections of the VB composite succession, leaving 20 samples for analyses. For the MPT section several different n-alkanes have been masked for distinct samples. The respective samples have been excluded for the records of these ratios.

5.4 Results5.4.1 PalynofaciesPreservation in terms of maturity is good to moderate in all samples and consistent throughout both successions, no sign of carbonization is observed for the palynomorphs. Preservation of particles with regard to fragmentation is in general better for the MPT. Stratigraphic distribution of palynofacies of the MPT can be separated into two major parts (Fig. 5.2, 5.3). The lower part up to 8.9 m is dominated by terrestrial particles, represented essentially by phytoclasts (max. ~90 %). These particles show six intervals of extremely enhanced abundances (at 2.0, 2.9, 3.8, 6.0, 7.6, and 8.8 m) and five of slightly enhanced abundances (at 0.5, 1.5, 4.8, 14.2, and 15.9 m). The marine fraction in this interval is represented with max. 16.7%. After a sharp shift the marine fraction increases up to max. 36.1% while translucent phytoclasts decrease and opaque phytoclasts remain relatively frequent (max. ~60 %). Within this upper interval a high amount of disrupted particles may to a large portion be of dinoflagellate cyst origin. The marine proportion remains prevalent up section, only interrupted between 14.4 to 15.9 m by an interval of reduced amounts of dinoflagellate cysts and foraminifera linings and a slightly increased amount of translucent phytoclasts. Amounts of AOM are in general low (max. 12.9%) and do not show significant changes in abundances. Cuticles are nearly absent (max. 0.7%). Spore-pollen show moderate abundances (avg. 10.5 %) with slightly enhanced abundances

Fluctuations in sea-level and terrestrial input at the NW Tethys and the Carpathian seaway during the Valanginian (Early Cretaceous) - Evidence from palynofacies and n-alkanes

97

between 4.0 to 7.0 m (max. 25.9 %). The occurrence of one acritarch, Leiofusa (Plate 5.1, Fig. 19), was identified but not quantified. Eight samples (at 0.8 m, 1.4 m, 2.2 m, 2.9 m, 6.4 m, 7.5 m, 8.8 m, and 14.3 m) are containing Leiofusa. Its occurrence is correlating with high amounts of translucent phytoclasts (except for the sample at 6.4 m). Enhanced amounts of charcoals (>3.5 %) occur in four samples (at 0.5, 1.8, 5.3, and 15.9 m). The marine/terrestrial ratio shows a sharp turnover from a terrestrial to a marine dominance at 8.9 m. From here onwards a marine dominance is established only interrupted by a phase of slightly reduced values around ~14 m.Palynofacies of the VB is mainly characterized by an alternating dominance of AOM and terrestrial particles in the form of phytoclasts (Fig. 5.2, 5.3). The palynofacies of the lowermost 16.4 m of the VB is dominated by AOM (max. 44.2 %) and non-cuticular membranes (max. 17.4 %). The next interval up to 81.6 m is dominated by AOM (max. 98.0 %) with lowest phytoclasts abundances within the Barrande layers around 40.0 m (<5% and ~10 %, respectively). This is followed by an interval with sharply decreasing AOM abundance, dominated by phytoclasts up to 102.8 m (max. ~47.4 %). This is followed by a short interval of mainly AOM up to 110.9 m (max. 54.1 %). Subsequently, phytoclasts are dominant again up to 147.0 m (max. 49.2 %). The upper part up to the top of the investigated section is again dominated by AOM (max. 58.2 %) and dinoflagellate cysts (max. 27.3 %). Four samples of the VB section evidence an extremely high amount of AOM which covers most of the palynodebris. These are three samples from the so called Barrande layers (at 17.5, 17.8, and 18.8 m; Reboulet, 2001; Reboulet et al., 2003), organic-rich finely laminated and non-bioturbated layers deposited under extreme sea-level conditions (especially high or low) and supposed anoxic conditions (e.g. Gréselle and Pittet, 2010; Westermann et. al., 2010; Kujau et al., 2012). Additionally, one sample from the Morenas section (at 29.7 m) as well shows comparable high amounts of AOM. This sample was sampled off the main sampling sites due to a slumped interval in the main sampled sections which may also point to extreme depositional conditions during the time of accumulation. High cuticle abundances (max. 5.1 %) are restricted to an interval between 131.2 to 147.0 m. Spore-pollen show fluctuating values (avg. 9.3 %) with a general upward decline. One sample (at 12.1 m) is containing Leiofusa. Charcoals are slightly enhanced in one sample (87.6 m, 1.2 %). For the marine/terrestrial ratio values are increasing from ~15.0 m up section, decreasing after ~50.0 m up section, and finally increasing again in the uppermost part of the section, from ~150.0 m up section, followed by a decrease.

5.4.2 N-alkanesFor the MPT, the CPI ranges between 1.2 and 2.5 and shows a shallow increase up to 6.0 m (CPI of max. 2.5, Fig. 5.4), stable conditions up to 11 m, a decrease up to 13.3 m (CPI of min. 1.6), another increase up to 16 m (CPI of max. 1.9) and subsequently lower values again (CPI of min. 1.2). Values for the nC31/(nC29+nC31) ratio vary between 0.3 and 0.4 and show a decreasing trend until 9.5 m, followed by an increasing trend. For the VB the CPI shows constantly low values (below 1.5). Values for the nC31/(nC29+nC31) ratio vary between ~0.6 and 0.7, with initially comparatively low values (min. ~0.6), an increase between 29.7-39.1 m (max. ~0.6) followed by a phase of comparatively stable values, and another increase after 56.0 m (max. ~0.7). The Barrande layers (at 17.5, 17.6 [this one was not analyzed for palynofacies], 17.8, and 18.8 m; Reboulet, 2001; Reboulet et al., 2003) evidence enhanced values compared to background values (all of ~0.6).

Chapter 5

98

Fig. 5.4. Biostratigraphy, thickness, lithology, composition of section (only VB) compared to n-alkane ratios CPI and nC31/(nC29+nC31) for the VB (left, excluding the La Charce section) and the MPT (right). Blue boxes mark phases of high sea-level. For stratigraphic legend see Fig. 5.2. B1-4 mark position of Barrande layers.

5.5 Discussion5.5.1 Interpretations of sea-level changes based on palynofacies characteristicsAccumulation and preservation of particulate organic matter in sediments depends on biomass production, degradation, and transport (e.g. Gorin and Steffen, 1991, and references therein). Changes in the palynofacies can be used to gain information about changes in paleoenvironmental conditions during accumulation of the organic matter (Tyson, 1995). Sea-level fluctuations strongly affect the position of the shoreline, which in turn has a strong impact on the amount of the marine versus continental fraction in the sedimentary deposits (Gorin and Steffen, 1991; Feist-Burkhardt and Götz, 2002). For the MPT phases of high dinoflagellate cyst and acritarch abundances associated with low amounts of the terrestrial fraction are interpreted as phases of high sea-level (e.g. Bombardiere and Gorin, 1998; Gorin and Steffen, 1991; Feist-Burkhardt and Götz, 2002; blue boxes in Fig. 5.2-5.4), prevailing between ~9.0 to 14.0 m, and between ~16.0 and 17.5 m. The low portion of translucent phytoclasts for both intervals affirms the interpretation, due to their potentially lower potential of preservation. They are of higher abundance closer to the shoreline (Fig. 5.3; Gorin and Steffen, 1991; Feist-Burkhardt and Götz, 2002). A relative sea-level lowstand hence can be interpreted for the lower half of the section, with high amounts of translucent phytoclasts. Overall low amounts of AOM for the MPT indicate a low sedimentation rate (Gorin and Steffen, 1991), consistent with a condensed sedimentation.For the VB phases of high AOM amounts and low amounts of translucent phytoclasts are used as an indication

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Fluctuations in sea-level and terrestrial input at the NW Tethys and the Carpathian seaway during the Valanginian (Early Cretaceous) - Evidence from palynofacies and n-alkanes

99

for phases of high sea-level (c.f. Batten and Stead, 2005, blue boxes in Fig. 5.2-5.4). A high fraction of AOM counted during these phases results in low abundances of the marine fraction which should otherwise be expected to be of high abundances under high sea-level. High amounts of AOM do in this case not serve as an indication of reducing conditions in bottom waters, since anoxic conditions have been excluded for the VB (Kujau et al., 2012). They do instead probably reflect high sedimentation rates and rapid organic matter burial (Gorin and Steffen, 1991). This is in accordance to the observation that sedimentation rates at the VB were about ten times higher compared to the MPT. Accordingly, phases of high sea-level are located between ~17.0 and ~70.5 m, around ~107.0 m and from ~147.0 m up section (Fig. 5.2, 5.3). Due to their specific depositional conditions Barrande layers are excluded from sea-level interpretations. Comparison of the sea-level patterns deduced from palynofacies records from the two sites clearly shows similar trends for the overlapping interval. At both sites rapid trend towards a higher sea-level before the early/late Valanginian boundary and high sea-level around this boundary is followed by a short phase of lower sea-level in the early late Valanginian. This is at both sites followed by another phase of high sea-level.

5.5.2 Terrestrial input and paleoenvironmental conditionsIn addition to sea-level interpretations the composition of terrestrial palynomorphs can be used to identify changes in paleoenvironmental conditions for adjacent continents. The terrestrial fraction of the palynofacies of the VB is of minor importance compared to the marine fraction, which is consistent with an organic-geochemical study by Kujau et al. (2012) who identified the OM of the VB as phytoplankton dominated. Strong fluctuations in the abundances of translucent phytoclasts at the MPT reflect most probably local variations in the input of land-plant derived debris, e.g. during extreme rainfall events (c.f. Batten and Stead, 2005). On the other hand, slightly enhanced amounts of charcoals at the VB and the MPT may point to at least seasonally dry conditions that favoured wildfires during these intervals (Hesselbo et al., 2003). An increasing amount of cuticles for the upper interval of the VB, before the initiation of the last interval of rising sea-level, may point to short distance transport, since these particles are comparatively large. Their large size in combination with an enhanced weight may cause reduced buoyancy and thereby a deposition in short distance to the detritus source in form of river deltas, or a high-energy transport (c.f. Batten and Stead, 2005). An enhanced occurrence of these particles may point to the progradation of deltas towards the site of deposition, consistent with low sea-level and/or an increase in humidity in the hinterland combined with high-energy fluviatil transport (e.g. Boulter and Riddick, 1986). The acritarch Leiofusa was referred to as “disaster species” e.g. by van de Schootbrugge et al. (2007). Like this, it can probably be interpreted as being an opportunist and thereby of selective advantage under extreme environmental conditions like eutrophication under high terrestrial input. This, and the fact that it is a dominant marine component under low sea-level in several samples of the MPT with high amounts of phytoclasts and in one sample of the VB with high amounts of AOM, may point to recurrent hostile marine conditions. A low sea-level may have caused sluggish conditions in combination with an enhanced terrestrial nutrient input. This may have caused unfavourable eutrophic conditions for other marine biota. Another scenario may be an increased transport of these particles from terrestrial sites in form of lakes and swamps under humid conditions and elevated fluviatile discharge, which may have transported Leiofusa towards the depositional sites. Biomarkers in the form on n-alkanes provide further information about specific OM sources and thereby about changes in terrestrial input (Fig. 5.4; e.g. Meyers, 1997). Barrande layers are excluded from interpretation since their different composition for n-alkanes compared to the rest of the samples is probably due to a

Chapter 5

100

better preservation caused by anoxic conditions during deposition (Westermann et al., 2010; Kujau et al., 2012). For the MPT, a comparatively high CPI points to a certain land plant input for the lower half of the section and to a reduction therein or dilution by other n-alkane sources from ~11.0 m up section, consistent with an increasing sea-level. The discrepancy between the indicated rise in sea-level around ~9 m based on palynofacies interpretations and the decrease in land-plant input followed with a delay may indicate that the distance to the shoreline remained constant during the initial phase of the transgression with a still high land-plant input. Biomarker are then still easily transported towards the site of deposition, while the input of terrestrial organic particles may be already have been reduced due to decay in the water column or more widespread transport and dispersal by marine currents. Another phase of a slightly increased CPI values between ~14.0 and 16.0 m corresponds to another phase of lowered sea-level. Overall low CPI values for the studied lower interval of the VB section point to only limited land plant input (e.g. Jeng and Huh, 2006; Bendle et al., 2007; Jauro et al., 2007), consistent with a high sea-level. Regional precipitation is supposed to have a strong influence on the chain length distribution of leaf-wax lipids (Schefuss et al., 2003 and references therein). For the MPT the trend in the nC31/(nC29+nC31) ratio is initially decreasing and gradually shifts to more positive values after ~9.50 m. In the VB record, the nC31/(nC29+nC31) ratio shows a first shift to longer-chain nC31-dominated n-alkanes above ~35.0 m and another positive shift after ~55.0 m. This overall trend towards a higher nC31 content may be caused by a replacement of floras within the vegetation around the sites (Zdravkov et al., 2011). A shift to longer chain n-alkanes within the land-plant fraction may as well point to an intensification of precipitation under an intensifying monsoon, since chain-length variations in land plants are probably correlated to moisture availability, with longer chains indicating more humid conditions (Schefuss et al., 2003; Bendle et al., 2007). This would point to an enhanced moisture availability, probably due to an intensification of monsoonal precipitations for both sites, for the VB especially after ~55.0 m.

5.5.3 Comparison to existing interpretations of Valanginian sea-level fluctuationsThe section of the VB, presented here, has already been interpreted in terms of sea-level changes based on a sequence stratigraphic approach coupled with cyclostratigraphic interpretations by Gréselle and Pittet (2010, Fig. 5.3). Accordingly, a sea-level fall occurred before the early/late Valanginian boundary, accompanied by the progradation of fluvio-deltaic systems. An initially low sea-level is as well identified by the here presented data, associated with indications for enhanced terrestrial nutrient influx. In both interpretations this is followed by a sea-level rise. This rise in sea-level may be explained by a thermal expansion of water (e.g. Handley et al., 2011) under a rise in temperatures indicated for this interval (e.g. McArthur et al., 2007). The phase of low sea-level (according to this study) between ~70.0 and 150.0 m for the VB would be consistent with a cooling interval (following Pucéat et al., 2003; McArthur et al., 2007; Barbarin et al., 2012), only interrupted by a short phase of enhanced sea level around ~105 m. Small-scale fluctuations reported for the interval between ~60.0 and ~110.0 m, interpreted as glacio-eustatic changes (Gréselle and Pittet, 2010) are not congruently reflected in the palynofacies data. A short-term increase at ~90.0 m proposed by sequence stratigraphy is not observed in the palynofacies data, which rather indicate a regressive trend due to the dominance of land-plant derived debris. This discrepancy may be due to the fact that for the facies analysis used for the sequence stratigraphic approach the depositional environment may be over-interpreted based on the absence of distinct taxa. They may be missing due to eutrophication or cooling rather than sea-level fluctuations. A sea-level lowstand proposed for the late Valanginian (at ~120.0 m) is as well reflected in the

Fluctuations in sea-level and terrestrial input at the NW Tethys and the Carpathian seaway during the Valanginian (Early Cretaceous) - Evidence from palynofacies and n-alkanes

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palynofacies data, it is however not as abruptly followed by increasing sea-levels as proposed by Gréselle and Pittet (2010).The here presented sea-level record is compared to a global sea-level record adopted from Mutterlose and Bornemann (2000), based on Haq et al. (1987) and recalibrated after Gradstein et al. (1994, Fig. 5.3). The increasing trend around the early/late Valanginian boundary is consistent with both, the VB and MPT records. A highstand proposed for the entire late Valanginian is however neither in accordance with this study nor with Gréselle and Pittet (2010). This may be due to local factors influencing the sea-level records. Highest Valanginian sea-level was proposed for the S. verrucosum ammonite zone by Melinte and Mutterlose (2001), based on a “boreal nannofossil excursion” reported for Rumania in this zone. The evolution of Boreal and Tethyan taxa was reported to have been highly endemic during the late Jurassic, with a changing trend towards a high mixing during the Valanginian. A first exchange of ammonite taxa in both directions in northwestern Europe occurred in the S. verrucosum zone under a high sea-level. For belemnites only a migration of Tethyan taxa northwards is reported, not the other way round (Melinte and Mutterlose, 2001). Jansen and Clément (2002) reported an extinction event in belemnites during the S. verrucosum zone for the VB, explained by an increased Boreal influence and a replacement by new taxa, especially Boreal cephalopods. This interpretation of high sea-level around the early/late Valanginian boundary is in accordance with palynofacies data. For the VB very low abundances of phytoclasts occur during the interval of the S. verrucosum zone, probably as well pointing to especially high sea-level.

5.5.4 Platform drowning – eutrophication or sea-level?The entire phase of the late Valanginian was interpreted as a phase of carbonate platform crisis (Gréselle and Pittet, 2010 and references therein). Platform drowning is supposed to initiate around the deposition of the Barrande layers (c.f. Fig. 5.2-5.4, following Gréselle et al., 2011). Low sea-level, which would have exposed the platforms, establishes not before the late Valanginian in the uppermost S. verrucosum zone (after ~70.0 m). The initial phase of platform drowning after Barrande layer deposition is, however, accompanied by high sea-level at both sites which may have drowned the platforms. Eutrophication was supposed as another potential cause for the observed Tethyan platform drowning, induced by increased weathering and continental runoff under enhanced pCO2 (Fölmi et al., 1994; Lini et al., 1992; Weissert et al., 1998). Both factors, sea-level fluctuations and eutrophication, could also have accelerated each other (van de Schootbrugge et al., 2003). No distinct changes in terrestrial input based on of the palynofacies data can be observed during the phase of high sea-level at the two studied sites to approve the theory of eutrophication under an enhanced terrestrial input. N-alkanes may point to a change in the vegetation or to high precipitation intensities under an intensifying monsoonal climate after the early/late Valanginian boundary. At the VB acritarchs are of highest abundance at the same time (until ~65.0 m). This may be due to sea-level rise but could as well be an indication of enhanced pCO2 and thereby caused acritarch blooms (van de Schootbrugge et al., 2007). The question if platform drowning in the Tethys was caused by sea-level or eutrophication can not be entirely solved by this study. A rise in sea-level may have been an important initial trigger under rising temperatures and thereby expanding waters. Subsequently, an exposure of platforms under low sea-level may have elongated the crisis in platforms.

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5.6 ConclusionsIt was shown that palynofacies investigations serve as a useful tool for investigating Valanginian relative sea-level fluctuations. An early Valanginian phase of low sea-level was followed by a sea-level rise around the early/late Valanginian boundary and probably a short phase of high-stand during the S. verrucosum ammonite zone. This is followed by a late Valanginian low sea-level, only turning into an increasing trend at the earliest Hauterivian. Both investigated sites, in the VB and the Carpathian seaway show comparable trends. Before the first rise in sea-level terrestrial nutrient input was probably enhanced. During times of high sea-level changes in terrestrial input based on palynofacies are hardly seen. N-alkanes point to a change in vegetation structure during the phase of high sea-level around the early/late Valanginian boundary probably due to an increase in humidity. Before the last phase of increasing sea-level at the VB a progradation of river deltas may have occurred. Platform drowning was probably initially caused by a rapid sea-level rise around the early/late Valanginian boundary and elongated during the late Valanginian caused by platform exposure under low sea-level. The role of nutrient influx for platform drowning cannot be estimated by this study.

AcknowledgementsThanks are due to Izabela Ploch, Polish Institute of Geology, Warsaw, Poland, for providing the possibility of sampling the Polish site, and to Benjamin Gréselle for assistance in the field at the French site. Further thanks are due to Christian-Ostertag Henning, Georg Scheeder, and Monika Weiss of BGR (Federal Institute for Geosciences and Natural Resources), Hannover, Germany, for the help with biomarker analyses. Financial support from the DFG project HE4467/2-1 is gratefully acknowledged.

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Plate 5.IExamples of palynofacies particles with indicated sample in that the respective pictures were taken. Latin numbers according to pictures on plates, indicating one (to three) example(s) of particles.

Marine fraction:1 Dinoflagellate cyst PIG1.1502 Dinoflagellate cyst PIG1.1823 Dinoflagellate cyst PIG1.1604 Acritarch LC1665 Acritarch PIG1.1826 Foraminifera lining PIG1.182

Terrestrial fraction:7 Translucent phytoclast PIG1.1828 Translucent phytoclast PIG1.659 Translucent phytoclast PIG1.15010 Opaque phytoclast PIG1.18211 Opaque phytoclast PIG1.18212 True charcoal PIG1.5613 Pollen PIG1.5614 Pollen PIG1.1315 Spore PIG1.17416 Spore PIG1.3517 Cuticle LC16618 Cuticle LC166

Others:19 Leiofusa PIG1.7820 AOM VER5621 AOM PIG1.15022 Non-cuticular membrane LC166

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6. Linking changes in pCO2 with environmental and climate dynamics during the Valanginian (Early Cretaceous)

AbstractMajor triggers for a Valanginian (Early Cretaceous) positive carbon isotope excursion (CIE), reflecting a global perturbation of the carbon cycle are a matter of ongoing debate. It has been previously proposed that this positive CIE was caused by oceanic anoxia leading to enhanced storage of organic matter (OM) in the marine realm and thereby causing the isotope shift in the active carbon reservoir since OM is 13C-depleted. There is, however, no evidence for widespread marine anoxia. Other potential triggers like continental carbon storage or carbonate platform demise have been considered, but the causes and consequences for this carbon cycle perturbation are still questionable. This study provides new insights into the leads and lags of climatic and environmental alterations that accompanied the Valanginian CIE. Carbon isotope records (δ13Corg and δ13Ccarb) are used to reconstruct coeval changes in the carbon cycle and trends in pCO2 (partial pressure CO2) in form of a ∆δ13C record. Compound specific measurements and RockEval pyrolysis provide a control of the reliability of the δ13Corg bulk record. The established ∆δ13C record represents the first high-resolution measure of Valanginian pCO2 fluctuations accompanying the positive CIE. A correlation of terrestrial environmental moisture levels with the CIE allows for the first time the investigation of coeval changes in terrestrial and marine environments based on stratigraphically well constraint sections. A humid phase indicated by vegetation structure occurs coeval to a major positive shift in the δ13C record. This interval is accompanied by a pCO2 drawdown and cooling, and may point to terrestrial carbon storage as a major trigger for the Valanginian CIE. According to the timing of events changes in pCO2 may have been caused system-intrinsically, since the occurrence of enhanced volcanic activity is still debated.

KeywordsValanginian CIE, pCO2, Δδ13C, pCO2, compound specific isotopes, Vocontian Basin

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6.1 IntroductionThe causes and consequences of a major perturbation in the carbon cycle during the Valanginian (~144.5-133.9 Ma; Ogg and Ogg., 2008), globally recorded in a severe positive carbon isotope excursion (CIE) of about 2.5‰, are still enigmatic (e.g. Cotillon and Rio, 1984; Lini et al., 1992; Hennig et al., 1999; Wortmann and Weissert, 2000; Weissert and Erba, 2004; Gröcke et al., 2005; Föllmi et al., 2006; Nunn et al., 2010). The Valanginian positive CIE represents the first of several Cretaceous carbon isotope anomalies. Its duration is estimated to be in the order of ~3 Ma (Gréselle et al., 2011), which is comparatively long in respect to other Cretaceous CIEs (Jarvis et al., 2011; Föllmi, 2012). Commonly, similar Cretaceous CIEs are interpreted to result from short-termed enhanced sequestration of organic matter (OM) in deep ocean basins or marginal settings. This OM sequestration is postulated to be caused by enhanced marine productivity and/or preservation probably under the establishment of anoxia (e.g. Nunn et al., 2010; Jarvis et al., 2011 and references therein). The Valanginian lacks evidence for widespread anoxic conditions and enhanced sequestration of OM in the marine realm (e.g. van de Schootbrugge et al., 2003; Reboulet et al., 2003; Erba et al., 2004; Gröcke et al., 2005; Westermann et al., 2010; Kujau et al., 2012). It was, however, a time of distinct paleoenvironmental and climatic changes, including fluctuations in the amount of atmospheric pCO2 (partial pressure CO2) supposed to be connected to volcanic activity (e.g. Lini et al., 1992; Price and Mutterlose, 2004; Weissert and Erba, 2004), an acceleration of the hydrologic cycle and therewith changes in terrestrial input into ocean margins (Weissert, 1989; Föllmi et al., 1994; Weissert et al., 1998; Weissert and Erba, 2004; Föllmi et al., 2006), a phase of cooling (e.g. Weissert and Lini, 1991; Podlaha et al., 1998; Pucéat et al., 2003; McArthur et al., 2007; Brassell, 2009), and a “biocalcification crisis” expressed by widespread carbonate platform demise and a decline in nannoconids (Wortmann and Weissert, 2000; Erba and Tremolada, 2004; Weissert and Erba, 2004; Barbarin et al., 2012). A detailed assessment of qualitative changes in pCO2 across the Valanginian CIE and its linkage to climatic and environmental change on high resolution is to date missing. In general, the level of pCO2 during the Early Cretaceous is supposed to have ranged between ~560-1500 ppm, compared to pre-industrial levels of ~285 ppm (Berner, 1994; Haworth et al., 2005). During the Valanginian volcanic activity may have introduced CO2 into the active carbon reservoir, causing global warming, enhanced weathering and runoff. Enhanced nutrient availablitiy and submarine volcanism could have enhanced primary production and carbon storage, thereby causing the positive CIE (Lini et al., 1992). Supposed potential volcanic sources for CO2 include the Paranà-Etendeka trapps (Renne et al., 1992; Weissert et al., 1998; Erba et al., 2004) and the Comei-Bunbury Large Igneous Province (Zhu et al., 2009). The co-occurrence of enhanced volcanic activity with the positive shift has, however, been called into question (e.g. Barbarin et al., 2012). The exact timing of the Paranà-Etendeka trapps (~134.7 ± 1 Ma) and the Comei-Bunbury Large Igneous Province (~132 Ma), formerly associated with the Valanginian CIE, is to date unknown, significant volcanic activity was probably post-dating the Valanginian CIE (Gibson et al., 2006; Thiede and Vasconcelos, 2010; Barbarin et al., 2012 and references therein). In this study, we use carbonate and organic carbon isotope records (δ13Ccarb, δ

13Corg) to reconstruct changes in the Valanginian carbon cycle and to provide a qualitative measure for changes in paleo-atmospheric pCO2. RockEval pyrolysis is used to decipher the composition of sedimentary OM, which could affect its isotopic signature. Compound specific carbon isotope analysis on isoprenoids and a supposed land plant derived n-alkane further back up the reliability of trends in the bulk δ13Corg record. The estimated pCO2 change is

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compared with existing records on changes in marine carbonate production, temperature, and vegetation and deduced terrestrial moisture levels. A concurrence of massive changes in vegetation structure and fluctuations in the global carbon cycle has been detected for several time intervals like the end-Permian mass extinction and the Triassic-Jurassic transition (Bonis et al., 2009; Larsson, 2009; Hermann et al., 2011). In contrast, the interplay of Valanginian terrestrial environments with the carbon cycle perturbation was not yet investigated. The potential role of the vegetation as an active modulator of the Valanginian carbon cycle, via terrestrial storage of carbon under enhanced humidity, will be discussed. This approach provides a new perspective on and a deeper understanding of triggers for this CIE.

6.2 Geologic settingFor this study an outcrop composite section located in the central part of the Vocontian Basin (southeast France, Fig. 6.1) has been chosen as sedimentary archive. Logged and sampled sections form a composite succession including La Charce, Vergol, and Morenas (Fig. 6.2); for details see Gréselle et al. (2011) and Kujau et al. (in prep.). Only the marly interbeds of marl-limestone alternations have been sampled to assure consistency within samples. Stratigraphic framework is based on biostratigraphy (ammonites and calcareous nannofossils) and chemostratigraphy (δ13Ccarb, Gréselle et al., 2011 and references therein; Kujau et al., in prep.). During the Early Cretaceous the VB was a marginal marine basin located in the northwestern Tethys, to the north separated from the Boreal realm by the archipelagic mid-European continents (Blakey, 2010).

Fig. 6.1. Location of sites (La Charce, Vergol, Morenas) within the Vocontian Basin (modified after Gréselle, 2007) marked by black asterisks. For colour codes see legend.

6.3 Material and MethodsThe δ13Ccarb record is adopted from Kujau et al. (in prep.), measured on the here presented material (295 samples). For the δ13Corg record measurements have been carried out on the organic fraction of 104 powdered bulk rock samples (~0.6 mg). An elemental analyser (Thermo scientific Trace GC Ultra) connected to an isotope ratio mass spectrometer (Finnigan Mat delta S via a Thermo scientific Finnigan GC Combustion III), available at the Ruhr-University Bochum, Germany, have been used for stable isotope analyses. Repeated analyses of a certified standard (USGS24) and graphite (NBS-21) show an external reproducibility of ± 0.09 ‰. Values are expressed in conventional delta notation, relative to the Vienna Pee Dee Formation belemnite (VPDB) international standard, in per mil. (‰). Duplicate measurements (24) show that the samples are homogenous (max. dev. δ13C ± 0.21).Additionally, carbon isotope measurements have been accomplished on specific compounds of the apolar

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campylotoxus

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Fig. 6.2. Chemostratigraphic correlation of δ13Ccarb and δ13Corg records with significant intervals marked with grey bars. Records are form left to right from the Vocontian Basin in southeast France (Kujau et al., 2010 and this study; given with sub-stanges, ammonite zones, nannofossil zones, and thickness), the Lombardian Basin in northern Italy (Lini et al., 1992), and the Gulf of Mexico (Site 535; mbsf=meters below sea floor; Wortmann and Weissert, 2000). Correlation with DSDP site 535 is hampered by the low resolution of the carbon isotope records, indicated by broken bars.

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fraction of the extracted bitumen (14 samples), including the isoprenoids pristane and phytane, as well as the n-alkane nC29. Extraction of the bitumen was performed at the BGR (Federal Institute for Geosciences and Natural Resources, Hannover, Germany). An aliquot of 15-28 g powdered sample material was extracted with dichloromethane using a Dionex accelerated solvent extractor (ASE 200). Elemental sulfur was removed with activated copper. The asphaltenes were separated from the bulk extract by precipitation in petroleum ether. The remaining maltene and resins fraction was weighed and further fractionated into an aliphatic, an aromatic and a heterocompound fraction using medium-pressure liquid chromatography with solvents of increasing polarity. Compound specific measurements were accomplished at the NIOZ (Royal Netherlands Institute for Sea Research, Texel, the Netherlands). Measurements were repeated twice to three times, on a gas chromatograph connected to a mass spectrometer (GC-MS). A Finnigan Delta C isotope-ratio-monitoring gas chromatography mass spectrometer was used equipped with a fuse silica capillary column coated with CP-Sil5, with helium used as a carrier gas. The oven was programmed at a starting (injection) temperature of 70°C, which rose to 130°C at 20°/min and then to 320° at 4°/min, at which it was maintained for 20 min. The δ13C values are reported in standard delta notation against VPDB with an average reproducibility of 0.5‰. Even though low response and coelution with compounds of the same weight hampers the identification of compounds and thereby causes a high variance of the δ13C values major trends in isotope records can be identified. For determining the type and thermal maturity of the organic matter (OM) and to assess whether migrated bitumen is present, samples with TOC ≥0.2% have been analyzed using a RockEvalTM6 at the BGR Hannover, Germany. For the Vergol and Morenas sections this has already been done by Kujau et al. (2012). Here, 18 additional samples of the La Charce section have been analyzed. The parameters quantified during pyrolysis include S1, S2, S3, Tmax, PI, HI, and OI (following Espitalié et al., 1977). The S1 and S2 peaks correspond to the maximum of hydrocarbon generation in mg/gsed, with S1 representing hydrocarbons volatilized from the powdered material and S2 representing hydrocarbons mainly formed during thermal cracking of the remaining kerogen. S3 corresponds to the amount of CO2 generated from the OM. Tmax (in °C) is the temperature of maximum rate of hydrocarbon generation during pyrolysis. Production index (PI) is calculated from S1/(S2+S2). Hydrogen index (HI) is given in mg HC/g TOC and oxygen index (OI) in mg CO2/g TOC.

6.4.1 The significance of the ∆δ13C record for reconstructing trends in Valanginian pCO2

The δ13Corg bulk signals vary between 26.1 and 28.4 ‰ and show a two-step positive CIE of about 2.5 ‰ (Fig. 6.2). The record does not show an offset between the top to bottom parts of different sections of the composite record. Hence, a composite δ13Corg curve can be established. In the NK3A nannofossil zone (0.0 to 32.0 m) a first increase to more positive values from ~-28.5 to ~-27.5 ‰ can be observed. In the lowest NK3B zone (32.0 to 47.0 m) values are stagnant around -27.7 ‰. A second increase up to ~-26.5 ‰ occurs within the lower to middle NK3B zone (47.0 to 74.0 m). Values remain high with maximum values up to ~-26.2 ‰ at around the transition of zone NK3B to NC4 (90.0 to 120.0 m). The observed trend in the δ13Corg record mimics the pattern visible in the δ13Ccarb record. The shape in the records is, however, not completely congruent, some intervals are significantly deviating (Fig. 6.2). These deviations are an initially less pronounced increase to more positive values in the δ13Corg record (lower grey bar, ~9-33 m), followed with delay by an increase in the δ13Ccarb record after a short negative shift in δ13Ccarb. This negative shift (at ~26 m) is not recorded in the δ13Corg record, maybe due to lower resolution

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of the record. The increase in values is followed by a decrease for the δ13Corg record (~-27.8 ‰) and a phase of relatively stable values for the δ13Ccarb record (33-39 m δ13Ccarb, 33-47 m δ13Corg). A second, more pronounced increase to more positive values in the δ13Ccarb record is reflected in the δ13Corg record following a short temporal lag (upper grey bar). Subsequently, the δ13Ccarb record shows a plateau-phase (up to ~107 m), followed by a decrease, while the δ13Corg record is further increasing (max ~-26.0 ‰ at ~93 m). Reconstruction of trends in paleo-atmospheric pCO2 is based on calculation of the ∆δ13C record. The ∆δ13C is calculated in the form δ13Ccarb - δ

13Corg, with both values measured on bulk material (e.g. Kump and Arthur, 1999; Jarvis et al., 2011). The applicability of this record as a proxy for changes in pCO2 is based on the fact that δ13Ccarb mirrors the surface ocean carbon reservoir and thereby the 12C/13C composition of sea-water, whereas δ13Corg reflects the isotopic composition of marine phytoplankton (Lini et al., 1992; Pagani et al., 1999). Photosynthetic fractionation by phytoplankton increases under an increase in CO2 (Dean et al., 1986). This explains for discrepancies occurring between δ13Ccarb and δ13Corg measured on stratigraphic records during intervals in Earth history when atmospheric and thereby oceanic pCO2 was changing (e.g. Heimhofer et al., 2004; Jarvis et al., 2011). The isotopic composition of organic carbon is however, furthermore influenced by several other factors. These include changes in the surface ocean carbon reservoir (changing abundances of HCO3-, CO2[aq], H2CO3, CO3

2-) variations in the isotopic composition of dissolved CO2 (CO2[aq]) metabolized by phytoplankton (Lini et al., 1992), and changes in growth rate of phytoplankton which is increasing under an increase in paleoproductivity and thereby causing the same trend in δ13Corg like decreasing CO2 (e.g. Pagani et al., 1999; Heimhofer et al., 2004). The latter highlights the importance of information on paleo-productivity. Variations in the contribution of terrestrial versus marine organic matter to the bulk OM also have a severe influence on the δ13Corg values (Lini et al., 1992), due to the fact that terrestrial OM shows a less negative δ13C composition in comparison to marine OM (Jarvis et al., 2011). If it is not about a calculation of absolute values, the reconstruction of trends in pCO2 from bulk material is, however, regarded as reliable as long as the composition of OM remains stable (e.g. Jarvis et al., 2011).Samples have only been taken from the same substrate represented by well consolidated marls which ensures consistency of the measured type of lithology. The source of the carbonate within the marls was shown to be of dominantly calcareous nannofossil origin (Gréselle et al., 2011). This assures that the δ13Ccarb record is not biased by the contribution of differing δ13C signatures from different carbonate precipitating organisms. The OM was shown to be of predominantly phytoplankton origin for the Vergol and Morenas sections, supporting the reliability of the δ13Corg bulk record (Kujau et al., 2012). RockEval pyrolysis data for the lower half of the composite section (up to 62.40 m; Vergol and Morenas; Kujau et al., 2012) plotted in a van Krevelen diagram reveal that the OM of the analyzed samples can be assigned to marine type II kerogen, close to type III, with low to moderate maturity (Fig. 6.3; HI above ~170 mg HC/g TOC,, Tmax of avg. 435°C Lüniger and Schwark, 2002). RockEval pyrolysis for the upper half of the composite section (La Charce) exhibits HI values between 133.9 and 296.6 mg HC/g TOC and OI values between 44.8 and 271.4 mg CO2/g TOC. Tmax values vary between 425° and 436° (avg. 430°C). Low PI values (avg. 0.02) indicate that no migrated bitumen is present in the investigated samples (e.g. Schwark et al., 2009). Plotted in a van Krevelen diagram the OM of the La Charce section can as well be assigned to marine type II kerogen, close to type III, with low to moderate maturity. Here, the maturity of the OM is probably lower compared to the lower part of the composite section seen in lower Tmax and HI values. To test the reliability of trends in the δ13Corg bulk record compound specific δ13C measurements were

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accomplished on two biomarkers of assumable marine phytoplankton origin, pristane and phytane, and one derived from land-plants, nC29 (e.g. Bianchi and Canuel 2011; Fig. 6.4). Low compound response to high background noise of non-identified compounds with similar weights, like steranes and hopanes, hampered compound identification of the nC29 n-alkane at the La Charce section and results in high variations of measured isotope values, which therefore are excluded. Values for the marine phytoplankton markers fluctuate between -32.0 to -36.0 ‰, values for the land-plant derived compound vary between -31.5 to -34 ‰. Up to ~25 m the marine biomarker show an opposing trend to the terrestrial one. Subsequently (~30 m), the δ13C trend is decreasing, which is less clear for the n-alkane, followed by an increase in values (~37 m), which is again less clear for the n-alkane. This is followed (~42 m) by a decrease for all records. The n-alkane record stops, while pristane and

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Fig. 6.3. Van Krevelen diagram depicts samples as type III kerogen with medium maturity. Samples derived from the Barrande layers correspond to white dots, the Vergol/Morenas section samples are represented by white dots (Kujau et al., 2012), samples from the La Charce section are shown as white quaders.

phytane show a general shift to more positive values, only decreasing with the last measured sample (~80 m). Thereby the records show a general similarity with the δ13Corg bulk record and further prove its reliability. Exceptions are the negative trend around 30 m, and the decreasing trend reflected in the last samples of the marine biomarkers which are not reflected in the δ13Corg bulk record. The negative trend around 30 m, however, mimics a preceding negative trend in δ13Ccarb. The data set of bulk δ13C records is thereby identified as most suitable for pCO2 reconstructions, with constraints for the La Charce section of a lower maturity and a higher admixture of various compounds (high background noise) hampering compound specific measurements, reflected in high fluctuations in values of repeated measurements of δ13C values for pristane and phytane. N-alkane identification for compound specific measurements is not possible due to masking by other compounds for the La Charce section.

6.4.2 Comparison with existing carbon isotope recordsFor the Valanginian CIE, only few δ13Corg records are available to date. However, the deviating pattern of the δ13Ccarb and δ13Corg record indicating changes in pCO2 was also observed at other sites for which both, δ13Corg and δ13Ccarb, records are available e.g., the Lombardian Basin in northern Italy (Lini et al., 1992) and DSDP Site 535 in the Gulf of Mexico (Wortmann and Weissert, 2000; Fig. 6.2). The comparatively low resolution of the δ13Corg records of these sites hampers a high-resolution correlation with the here presented secords. But an overall consistency in the trends of the δ13Ccarb and δ13Corg records as well as in the deviation between the different isotopic substrates proves that the observed patterns as reflect a true paleoenvironmental signal that can be interpreted for pCO2 reconstructions. An initial increase towards more positive values in δ13Corg (lower grey bar) is followed with a certain delay by an increase in δ13Ccarb (no data available for DSDP site 535). A subsequent decrease in δ13Corg values is accompanied by stable δ13Ccarb records (no data available for DSDP site 535). A second, more pronounced increase to less negative values in δ13Corg is visible in all three records (upper grey bar) and shows a temporal delay compared to the second positive shift in δ13Ccarb. For all presented records the positive shift amounts

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to ~2.5‰ except for the δ13Corg record at DSDP site 535, where a shift of only ~2.0 ‰ has been measured. The following phase of decreasing δ13Ccarb values is not paralleled by δ13Corg. Since in the Vocontian Basin the upper half of the δ13Corg record may be biased by the maturity and composition of the OM (terrestrial/marine, high abundances of compounds masking the n-alkanes, high fluctuations in duplicate measurements) this part will be excluded from interpretations of changes in pCO2. The interpretations will focus on the initial phase of the Valanginian CIE. Trends in pCO2 based on the Δδ13C record represent as follows: PCO2 is decreasing until ~26 m (from to ~29.2 to 28 ‰), followed by increasing values up to ~48 m (to ~29.7 ‰), followed by a decreasing trend up to ~100 m (to ~28 ‰). Subsequently the record shows highly fluctuating values on a low level (around ~28.2 ‰; Fig. 6.4).A change in Valanginian pCO2 has long been discussed as an important trigger for the CIE (e.g. Lini et al., 1992; Wormann and Weissert, 2000; Erba and Tremolada, 2004). So far, no high resolution estimation has

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been accomplished due to the lack of appropriate δ13Corg records. According to Gröcke et al. (2005) there is no indication for a shift in Valanginian pCO2 based on Δδ13C records before the CIE and during its peak for marine records at DSDP site 535 (Patton et al., 1984), Polaveno, and Rio Corna (Lini et al., 1992). In contrast, a land-plant record from Crimea studied by Gröcke et al. (2005) revealed a shift of 2 ‰ in the Δδ13C record and has been interpret as a drop in pCO2 from the pre-peak to peak CIE. The problem of the data set from Crimea is twofold: (i) Due to the absence of non-altered carbonate in the Crimean succession the Δδ13C record has been calculated with the help of δ13Ccarb data adopted from Tethyan marine composite records. (ii) Changes in the photosynthetic fractionation as a function of pCO2 are well established for marine photosynthetic organisms (e.g. Pagani et al., 1999), but not for land plants, which have a better control on incoming pCO2 (e.g. O’Leary, 1988). Hence, the use of the 13Corg signature derived from fossil wood may cause a certain bias in the resulting Δδ trend. In contrast to existing data sets the δ13Ccarb and δ13Corg records of this study are of high resolution and well correlatable since they are measured from the same samples. Furthermore, the OM, at least for the initial phase of the CIE, is of a dominantly phytoplankton origin. A drop of ~2 ‰ in the Δδ identified by Gröcke et al. (2005) is consistent with the here presented data, but in contrast to the Crimean data set it only establishes after the peak in pCO2 not before the peak phase. Based on paleofluxes of nannofossils Erba and Tremolada (2004) assume a 2-3 times increase in pCO2 during the Valanginian CIE, which would be in accordance to a proposed increase for this interval based on this study.

6.5.1 Climate settings during the initiation of the CIEValanginian humid phases are reflected in vegetation structure changes, here shown in form of a spore-pollen ratio (Kujau et al., in prep.; Fig. 6.5). A gradual trend to more humid conditions for the lower Valanginian, indicated for a site in the Carpathian seaway (central Poland) probably indicates an intensification of monsoonal moisture availability for southern Paleo-Europe (Kujau et al., in prep.). A subsequent phase of maximum humidity after the early/late Valanginian boundary is also reflected at the Vocontian Basin and is coinciding with the second pronounced increase in the δ13C records. It is probably caused by most extreme monsoonal climates (Kujau et al., in prep. and references therein). Factors like a positive vegetation-precipitation feedback via albedo (e.g. Ganopolsi et al., 1998) and/or an enhanced solar output (e.g. Vaughan, 2007) may have further intensified moisture availability. A northward shift of the Inner Tropical Convergence Zone (ITCZ), controlled by changes in the orbital parameters obliquity and precision, may have additionally enhanced solar insolation for the northern hemisphere (NH) and thereby potentially elongated the growing season by a seasonally longer lasting humid period (Kujau et al., in prep.).A short-termed drying, reflected at the Vocontian Basin before supra-regionally extreme humid conditions are reached, may be explained by a temporal northward shift and/or expansion of the NH arid belt under descending air masses of expanding Hadley cells (c.f. Hasegawa et al., 2010). This may have influenced moisture availability for the hinterland of the Vocontian Basin, which then would have annually longer been under the influence of dry air masses. The phase of supra-regionally humid conditions would then be reflecting a southward shift of this arid belt again. In current meteorological literature mechanisms determining the latitudinal position of the descending arms of Hadley cells are still described as an unknown (e.g. Roedel and Wagner, 2011). However, recent paleoclimatological studies suggest that an expansion of the Hadley cells is probably to an important degree controlled by the amount of atmospheric pCO2 and temperature and thereby the energetic level of the atmosphere (Brierly et al., 2009; Reichler, 2009; Hasegawa

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et al., 2011). Accordingly, the Hadley cells would be expanding poleward under high temperature and/or enhanced pCO2, like they do today (e.g. Hu and Fu, 2007; Sheerwood et al., 2010). But the cells seem to react non-linear. Under even more enhanced pCO2 (pCO2 ~1000-1500 ppm) and/or temperature of “super greenhouse” conditions of a certain threshold the cells would show an opposite trend and be rapidly shrinking equatorward again (Fig. 6.6, Hasegawa et al., 2011). Considering changes in Valanginian pCO2 this mechanism probably played an important role for Valanginian climate settings of the mid-latitudes.

VBVB

+Temp./CO2 ++Temp./CO2

Hadley cells

Northern hemispherearid belt extent

Fig. 6.6. Schematic globes with circulation of Hadley cells (black lines with arrows) and poleward expansion of arid belts (as black lines). Left: Hadley cells expanded under high temperature and/or atmospheric pCO2, arid belts located further polewards. Right: Hadley cells shrunk under even more elevated temperature and/or pCO2. Asterisk marks position of the Vocontian Basin. Not to scale.

6.5.2 Causes and consequences of changes in Valanginian pCO2

It has been noted earlier that coupled changes in both, the marine and terrestrial organic and carbonate carbon system may have contributed to CIEs (e.g. Weissert et al., 1998; Erba et al., 1999). This is to be tested for the Valanginian CIE. A scenario that explains the positive Valanginian CIE beyond marine OM storage was e.g. suggested by van de Schootbrugge et al. (2000), Price and Mutterlose (2004), and Westermann et al. (2010) since evidence for enhanced marine OM storage is missing (e.g. Weissert et al., 1998; Weissert and Erba, 2004; Westermann et al., 2010). The burial of organic carbon in terrestrial environments has long been identified as an important controlling factor for the gas composition of the atmosphere (Berner and Canfield, 1989). Westermann et al. (2010) propose a combination of OM storage in marginal marine seas and on continents coupled with carbonate platform demise to explain the positive isotope shift. Since organic carbon is depleted in 13C, a higher amount of 13C would thereby be left in the active carbon reservoir and 13C-enrichment in the atmosphere and the oceanic dissolved inorganic carbon reservoir reservoir would be the consequence (Arthur et al., 1985; Scholle and Arthur 1980). The records of pCO2 and environmental and climatic change around the initiation of the CIE are described based on segmentation into four specific intervals, indicated before each description (Fig. 6.4, 6.5). Thereby the question of causes and consequences of the CIE is concerned. The La Charce section is excluded from pCO2 interpretations due to its problematic constitution of the OM and high variance of measured δ13C values. For comparison with changes in temperature, three mid-latitudinal sea surface temperature reconstructions have been chosen, including that of Pucéat et al. (2003, based on oxygen isotopes of fish tooth enamels), McArthur et al. (2007, based on oxygen isotopes and Mg/Ca ratios of belemnites), and of Barbarin et al. (2012, as well based on oxygen isotopes of fish tooth enamels). For comparison with carbonate production a CaCO3 % record for the lower part of the section from Kujau et al. (2012) as well as

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a record of nannoconid absolute abundance of Barbarin et al. (2012) and an estimation of the time interval of carbonate platform demise following Gréselle et al. (2011) are shown. Changing moisture levels are indicated based on high abundances of spores in the floral composition of the vegetation shown in a spore-pollen ratio (Kujau et al., in prep.).

(I) 0-25 m: Comparatively stable δ13Ccarb values indicate that no severe changes in the 12C/13C ratio of the seawater occurred. Carbonate production and nannoconid absolute abundance are as well stable after an increasing trend in CaCO3% and a decrease in nannoconid abundance at the very beginning of the zone. Platform demise is supposed to initiate towards the end of the zone (according to Gréselle et al., 2011) contemporaneous to a decrease for δ13Ccarb pointing to environmental stress for carbonate producing organisms. The bulk δ13Corg record evidences a 1st positive shift. Accordingly, the pCO2 record (∆δ13C) shows a decreasing trend. A first of three possible scenarios for explaining this may be (i) a sequestration of 12C out of the active carbon cycle by marine or terrestrial carbon storage, leaving heavier values to the rest of the system and lowering atmospheric pCO2. There is, however, no evidence for severely enhanced marine carbon sequestration for this interval except for a few centrimetric layers (Erba et al., 2004), in the Vocontian Basin expressed by the so-called Barrande layers, most probably deposited during a short phase of temporary dys- to anoxic conditions (B1-4, c.f. Kujau et al., 2012 and references therein). Terrestrial environmental moisture conditions are increasing at a site located in the Carpathian seaway (central Poland; Kujau et al., in prep.) northeast of the Vocontian Basin. At the Vocontian Basin, however, the decreasing trend in moisture availability does not support a widespread terrestrial carbon sequestration. Moreover, in case of carbon sequestration, both, the carbonate and organic carbon system would be expected to react with a positive shift. (ii) An alternative explanation may be that in this interval the ∆δ13C record does not reflect changes in pCO2 but rather a change in sea surface productivity. An increased nutrient influx may have caused enhanced phytoplankton fractionation. Sea-surface waters then would have become 12C-depleted, leading to heavier isotope values for phytoplankton, leaving the carbonate carbon system relatively unaffected. Nutrient influx was, however, not especially enhanced during this interval probably except for the short interval of Barrande layer accumulation (Kujau et al., 2012 and references therein). (iii) A further scenario is that under high temperatures fractionation between HCO3- and CO2[aq] is decreasing, causing an increase in phytoplankton δ13C (Méhay et al., 2009 and references therein). Temperatures are indeed high during this interval (Pucéat et al., 2003; McArthur et al., 2007; Barbarin et al., 2012). The latter scenario or a combination of the last two can probably be assumed as the most plausible explanation, and putative changes in pCO2 seen in Δδ13C as not entirely reliable. For biomarkers the 1st shift in the bulk Corg records is only reflected in nC29. This is a further evidence against severe changes in atmospheric pCO2, on that the marine Corg system would probably have reacted homogenously. High temperatures (induced by orbital forcing) were probably also responsible for the initiation of the northward expansion of the NH arid belt by expanding Hadley cells causing a drying around the Vocontian Basin.

(II) 25-33 m: An abrupt short-termed negative shift of about 0.5 ‰ is seen in δ13Ccarb, with a short time lag also reflected in pristane and phytane. This negative shift was also recorded from other Tethyan (e.g. Lini et al., 1992; Erba et al., 2004; McArthur et al., 2007; Gréselle et al., 2011) and non-Tethyan sites (e.g. Cotillon and Rio, 1984) proving its reliability as a paleo-environmental signal. It is, however, not reflected in bulk δ13Corg. The negative shift may have been caused by an abrupt release of marine methane clathtrades into the

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active carbon cycle (CH4 has light isotope values of ~-60‰) destabilized under high temperatures (Jahren et al., 2001; Beerling and Berner, 2002; Galli et al., 2005; Retallack and Jahren, 2008; Misumi et al., 2009). A here recorded shift in 0.5 ‰ may reflect an input of ~800 Gt carbon in form of CH4, after oxidation leading to an increase of max. 100 ppm in pCO2 (c.f. Beerling and Berner, 2002). PCO2 is subsequently abruptly increasing. The negative shift is probably too short and abrupt for being induced by pCO2 release from volcanic activity, also 13C-depleted (c.f. Jahren et al., 2001; McElwain et al., 2005). The abrupt increase in δ13Ccarb following the negative shift can probably be explained by a “response effect” of the carbonate carbon system to the disturbance and elevated stress under an abrupt increase in pCO2, a thereby caused ocean acidification, and/or an enhanced nutrient influx from continents, and may have caused a shelf ecosystem change from oligotrophic to eutrophic and platform demise (Wortmann and Weissert, 2000). Carbonate production is indeed decreasing and carbonate platform demise continues. Evidence for an increase in nutrients comes e.g. from an increase in high-fertility nannofossil taxa (Erba and Tremolada, 2004; Duchamp-Alphonse et al., 2007; Barbarin et al., 2012), a shift from autotrophic to heterotrophic producers (Géselle et al., 2011 and references therein), high phosphorous accumulation rates (van de Schootbrugge et al., 2003), enhanced kaolinite contents (Fesneau, 2008) and accumulation of Mn and Fe (Kuhn et al., 2005). Arid conditions are established at the Vocontian Basin (Kujau et al., in prep.). The increase in pCO2 may have contributed to the expansion of Hadley cells, generating a more widespread establishment of arid conditions. The δ13Corg record shifts to more positive values consistent with an increase in pCO2.

(III) 33-47 m: This interval of most rapid increase in pCO2 is coeval to stagnant to decreasing values for the δ13C records, confirmed by biomarker evidence. Following Gröcke et al. (2005) an increase in ∆δ13C of ~1.5‰, seen here, points to an increase in pCO2 of ~30%. This calculation is based on plant material but is in rough agreement with pCO2 calculations of Jarvis et al. (2011) for the Cenomanian-Turonian oceanic anoxic event, also based on bulk measurements. According to the decreasing to stable δ13C records the introduced CO2 must have been 13C-depleted. The source of CO2 may have been a combination of the oxidation of terrestrial and marine 13C-depleted OM e.g. by (seasonally enhanced) wildfires and drying, a degassing of swamplands etc., and changes in ocean circulation, under high temperatures (e.g. Zachos et al., 2001 and references therein). An overturning of the Tethys may have served as an OM source by oxidation of marine OM-rich deposits (McElwain et al., 2005). Carbonate production is most rapidly decreasing towards the end of this zone and thereby probably contributed to the pCO2 increase. There is no confirmed evidence for volcanic activity peaking around the Valanginian CIE (e.g. Barbarin et al., 2012 and references therein). Volcanic activity as a source for CO2 can, however, not be neglected. There is evidence for potential Valanginian volcanic activity along the northern active margin of the Tethys around the early/late Valanginian boundary based on volcanic ash (Fesneau et al., 2009), or in arctic Canada (Emby, 1988). This needs further investigation. The preceding and following processes around the pCO2 increase may support a non-volcanic scenario where an increase in pCO2 is system-intrinsically controlled by terrestrial and marine OM oxidation under high temperatures.An inversion of the trend in Hadley cell expansion probably occurred under high temperatures and rising pCO2 based on the assumption of a non-linear reaction of the cells on an increase in pCO2 and temperature. Accordingly, this may point to an increase in pCO2 above at least 1000 ppm (Hasegawa et al., 2011). Consequently, the NH arid belt was probably retreating equatorward, causing the initiation of a humid

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trend at the Vocontian Basin. A connection of drastic shrinking of the Hadley cells and extreme humid conditions for the mid-latitudes was also shown for the mid-Cretaceous (Hasegawa et al., 2011). The abrupt increase in δ13Ccarb following the stable to decreasing values can probably again be explained by a “response effect” of the carbonate carbon system in form of the carbonate producing community to the disturbance by acidifying CO2 and nutrient input and elevated stress (c.f. Weissert and Erba, 2004).

(IV) 47 m upsection: An abrupt increase in δ13Corg follows that in δ13Ccarb. It is probably caused by terrestrial OM storage and 12C sequestration under the widespread, abrupt establishment of especially humid conditions, reflected in the spore-pollen record, coeval to one at the Carpathian seaway (Kujau et al., in prep.). This probably further promoted the increase in δ13Ccarb. Under extreme humid conditions, probably caused by most intense monsoonal climates with enhanced terrestrial Corg burial and silicate weathering, a negative feedback loop turns increasing pCO2 into a rapid drawdown. A transition from an increasing to a decreasing trend in pCO2 can cause a temporary further increase in global average precipitation (Wu et al., 2010), promoting humid conditions. This 2nd positive shift in δ13Corg bulk records is affirmed by a similar two-fold increase in pristane and phytane. Contemporaneously, a severe transient cooling occurs reflected in all three presented temperature records, consistent with the drawdown in pCO2 and a potential terrestrial Corg-storage. The return to pre-CIE δ13C values only establishes in the lower Hauterivian (Sprovieri et al., 2006). The interval covered by the La Charce section is not further interpreted in terms of pCO2 changes due to the problematic composition of the OM probably affecting the bulk δ13C. The indicated decreasing trend in pCO2 is, however, probably reliable to some extent, consistent with a return to pre-CIE conditions.

6.5.3 Implicati ons for triggers of the Valanginian CIEThe fact that plants as terrestrial primary producers are able to store carbon (preferentially 12C) makes them an active part of the short-term global carbon cycle. An increased availability of moisture furthermore increases discrimination against 13C and therefore causes plants taking up even more of the lighter carbon isotope 12C (Farquhar et al., 1989). Terrestrial carbon storage is enhanced under humid conditions by a more dense vegetation cover and a thereby enhanced production of biomass. Under humid conditions carbon is stored on continents by preventing plant organic matter from being decayed in swamps etc. in a first stage, occasionally followed by coal formation (when precipitation>evaporation; Ziegler et al., 1987), affecting the long-term carbon cycle. Year-round continuity of rainfall is the most important factor in forest growth, swamp preservation, and therewith coal formation (Ziegler et al., 1987), which is given during the here described interval of most extreme humid conditions and most pronounced shifts in the δ13C records (beginning segment IV). This study shows for the first time a congruent change in vegetation structure and derived environmental moisture levels and a severe positive shift in the carbon isotope records of the Valanginian. The 2nd pronounced positive shift in δ13Ccarb and δ13Corg was probably mainly triggered by terrestrial carbon storage. Contrastingly, the 1st positive shift in δ13Ccarb can probably be explained by a “response effect” of the carbonate system after methane release, and the 1st less pronounced shift in δ13Corg by decreasing fractionation between HCO3- and CO2[aq] under high temperatures in combination with nutrient influx and phytoplankton productivity. Thereby the importance of terrestrial changes for the Valanginian long-term carbon cycle is pointed out. Evidence for coals of potential Valanginian age was given by studies of Budyko et al. (1987) and McCabe and Parrish (1992) who assume coastal coal formation and a maximum of Corg storage on continents for the

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Valanginian. Valanginian terrestrial coals are e.g. known from low latitudes and sites in Asia and Greenland (e.g. Ziegler et al., 1987; Rees Mc Allister, 2004). In general, mid- and high latitudes were identified as sites of coal deposition and maximum plant productivity (Rees McAllister et al., 2004). Exact dating of terrestrial coals is, however, hard to provide which hampers a direct correlation to δ13C records. The extraction of 12C from the active carbon cycle resulting in pCO2 drawdown (probably accelerated by silicate weathering) by terrestrial carbon storage in humid mid- and high-latitudes is, however, a logical consequence of the here described scenario. During this time interval of pCO2 drawdown, humid conditions were probably expressed more widespread, and probably expanded more equatorward (retread of arid belt). Thereby, a more widespread swamp formation and potential terrestrial carbon storage could have occurred. A larger area would then be affected by vegetation restructuring, enhanced discrimination, and enhanced biomass production. Besides the well correlation of the pronounced isotope shift with the spore-pollen ratio, the interval around the positive CIE was shown to be a time of enhanced swamp habitat occurrence (Kujau et al., in prep.), which further supports the idea of an enhanced terrestrial Corg storage. A link between a strengthened NH summer monsoon and positive shifts in the δ13Corg records, which can be assumed for the time interval of the Valanginian CIE, was also shown for small scale monsoonal fluctuations during the Last Interglacial (Yong et al., 2002). The two-step shift to more positive carbon isotope values coincides with a phase of high sea-level (Gréselle and Pittet, 2010). This may have been a further factor for carbon sequestration on marine shelfs and in epicontinental seas. Another location for terrestrial Corg storage may have been large rift lakes between South America and Africa (Chang, 1991; Goncalves, 2001).

6.6 ConclusionsFor this study a site of the mid-latitudes of Early Cretaceous Europe, located in the Vocontian Basin (southeast France), was investigated regarding carbon cycling and changes in pCO2, in correlation with terrestrial moisture levels, and marine carbonate production. Enhanced humidity was shown to be parallel to a major positive shift in the δ13C records, pointing to enhanced terrestrial OM storage, probably causing an identified pCO2 drawdown and a contemporaneous cooling. Therewith, continental environments can be supposed to have played a pivotal role for observed Valanginian changes in the global carbon cycle reflected in a pronounced positive CIE. A severe increase in pCO2 (of ~30 %) accompanies the initial phase of the Valanginian CIE, affirming its importance for recorded environmental and climatic perturbations associated with this carbon cycle event.

AcknowledgementsThanks are due to Benjamin Gréselle for his assistance in the field. Great thanks are also due to Christian Ostertag-Henning, Georg Scheeder, Monika Weiß, and Annegret Tietjen for their help with biomarker extraction, and RockEval measurements and to Stefan Schouten and Jort Ossebaar for their help with specific isotope measurements. Financial support from the DFG project HE4467/2-1 is gratefully acknowledged.

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7. Synthesis and future perspectives

7.1 The applicability of the chosen approachIn this study, climatic and environmental dynamics that accompanied a Valanginian carbon cycle perturbation, reflected in a positive carbon isotope excursion (CIE), were investigated based on a multi-proxy approach. An assembled high-resolution carbonate carbon isotope record of the Vocontian Basin (SE France) reflects the already well established positive excursion initiating during the lower/upper Valanginian boundary interval. It was possible to further assemble a δ13C record based on bulk organic matter (OM) on the very same sample material, which is of special interest regarding the scarcity of these kind of records for the Valanginian. Since bulk OM bears the problematic that various factors can bias its carbon isotope values it was especially usefull that it was possible to prove the paleoenvironmental significance of the δ13Corg record by geochemical analyses and compound specific isotope measurements, at least over the initial phase of the CIE. These anaylses in the form of biomarker investigations and RockEval pyrolysis revealed that the bulk organic matter is of mainly phytoplankton origin and therewith its trends are reliable. This was taken advantage of for the establishment of a record on qualitative changes in Valanginian atmonspheric pCO2 over this event. Changes in Valanginan terrestrial plant community structures were investigated for both investigated mid-latitude sites (Vocontian Basin, SE France; Mid-Polish Trough, central Poland) and allowed for a comparison of supra-regional similarities as well as specfic local characteristics of the vegetation. Based on this, changes in moisture and climate were deduced. Palynofacies investigations for both sites were used to reconstruct changes in sea-level. Additionally, biomarkers in the form of n-alkanes were investigated to further reveal trends in terrestrial input. By the chosen approach it was possible to gain a most complex insight into various paleoenvironmental and climatic dynamics that were of importance for a better understanding of causes and consequences of the Valanginian CIE. Unfortunately, the compound specific measurements were hampered by the occurrence of various compounds of “background noise” and variations in the maturity for the different sites. But still, for the important interval of the initiation of the CIE, trends could be established.

7.2 A marine or terrestrial trigger for the Valanginian CIE?The investigation of biomarkers and OM content for the Vocontian Basin challenges the already questioned occurrence of enhanced Valanginian marine OM burial and a thereby caused 13C-depleted carbon sequestration as a major trigger for the observed positive shift in the carbon isotope records. The investigation of one single basin of course provides only a local view on changes in marine conditions. But due to the fact that this basin did indeed serve as a site of enhanced OM accumulation during younger Cretaceous OAEs and in regard of a generally observed absence of enhanced OM storage on a global scale, the observation of this basin probably did reveal a general truth. Enhanced terrestrial nutrient influx may have caused a temporary increased marine productivity and may even have caused enhanced OM burial in restricted basins. This probably contributed to the isotope shift. A widespread storage of carbon in the marine realm as a major cause for the Valanginian CIE becomes, however, more and more unlikely. A special focus of this study was on the terrestrial realm and its potential as a major sink for carbon. Based

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on the chosen approach of a palynological investigation of plant communities this can of course only be answered indirectly. A direct approach would be the identification of terrestrial coal bearing archives and their correlation with the CIE, which is hampered by the problematic dating of terrestrial archives, especially coals. But indeed, a major change in the plant community structures indicating enhanced humidity can be well correlated with a major increase in both, the organic and carbonate carbon isotope records. The occurrence of widespread humid conditions can be assumed, at least for the mid-latitudes, were the investigated sites were located in. This is based on various indications for the occurrence of swamp habitats, the absence of xerophytic taxa, and high fern spore abundances. The establishment of especially humid conditions was interpreted as being potentially caused by the influence of a monsoonal circulation on the studied sites of southern Paleo-Europe that may have been temporary most intense during the major shift in carbon isotopes causing most intense humidity. N-alkane ratios that indicate changes in chain-lengths of land-plants, which can be interpreted in form of changes in moisture, provide a further hint for an increase in humidity around the lower/upper Valanginian boundary, which is the interval of the initiation of the positive CIE. This increase in humidity can then be indirectly assumed to have caused a widespread formation of swamps, ponds, and mires that store OM and may occassionally have formed coal deposits. Sea-level was high during the initial phase of the CIE, indicated by palynofacies data, which probably enlarged the area of evaporation around the continents and thereby may have been an important factor for an increased humidity. High sea-level may furthermore have caused the formation of epicontinental seas and the flooding of coastal habitats, both further factors for an enhanced carbon storage. This scenario provides the possibility of an enhanced medium- and long-term storage of carbon in the terrestrial realm during the Valanginian and thereby highlights the high potential of terrestrial sites as major or at least severe sinks for 13C-depleted carbon.

7.3 Fluctuations in the atmospheric carbon contentBased on a qualitative measure of trends in fluctuations of atmospheric pCO2 it was possible to estimate changes therein over the Valanginian CIE. A severe increase in pCO2 was identified to accompany the isotope shift and can thereby be interpreted as an important controlling factor for the various recorded paleoenvironmental alterations. It may, for example, have increased temperatures and thereby may have accelerated the hydrologic cycle. Furthermore, an elevated amount of atmospheric pCO2 seems to be an important controlling factor for atmospheric circulation systems like the Hadley cells, which react non-linear to atmospheric pCO2 by an expansion or shrinking. These effects then have a profound impact on environmental change, too, by controlling the location of arid belts. A change in these belts was probably identified in this study by the phase of arid conditions that preceded the supra-regional especially humid conditions at the Vocontian Basin. Even if the source for the introduced CO2 remains enigmatic it is possible to speculate that even here the terrestrial realm may have played an important role via the oxidation of OM. A drawdown in pCO2 can, like mentioned before, however, most probably be directly assigned to an enhanced terrestrial storage of OM under especially humid conditions and high sea-level, in combination with an enhanced silicate weathering.

7.4 Future perspectivesThis study can be understood as a first step towards a more complete understanding of the Valanginian CIE that includes the role of the terrestrial realm. This, of course, provides information that can help in general

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to better understand environmental perturbations of this kind, like other intervals in Earth’s history that revealed comparable shifts in the isotope records but lack evidence for a widespread carbon sequestration in the marine realm. Regarding future climate change it is clear that the complex interactions and feedback loops within the atmosphere-ocean-biosphere system do not allow for simple predictions of consequences of a continuous fossil carbon oxidation and associated global warming. Regarding the non-complete understanding of the operation mode of the carbon cycle and its interaction with climate and the environment, ongoing research is needed. A more complete estimation of sources and sinks for carbon in oceans and in the terrestrial realm is still in progress. Paleoenvironmental and climatic studies are still a most valuable tool for solving unanswered questions. The terrestrial realm should become a more important part of this research, which can further be concluded from this study. It was mentioned that the Early Cretaceous received less attention regarding changes in vegetation patterns and floral compositions of different climate belts, compared to younger Cretaceous time intervals. This is remarkable regarding the fact that the rise of angiosperms, which fastly became the most important line in plant evolution, was initiated in this time interval and which is still not fully understood. Further investigations of Valanginian sediments should therefore include palynological approaches, as far as possible.It should have become clear that general conclusions regarding changes in large-scale climate patterns like the monsoonal circulation or the Hadley cell expansion would profit from a more dense data base in form of a higher number of investigated sites to better understand changes of these large scale phenomenons, reflected in vegetation change. More globally distributed sites would provide a more general picture. Of course this is hampered by the scarcity in Valanginian sediments, large hiatuses, and condensed sedimentation. Worth a try would be the search for and investigation of rare continental archives like ancient lake sediments, paleosoils or, like referred to in this study, the further investigation of the varying location of eolian dunes. A further interesting task would be the quantitative estimation of changes in Valangininan pCO2 which would help to further assess sources and sinks for carbon.

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Acknowledgements

First of all I would like to give my deepest thanks to my supervisors Ulrich Heimhofer and Jörg Mutterlose for giving me the opportunity to write this thesis. Thank you for all the help through the steps of data collection, understanding, and finally writing this up, and for continuously keeping me on track. Not the least, I am very thankful to you for giving me the opportunity to attend several very interesting and important conferences and summer schools. Thank you both very much for believing in me, for all your support, and all the things I learned from you!!

I further would like to very much thank Peter A. Hochuli for introducing me into the field of Cretaceous Palynology. Stefan Schouten for hosting me in his laboratory at the northwestern edge of the Netherlands, Jort Ossebaar for sharing his office and his knowledge on biomarkers there with me. Benjamin Gréselle for his assistance in the field in France and ongoing enthusiasm for the Valanginian. Christian Ostertag-Henning, Georg Scheeder, Monika Weiß, and Annegret Tietjen from the BGR in Hannover for their support with my biomarker studies. Izabela Ploch for giving me the opportunity to sample a core in Wąwał, Poland, and for being such an especially nice host. Thierry Adatte and Chloé Morales who went to Poland with me. Christoph Hartkopf-Fröder from the Geological Survey NRW, Germany, for preparing the palynology samples and providing me an insight into his lab. Other scientists that were important to me during my PhD thesis in one way or the other and that I would like to give my thanks to are (especially!!) Adrian Immenhauser, Silke Voigt, Helmi Weissert, Stéphane Bodin, Stéphane Westermann, André Bahr, René Hoffmann, Dieter Buhl, Niels Rameil, and Rolf Neuser.

Special thanks still go to the supervisors of my diploma thesis, written years ago, who made me believe in the first place that I could be someone doing a PhD, Dirk Nürnberg and Christoph Zielhofer.

Great thanks are due to my dear PhD (+ pre- and post-) fellows for all the fun, experiences, beers, thoughts, and coffees we shared within these years in Bochum: Jasper, Nico, Anthony, Susanne, Stefan, FranÇois, Mélanie, Baris, Sylvia, Sabine, Christian, Sebastian, Anna, Sara, Agnes, Nils, Jean, Andrea, Dana, Rute, Wawa, Steffi, and Melody.I would like to very much thank our secretaries Conny Mell and Sabine Sitter for always helping me out when needed. And thanks are also due to Beate Gehnen and Ulrike Schulte for their help in the lab in Bochum.

Biggest thanks go to my family for always supporting me in following my interests, and for their continuous encouragement and support, and their never ending believe in me. My parents Ilse and Sandor, my sister Britta with Rainer, and my brother Daniel with Kathrin. At this point I would also like to give my warmest thanks to Rita, Fritz, and Lisa Wiesmann for their ongoing interest and motivation. Too many people left during these years of writing, among them my beloved aunt and grandmothers Karin, Aleida, and Waltraut. But I know you would like what I am doing here.

I am also gratefully thankful to my friends for accompanying me during this time of my life. Especially Annika, Kerstin, Julia, Jacek, Nicklas, Priggi, Astrid with Ingo and Janosch, Jojo, Axel, Jana, Michi, Jan-Rainer, Ludger, Jan, Wolle, Heiko, Alex, Felix, Alev, Jens, Janne, and Robbe. Even if you sometimes did not give a damn about geosciences that may especially have helped me through this. And I would like to thank Stefan R.

Most thank of all is due to the most important person in my life, Hanno Wiesmann, for always being there for me, whatever stage I am in (…). I love you.

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AppendixCorrected values, carbon and oxygen isotopes:sample /d13CVPDB(‰)/d18OVPDB(‰)/duplicatesVER001 0,4 -1,3 VER002 0,8 -1,3 VER003 0,4 -1,3 VER004 1,0 -1,3 VER005 0,9 -1,1 VER006 0,7 -0,8 VER007 0,5 -1,1 VER008 0,7 -1,1 VER009 0,6 -1,0 VER010 0,6 -0,9 VER011 0,8 -1,1 VER012 0,7 -0,9 VER013 0,8 -1,2 VER014 0,7 -1,0 VER015 0,7 -1,1 VER016 0,6 -0,9 VER017 0,6 -1,4 VER018 0,8 -0,9 VER019 0,8 -1,2 VER020 0,9 -1,3 VER021 1,0 -1,4 VER022 0,4 -1,1 VER023 0,9 -1,6 VER024 0,7 -2,1 VER025 0,7 -2,1 VER026 0,9 -2,1 VER027 0,4 -1,5 VER028 0,6 -1,3 VER029 0,7 -2,0 VER030 0,8 -1,2 VER031 0,7 -1,2 VER032 0,8 -1,2 VER033 0,6 -1,6 VER034 0,6 -1,1 VER035 0,6 -1,4 VER036 0,6 -1,0 VER037 0,5 -1,7 VER038 0,5 -1,8 VER039 0,1 -1,9 0,3 - 1 , 8VER040 0,5 -1,5 MOR001 0,7 -1,5 MOR002 0,7 -1,2 MOR003 0,9 -1,4 0,9 - 1 , 4MOR004 1,0 -2,0 MOR005 1,2 -1,0 MOR006 1,1 -0,9 MOR007 1,1 -1,9 MOR008 1,0 -1,7 MOR009 1,1 -1,7 1,1 - 1 , 7MOR010 1,1 -1,8 VER041 1,2 -1,3 VER042 1,4 -1,3 VER043 1,6 -1,7 VER044 1,5 -1,4 VER045 1,7 -1,8 VER046 1,3 -1,5 VER047 1,7 -1,8 VER048 1,4 -1,9 VER049 1,5 -1,8 VER050 1,4 -1,8 VER051 1,4 -1,6 VER052 1,6 -1,7 VER053 1,2 -1,6 VER054 1,5 -1,5 1,5 - 1 , 5VER055 1,3 -1,5 VER056 1,7 -1,8 VER057 1,6 -1,9 VER058 1,5 -1,6 VER059 1,7 -1,8 VER060 1,5 -1,4 1,5 - 1 , 4VER061 1,7 -1,9 VER062 1,7 -1,7 1,7 - 1 , 7VER063 1,5 -1,6 1,6 - 1 , 6VER064 1,9 -1,8 VER065 2,1 -2,0 VER066 2,0 -1,6 VER067 2,2 -2,1 VER068 2,3 -1,8

Corrected values, carbon and oxygen isotopes:sample /d13CVPDB(‰)/d18OVPDB(‰)/duplicatesVER069 2,4 -1,7 VER070 2,5 -1,8 VER071 2,5 -2,0 VER072 2,3 -2,0 VER073 2,4 -1,9 VER074 2,3 -2,0 VER075 2,3 -1,7 VER076 2,2 -1,7 VER077 2,1 -1,8 VER078 1,9 -2,2 VER079 2,1 -2,0 VER080 2,0 -1,9 VER081 2,4 -2,1 VER082 2,2 -1,8 VER083 2,3 -2,0 VER084 2,5 -2,0 VER085 2,1 -1,9 VER086 2,5 -2,1 VER087 2,6 -1,8 LC001 2,0 -0,8 LC002 1,7 -1,0 1,7 - 1 , 0LC003 1,9 -1,2 LC004 2,4 -1,1 LC005 2,0 -1,2 LC006 2,0 -1,2 LC007 1,7 -1,0 LC008 2,6 -1,2 LC009 2,2 -0,9 LC010 2,3 -1,0 LC011 2,1 -1,0 LC012 2,4 -1,6 LC013 2,0 -1,1 LC014 2,3 -1,0 2,3 - 1 , 1LC015 2,0 -1,2 LC016 2,1 -1,0 2,1 - 1 , 0LC017 2,4 -1,2 LC018 1,8 -0,9 1,7 - 0 , 9LC019 2,3 -1,0 LC020 2,7 -1,3 LC021 2,1 -1,3 LC022 2,4 -1,4 LC023 1,8 -1,0 LC024 2,4 -1,3 LC025 1,8 -1,0 LC026 2,2 -1,2 LC027 2,5 -1,2 LC028 1,9 -1,3 LC029 2,4 -1,4 LC030 2,2 -1,1 LC031 1,8 -1,2 LC032 1,4 -1,2 LC033 1,4 -1,1 LC034 1,8 -1,2 LC035 1,4 -1,0 LC036 1,7 -1,0 LC037 2,0 -1,1 LC038 1,9 -1,1 1,9 - 1 , 1LC039 1,5 -1,1 LC040 1,7 -0,9 LC041 1,6 -1,2 LC042 1,8 -0,9 LC043 1,7 -0,9 LC044 1,6 -1,0 LC045 1,8 -1,1 LC046 1,5 -1,1 LC047 1,5 -1,2 LC048 1,6 -1,2 LC049 1,5 -1,1 LC050 1,5 -1,0 LC051 2,1 -1,0 LC052 2,2 -1,3 2,2 - 1 , 3LC053 2,1 -1,1 LC054 2,1 -1,1 LC055 2,2 -1,1 LC056 2,2 -0,9 2,2 - 0 , 8LC057 2,2 -1,0 LC058 1,9 -1,0 LC059 1,9 -1,2

Corrected values, carbon and oxygen isotopes:sample/d13CVPDB(‰)/d18OVPDB(‰)/duplicatesLC060 1,5 -1,2 1,5 - 1 , 4LC061 1,8 -1,3 LC062 1,9 -1,0 LC063 2,2 -1,0 LC064 2,4 -1,2 LC065 2,4 -1,0 LC066 2,4 -1,0 2,3 - 0 , 9LC067 2,2 -1,1 LC068 2,2 -1,2 LC069 2,0 -1,5 LC070 2,1 -1,5 LC071 2,2 -1,0 LC072 1,9 -1,1 LC073 1,8 -1,1 LC074 1,7 -1,1 LC075 1,6 -1,1 LC076 1,4 -0,6 LC077 1,5 -1,3 LC078 1,7 -1,1 LC079 2,2 -1,2 LC080 2,0 -1,2 LC081 2,5 -1,6 LC082 2,4 -1,2 LC083 2,4 -1,1 LC084 2,1 -1,3 LC085 1,9 -1,5 LC086 1,9 -1,3 LC087 1,8 -1,2 LC088 1,7 -1,1 LC089 1,5 -1,2 LC090 1,8 -1,2 LC091 1,8 -1,3 LC092 2,0 -0,9 LC093 1,8 -1,2 LC094 1,7 -1,2 LC095 2,0 -1,3 LC096 1,8 -1,1 LC097 2,1 -1,0 LC098 1,9 -0,9 LC099 1,9 -1,1 LC100 1,9 -1,1 LC101 2,0 -0,9 LC102 1,9 -1,0 LC103 1,8 -1,2 LC104 1,9 -0,8 LC105 1,9 -1,0 LC106 2,0 -1,0 2,0 - 1 , 1LC107 2,0 -1,0 LC108 1,7 -1,2 LC109 1,9 -1,2 LC110 2,1 -0,8 LC111 1,8 -1,1 LC112 1,7 -1,3 LC113 1,6 -1,2 LC114 1,6 -1,1 LC115 2,0 -1,0 LC116 1,6 -1,2 LC117 1,7 -1,1 LC118 1,7 -1,1 LC119 1,7 -1,2 LC120 1,6 -1,2 LC121 1,5 -1,4 LC122 1,5 -1,1 LC123 1,6 -1,2 LC124 1,8 -1,2 LC125 1,9 -1,2 LC126 1,8 -1,5 LC127 2,0 -1,6 LC128 1,7 -1,5 1,8 - 1 , 3LC129 2,0 -1,1 LC130 1,8 -1,2 1,9 - 1 , 3LC131 2,0 -1,2 2,0 - 1 , 3LC132 1,8 -0,9 LC133 1,5 -1,2 LC134 1,6 -1,2 LC135 1,6 -1,3 LC136 1,4 -1,4 1,4 - 1 , 4LC137 1,8 -1,5 LC138 1,9 -1,2 LC139 1,6 -1,2 LC140 1,8 -1,3

Corrected values, carbon and oxygen isotopes:sample/d13CVPDB(‰)/d18OVPDB(‰)/duplicatesLC141 1,8 -1,1 LC142 1,5 -1,4 LC143 1,5 -1,2 1,5 - 1 , 2LC144 1,8 -1,8 1,8 - 1 , 8LC145 1,5 -1,3 LC146 1,0 -1,3 0,9 - 1 , 3LC147 1,2 -1,0 LC148 1,5 -1,1 LC149 1,6 -1,1 LC150 1,8 -1,3 LC151 1,7 -1,2 LC152 1,8 -1,3 1,8 - 1 , 2LC153 1,6 -1,2 LC154 1,9 -1,6 LC155 1,0 -1,3 LC156 1,3 -1,4 LC157 1,4 -1,0 LC158 0,9 -1,1 LC159 1,1 -1,5 LC160 1,4 -1,5 1,4 - 1 , 4LC161 1,1 -1,0 LC162 1,3 -0,9 LC163 1,6 -1,5 LC164 1,3 -1,1 LC165 0,7 -1,0 LC166 1,3 -0,6 LC167 0,8 -1,5 LC168 1,1 -1,2 1,1 - 1 , 2LC169 1,3 -1,2 LC170 0,8 -1,2 0,8 - 1 , 2LC171 0,8 -1,1 LC172 0,8 -1,7 LC173 1,1 -1,4 LC174 1,1 -1,0 LC175 1,2 -0,9 LC176 1,1 -1,2 LC177 1,0 -0,9 1,0 - 1 , 0LC178 1,3 -1,3 LC179 1,2 -1,5 LC180 0,8 -1,1 LC181 1,0 -1,3 0,9 - 1 , 3LC182 1,2 -1,1 LC183 1,1 -1,0 LC184 1,5 -1,3 LC185 1,1 -1,1 LC186 0,9 -0,9 LC187 1,2 -1,3 LC188 1,3 -1,2 LC189 1,0 -1,1 LC190 1,1 -1,2 LC191 1,5 -1,5 LC192 1,1 -1,2 LC193 1,2 -1,0 1,2 - 1 , 1LC194 1,4 -1,2 LC195 1,4 -1,2 LC196 1,2 -1,1 LC197 0,8 -1,1 LC198 0,9 -1,2

142

Appendix

143

TIC %, TOC %, and stable organic carbon isotope values in ‰:depth sample TIC TOC delta13Corg m % % duplicate177,5 LC198 -26,5 176,9 LC197 176,0 LC196 175,2 LC195 62,7 1,2 -26,6 174,5 LC194 173,6 LC193 66,2 0,7 -26,5 172,8 LC192 171,3 LC191 -27,0 170,2 LC190 169,2 LC189 60,8 0,5 -26,8 - 2 6 , 8168,1 LC188 167,2 LC187 165,9 LC186 165,0 LC185 65,0 -26,7 163,9 LC184 162,9 LC183 75,0 0,6 -26,6 162,0 LC182 161,2 LC181 160,4 LC180 159,8 LC179 56,5 -26,5 158,7 LC178 157,9 LC177 69,7 0,4 -26,8 157,2 LC176 154,1 LC175 153,0 LC174 80,0 0,4 -27,0 - 2 7 , 0152,0 LC173 151,1 LC172 150,8 LC171 149,8 LC170 65,3 0,6 -27,0 148,7 LC169 148,0 LC168 147,6 LC167 147,0 LC166 73,4 -27,0 146,1 LC165 145,6 LC164 144,9 LC163 144,0 LC162 68,2 0,2 -26,2 143,5 LC161 143,0 LC160 50,0 -26,7 142,0 LC159 141,1 LC158 69,5 0,2 -26,5 140,5 LC157 139,5 LC156 57,1 -26,6 138,3 LC155 137,4 LC154 40,0 0,6 -26,4 135,0 LC153 134,0 LC152 51,5 0,6 -26,5 - 2 6 , 4133,4 LC151 132,6 LC150 132,0 LC149 131,2 LC148 55,0 0,5 -26,4 130,9 LC147 129,9 LC146 129,4 LC145 -27,0 128,7 LC144 128,2 LC143 127,4 LC142 65,0 0,6 -26,7 127,0 LC141 126,2 LC140 125,7 LC139 125,2 LC138 124,8 LC137 124,1 LC136 123,2 LC135 122,4 LC134 60,0 0,5 -26,8 - 2 6 , 7121,9 LC133 121,3 LC132 69,5 -26,3 - 2 6 , 3121,1 LC131 120,9 LC130 120,5 LC129 120,1 LC128 119,9 LC127 119,6 LC126 119,2 LC125 -26,6 119,0 LC124 118,8 LC123 118,5 LC122 61,2 0,4 -26,3 - 2 6 , 4118,2 LC121 117,6 LC120 117,3 LC119

TIC %, TOC %, and stable organic carbon isotope values in ‰:depth sample TIC TOC delta13Corg m % % duplicate117,0 LC118 116,8 LC117 116,2 LC116 115,9 LC115 115,7 LC114 115,2 LC113 -26,4 115,0 LC112 114,8 LC111 114,6 LC110 55,3 0,4 -26,3 - 2 6 , 5114,4 LC109 114,1 LC108 113,7 LC107 -26,5 113,4 LC106 113,0 LC105 112,7 LC104 61,2 -26,5 112,4 LC103 112,1 LC102 111,9 LC101 -26,8 111,7 LC100 111,2 LC099 110,9 LC098 68,2 0,3 -26,3 - 2 6 , 4110,6 LC097 110,0 LC096 109,9 LC095 109,3 LC094 108,9 LC093 108,5 LC092 60,0 0,4 -26,1 108,1 LC091 108,0 LC090 107,4 LC089 -26,4 107,1 LC088 106,6 LC087 106,0 LC086 -26,9 105,1 LC085 104,5 LC084 55,0 0,5 -26,5 103,9 LC083 103,1 LC082 102,8 LC081 -26,4 102,1 LC080 101,6 LC079 101,0 LC078 100,1 LC077 99,5 LC076 65,0 -26,2 99,0 LC075 98,1 LC074 97,9 LC073 97,2 LC072 -26,4 96,9 LC071 96,3 LC070 96,0 LC069 -26,3 95,6 LC068 95,1 LC067 94,7 LC066 60,7 0,5 -26,1 - 2 6 , 194,3 LC065 93,5 LC064 92,9 LC063 -26,3 92,6 LC062 91,8 LC061 91,4 LC060 -26,7 90,9 LC059 90,6 LC058 90,1 LC057 89,8 LC056 68,0 0,3 -26,6 - 2 6 , 689,1 LC055 89,0 LC054 88,7 LC053 -26,6 88,5 LC052 88,1 LC051 87,6 LC050 65,3 0,4 -26,6 - 2 6 , 587,3 LC049 86,6 LC048 86,1 LC047 85,5 LC046 85,0 LC045 84,6 LC044 62,5 0,5 -26,6 84,0 LC043 83,6 LC042 82,8 LC041 82,3 LC040 81,9 LC039

TIC %, TOC %, and stable organic carbon isotope values in ‰:depth sample TIC TOC delta13Corg m % % duplicate81,6 LC038 53,6 0,6 -26,6 80,8 LC037 80,0 LC036 79,0 LC035 78,1 LC034 77,7 LC033 77,5 LC032 77,2 LC031 76,9 LC030 60,0 0,5 -26,8 76,5 LC029 76,0 LC028 75,7 LC027 75,3 LC026 55,0 -26,9 75,0 LC025 74,7 LC024 74,3 LC023 74,0 LC022 73,4 LC021 -27,0 73,1 LC020 72,9 LC019 72,4 LC018 60,7 0,3 -26,9 - 2 7 , 072,0 LC017 71,6 LC016 57,3 -26,9 71,0 LC015 70,7 LC014 70,2 LC013 69,8 LC012 46,3 1,0 -26,8 69,4 LC011 69,0 LC010 45,0 0,9 -26,8 68,6 LC009 68,1 LC008 67,8 LC007 -26,7 66,9 LC006 66,1 LC005 65,2 LC004 52,9 0,6 -26,6 - 2 6 , 664,3 LC003 63,7 LC002 65,3 -27,0 63,0 LC001 62,4 VER087 61,9 VER086 47,7 -27,1 61,1 VER085 60,6 VER084 52,7 0,7 -26,8 - 2 6 , 960,1 VER083 59,5 VER082 54,2 0,5 -26,7 59,1 VER081 58,8 VER080 58,0 VER079 57,0 VER078 53,5 -27,4 56,7 VER077 55,9 VER076 45,8 0,4 -27,1 55,2 VER075 53,9 VER074 52,5 -27,2 52,9 VER073 51,9 VER072 43,8 0,4 -26,6 51,1 VER071 50,5 VER070 41,9 -26,9 49,8 VER069 48,9 VER068 50,0 0,5 -27,1 - 2 7 , 148,1 VER067 47,3 VER066 56,1 -27,8 46,4 VER065 45,6 VER064 60,6 0,5 -27,6 44,9 VER063 43,6 VER062 61,2 -27,8 43,2 VER061 42,6 VER060 65,9 0,3 -27,8 42,1 VER059 41,8 VER058 41,2 VER057 40,6 VER056 55,0 0,6 -27,7 - 2 7 , 939,9 VER055 39,1 VER054 66,2 0,4 -27,6 38,8 VER053 37,6 VER052 36,9 VER051 36,4 VER050 67,0 0,6 -27,6 36,0 VER049 35,5 VER048 35,2 VER047 35,0 VER046

TIC %, TOC %, and stable organic carbon isotope values in ‰:depth sample TIC TOC delta13Corg m % % duplicate34,3 VER045 34,0 VER044 58,5 0,4 -27,4 33,5 VER043 32,9 VER042 62,1 0,5 -27,5 32,3 VER041 31,8 MOR010 31,4 MOR009 59,4 0,4 -27,6 - 2 7 , 831,1 MOR008 30,8 MOR007 30,2 MOR006 74,5 0,2 -27,8 29,7 MOR005 69,4 0,3 -27,9 - 2 7 , 729,4 MOR004 28,9 MOR003 65,2 -27,9 28,1 MOR002 -27,3 27,5 MOR001 69,5 -28,0 26,8 VER040 69,0 -28,1 26,1 VER039 25,2 VER038 67,3 0,5 -27,7 24,3 VER037 23,1 VER036 63,2 0,3 -28,1 22,2 VER035 21,6 VER034 66,5 0,4 -28,0 - 2 8 , 220,8 VER033 20,3 VER032 75,1 0,2 -28,2 20,0 VER031 19,6 VER030 55,8 0,5 -28,1 18,8 VER029=B4 48,5 3,5 -27,5 - 2 7 , 518,2 VER028 63,1 0,5 -28,1 18,0 VER027 70,9 0,5 -28,3 - 2 8 , 517,8 VER026=B3 44,1 4,1 -27,4 17,6 VER025=B2 46,3 2,8 -27,8 - 2 7 , 717,5 VER024=B1 49,0 2,9 -27,6 17,1 VER023 16,4 VER022 67,1 0,3 -28,4 15,9 VER021 14,9 VER020 59,9 0,7 -28,0 14,3 VER019 13,6 VER018 75,1 0,4 -27,9 12,4 VER017 12,1 VER016 69,1 0,3 -28,0 11,6 VER015 11,3 VER014 78,0 0,6 -28,1 - 2 8 , 210,3 VER013 9,7 VER012 66,2 0,5 -28,3 9,3 VER011 8,8 VER010 71,4 0,3 -28,3 - 2 8 , 38,5 VER009 7,8 VER008 58,7 0,6 -28,1 7,0 VER007 6,3 VER006 72,9 0,3 -28,3 5,8 VER005 5,1 VER004 46,9 0,7 -28,1 4,5 VER003 4,1 VER002 52,0 0,6 -28,4 3,3 VER001

144

145

Appendix

Rock-Eval 6 data, externally re-calibratedsample S1 S2 S3 Tmax C org PI HI OI mg/g mg/g mg/g °C % (100*S2)/TOC (100*S3)/TOCVER002 0,1 1,4 0,4 437,0 0,6 0,0 232,8 72,1VER004 0,1 1,5 0,4 437,0 0,7 0,1 221,2 60,6VER006 0,0 0,7 0,3 438,0 0,3 0,1 248,1 125,9VER008 0,1 1,2 0,4 437,0 0,6 0,1 209,1 76,4VER010 0,0 0,8 0,3 436,0 0,3 0,0 242,4 90,9VER012 0,1 1,2 0,3 437,0 0,5 0,0 241,7 66,7VER014 0,1 1,4 0,4 436,0 0,6 0,0 243,9 68,4VER016 0,0 0,6 0,3 434,0 0,3 0,0 210,3 110,3VER018 0,0 0,7 0,4 438,0 0,4 0,1 208,6 111,4VER020 0,1 1,7 0,4 435,0 0,7 0,0 254,4 64,7VER022 0,0 0,7 0,4 437,0 0,3 0,0 219,4 119,4VER024 0,3 6,9 1,1 430,0 2,9 0,0 239,8 38,8VER025 0,2 7,9 0,8 428,0 2,8 0,0 286,2 29,5VER026 0,5 13,9 0,9 428,0 4,1 0,0 344,0 21,2VER027 0,0 1,2 0,4 439,0 0,5 0,0 246,9 77,6VER028 0,0 1,2 0,3 437,0 0,5 0,0 258,7 71,7VER029 0,5 13,4 0,7 427,0 3,5 0,0 382,8 19,5VER030 0,0 1,3 0,4 436,0 0,5 0,0 245,3 69,8VER032 0,0 0,5 0,3 440,0 0,2 0,0 220,8 108,3VER034 0,0 0,9 0,3 435,0 0,4 0,0 236,1 88,9VER036 0,0 0,7 0,3 435,0 0,3 0,1 221,2 100,0VER038 0,0 0,8 0,5 434,0 0,5 0,0 172,3 106,4MOR005 0,0 0,6 0,4 434,0 0,3 0,0 213,3 123,3MOR006 0,0 0,5 0,3 433,0 0,2 0,0 235,0 160,0MOR009 0,0 0,7 0,4 432,0 0,4 0,0 194,3 108,6VER042 0,0 1,0 0,4 438,0 0,5 0,0 212,2 87,8VER044 0,1 0,7 0,4 438,0 0,4 0,1 192,1 115,8VER050 0,0 1,4 0,5 438,0 0,6 0,0 220,3 76,6VER054 0,0 0,9 0,4 436,0 0,4 0,0 204,7 102,3VER056 0,1 1,5 0,4 435,0 0,6 0,0 241,9 62,9VER060 0,0 0,6 0,4 430,0 0,3 0,0 206,7 116,7VER068 0,0 0,8 0,5 432,0 0,5 0,0 186,7 108,9VER076 0,0 0,6 0,5 430,0 0,4 0,0 152,5 112,5VER082 0,0 0,9 0,4 437,0 0,5 0,0 178,4 78,4VER084 0,1 1,4 0,4 437,0 0,7 0,0 190,4 54,8LC004 0,0 1,2 0,4 431,0 0,6 0,0 194,9 66,1LC010 0,1 1,9 0,4 428,0 0,9 0,0 225,6 48,8LC012 0,1 2,1 0,5 428,0 1,0 0,0 223,2 49,5LC030 0,0 1,0 0,4 434,0 0,5 0,0 183,0 67,9LC038 0,0 1,0 0,4 431,0 0,6 0,0 177,6 62,1LC050 0,0 0,6 0,3 426,0 0,4 0,0 150,0 72,5LC076 n.n. n.n. n.n. LC084 0,0 0,8 0,4 427,0 0,5 0,0 153,1 77,6LC092 0,0 0,6 0,3 428,0 0,4 0,0 145,0 72,5LC110 0,0 0,6 0,3 425,0 0,4 0,0 137,2 67,4LC122 0,0 0,6 0,3 427,0 0,4 0,0 139,0 75,6LC134 0,0 0,8 0,4 432,0 0,5 0,0 174,5 83,0LC142 0,0 0,8 0,6 432,0 0,6 0,0 136,1 103,3LC148 0,0 0,7 0,4 428,0 0,5 0,0 146,8 78,7LC154 0,0 0,8 0,6 429,0 0,6 0,0 133,9 101,8LC162 0,0 0,4 0,6 436,0 0,2 0,0 181,0 271,4LC166 n.n. n.n. n.n. LC170 0,0 1,2 0,4 433,0 0,6 0,0 188,5 62,3LC183 0,0 1,7 0,3 432,0 0,6 0,0 296,6 44,8LC189 0,0 1,2 0,3 430,0 0,5 0,0 228,3 56,6

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16

,5

19,5

23,0

33

,5

31,3

23

,0

21,5

23

,5

22,6

40

,1

27,6

PIG

1.6

3

MR

M

area

83

,3

67,6

65

,0

209,

5 17

4,3

183,

6 94

,2

98,6

81

,2

76,8

14

5,7

89,9

PIG

1.5

6

MR

M

area

12

1,1

87,7

85

,1

271,

1 15

7,7

195,

3 82

,5

97,4

63

,4

63,6

88

,8

54,3

PIG

1.4

5

MR

M

area

66

,1

51,3

65

,8

173,

5 18

0,5

145,

8 95

,8

84,5

87

,1

78,0

13

7,1

85,1

PIG

1.3

5 1,

51

437

0,01

83

29

M

RM

ar

ea

15,9

14

,1

18

,5

29,1

36

,0

30,9

29

,4

31,1

28

,7

37,1

25

,3PI

G 1

.28

M

RM

ar

ea

36,6

45

,1

60

,5

59,3

93

,9

36,4

48

,0

31,5

35

,3

40,9

33

,1PI

G 1

.17

0,57

43

5 0,

02

91

153

MR

M

area

26

,5

21,3

16

,5

55,5

36

,9

34,5

14

,2

18,5

13

,4

15,1

16

,4

16,1

PIG

1.1

3

MR

M

area

34

,7

38,3

59,2

60

,5

77,8

21

,4

28,3

12

,7

13,9

13

,8

11,7

PIG

1.5

MR

M

area

83

,6

112,

4 31

,6

134,

0 11

9,5

117,

0 47

,1

46,1

34

,2

42,1

35

,0

29,6

146

147

Appendix

N-a

lkan

es M

id-P

olis

h Tr

ough

, Pol

and

sam

ple

TOC

%

Tmax

PI

H

I O

I M

essu

ng

°C

(1

00*S

2)/T

OC

(10

0*S3

)/TO

C

n-al

kane

s

Ty

p

C25

C

26

C27

C

28

C29

C

30

C31

C

32

C33

C

34

C35

C

36A

lkan

mix

10

MR

M

area

317,

5

314,

8

287,

1A

lkan

mix

4

MR

M

area

119,

5

117,

2

105,

2A

lkan

mix

2

MR

M

area

59,0

57,1

52,2

Alk

anm

ix 1

M

RM

ar

ea

29

,3

28

,2

25

,9PI

G 1

.182

3,

8 43

8 0,

01

67

88

MR

M

area

16

,8

12,1

18

,0

14,2

15

,6

8,0

11,3

4,

8 10

,2

3,8

9,0

3,6

PIG

1.1

74

M

RM

ar

ea

33,7

18

,5

33,5

26

,2

25,4

8,

6 12

,4

8,9

27

,0

13,3

6,

2PI

G 1

.166

MR

M

area

42

,6

23,7

47

,8

30,1

40

,9

13,9

25

,3

9,8

19,8

6,

6 17

,4

6,5

PIG

1.1

59

M

RM

ar

ea

56,4

33

,9

67,8

35

,8

60,4

21

,1

24,4

12

,1

25,9

7,

2 20

,3

9,1

PIG

1.1

50

M

RM

ar

ea

30,6

22

,2

40,9

37

,0

40,6

12

,7

25,1

8,

4 16

,6

PIG

1.1

40

2,2

444

0,01

73

14

2 M

RM

ar

ea

15,2

10

,5

18,0

12

,0

13,4

5,

5 8,

0 3,

5 4,

9 2,

3 5,

7 2,

3PI

G 1

.131

MR

M

area

42

,3

23,7

46

,5

30,1

39

,2

14,5

23

,7

9,7

22,8

4,

4 19

,2

9,3

PIG

1.1

22

M

RM

ar

ea

49,7

27

,6

65,1

41

,5

53,6

16

,7

30,2

9,

4 18

,9

6,7

17,1

6,

2PI

G 1

.114

MR

M

area

13

8,7

83,0

22

9,1

107,

1 20

3,9

73,7

12

0,2

38,6

83

,7

30,5

60

,3

6,8

PIG

1.1

06

M

RM

ar

ea

150,

1 80

,3

227,

2 11

2,4

181,

1 60

,2

99,3

31

,7

68,4

19

,8

52,7

15

,4PI

G 1

.98

M

RM

ar

ea

96,3

52

,3

141,

2 73

,8

107,

9 33

,7

57,4

19

,3

40,0

10

,1

33,0

9,

2PI

G 1

.91

M

RM

ar

ea

138,

5

10

7,9

165,

4 58

,0

88,8

33

,2

65,0

29

,0

50,3

13

,8PI

G 1

.86

11,7

42

2 0,

01

94

39

MR

M

area

48

,5

25,1

69

,3

31,5

51

,4

17,9

27

,7

9,3

19,9

6,

0 16

,6

4,5

PIG

1.7

8

MR

M

area

12

9,6

73,0

20

5,8

98,9

17

1,2

62,6

95

,4

31,0

67

,6

18,6

52

,0

14,0

PIG

1.7

1 4,

76

428

0,01

10

3 40

M

RM

ar

ea

59,9

33

,8

98,2

44

,8

70,6

25

,4

37,6

12

,8

27,1

8,

0 21

,0

6,4

PIG

1.6

3

MR

M

area

20

4,3

116,

7 35

4,0

120,

5 25

7,0

90,5

14

3,1

45,4

10

2,8

25,6

77

,2

3,8

PIG

1.5

6

MR

M

area

10

7,7

61,3

18

8,6

92,3

14

1,6

52,1

83

,1

26,6

60

,3

15,5

49

,7

13,8

PIG

1.4

5

MR

M

area

18

3,8

109,

4 34

0,7

132,

4 26

9,0

96,7

14

6,7

50,2

10

5,5

32,1

84

,3

25,2

PIG

1.3

5 1,

51

437

0,01

83

29

M

RM

ar

ea

35,4

23

,2

45,0

25

,0

38,3

15

,4

22,9

9,

1 19

,8

6,2

17,7

6,

2PI

G 1

.28

M

RM

ar

ea

40,2

32

,4

51,9

33

,2

52,5

20

,9

34,2

12

,8

27,5

7,

6 21

,9

5,4

PIG

1.1

7 0,

57

435

0,02

91

15

3 M

RM

ar

ea

18,4

17

,3

17,7

2,

7 17

,4

12,8

PIG

1.1

3

MR

M

area

12

,2

10,1

11

,0

16,3

11

,3

6,6

11,9

PI

G 1

.5

M

RM

ar

ea

39,0

31

,1

33,2

N-a

lkan

es M

id-P

olis

h Tr

ough

, Pol

and:

sam

ple

TOC

%

Tmax

PI

H

I O

I M

essu

ng

°C

(100

*S2)

/TO

C

(100

*S3)

/TO

C

n-

alka

nes

Ty

p

C37

C

38

C39

C

40

C41

Alk

anm

ix 1

0

M

RM

ar

ea

19

7,8

Alk

anm

ix 4

M

RM

ar

ea

70

,3

Alk

anm

ix 2

M

RM

ar

ea

38

,6

Alk

anm

ix 1

M

RM

ar

ea

17

,0

PIG

1.1

82

3,8

438

0,01

67

88

M

RM

ar

ea

5,9

2,5

6,8

1,6

3,2

PIG

1.1

74

M

RM

ar

ea

7,6

3,9

6,7

PI

G 1

.166

MR

M

area

8,

6 3,

0 7,

5 1,

2 2,

6PI

G 1

.159

MR

M

area

9,

6 3,

0 7,

0

PIG

1.1

50

M

RM

ar

ea

PIG

1.1

40

2,2

444

0,01

73

14

2 M

RM

ar

ea

4,5

1,9

5,1

1,8

1,9

PIG

1.1

31

M

RM

ar

ea

9,7

3,9

12,5

PIG

1.1

22

M

RM

ar

ea

10,6

3,

1 10

,1

2,1

4,1

PIG

1.1

14

M

RM

ar

ea

20,3

5,

8 12

,2

5,2

5,3

PIG

1.1

06

M

RM

ar

ea

20,9

6,

8 15

,3

4,6

6,6

PIG

1.9

8

MR

M

area

13

,9

4,3

10,4

2,

6 5,

2PI

G 1

.91

M

RM

ar

ea

15,6

5,

2 9,

0

PIG

1.8

6 11

,7

422

0,01

94

39

M

RM

ar

ea

7,3

2,5

6,0

2,4

2,8

PIG

1.7

8

MR

M

area

16

,2

4,5

10,5

PIG

1.7

1 4,

76

428

0,01

10

3 40

M

RM

ar

ea

8,3

2,1

5,5

PI

G 1

.63

M

RM

ar

ea

21,3

6,

5 14

,8

PI

G 1

.56

M

RM

ar

ea

16,3

4,

8 13

,8

5,0

5,6

PIG

1.4

5

MR

M

area

22

,9

6,5

17,0

3,

6 PI

G 1

.35

1,51

43

7 0,

01

83

29

MR

M

area

8,

2 2,

7 6,

1 1,

7 2,

6PI

G 1

.28

M

RM

ar

ea

7,3

2,4

4,5

1,0

1,9

PIG

1.1

7 0,

57

435

0,02

91

15

3 M

RM

ar

ea

PIG

1.1

3

MR

M

area

PI

G 1

.5

M

RM

ar

ea

148

149

Appendix

N-a

lkan

es, i

sopr

enoi

ds, V

ocon

tian

Bas

in, F

ranc

e:m

ass

mik

rol.

sam

ple

TOC

%

Tmax

PI

H

I O

I m

easu

rem

ent

Isop

reno

ids

n-

alka

nes

g µ

l

°C

(

100*

S2)/T

OC

(1

00*S

3)/T

OC

Typ

Pr

ista

n Ph

ytan

C

15

C16

C

17

C18

C

19

C20

C

21

C22

A

lkan

mix

10

MR

M

area

23

5,3

267,

0 27

3,5

271,

6 4,

8 27

9,0

3,

0

MG

S-1-

1 1.

5

M

RM

ar

ea

721,

7 64

7,0

262,

0 38

0,4

469,

8 39

2,6

395,

8 35

2,8

320,

2 27

3,0

19,8

10

00,0

V

ER00

4 0,

7 43

7,0

0,1

221,

2 60

,6

MR

M

area

99

3,3

348,

5 55

5,7

957,

6 12

93,6

11

75,9

12

38,9

10

79,0

10

48,5

90

6,4

22,6

50

0,0

VER

012

0,5

437,

0 0,

0 24

1,7

66,7

M

RM

ar

ea

944,

4 34

4,1

602,

2 10

25,7

13

53,3

11

91,0

12

33,1

10

53,2

10

13,5

86

4,5

24,1

15

00,0

V

ER02

0 0,

7 43

5,0

0,0

254,

4 64

,7

MR

M

area

98

6,9

313,

5 84

7,0

1018

,3

1140

,9

988,

7 10

06,5

87

0,7

835,

3 71

0,7

25,5

10

00,0

V

ER02

2 0,

3 43

7,0

0,0

219,

4 11

9,4

MR

M

area

72

9,1

307,

3 50

9,0

918,

0 12

08,6

10

65,3

10

51,3

90

8,1

847,

0 73

0,7

16,3

15

00,0

V

ER02

4 2,

9 43

0,0

0,0

239,

8 38

,8

MR

M

area

18

16,0

44

5,8

422,

2 47

5,4

518,

6 39

1,5

409,

1 31

5,9

284,

4 24

9,4

19,5

15

00,0

V

ER02

5 2,

8 42

8,0

0,0

286,

2 29

,5

MR

M

area

16

81,3

36

0,8

391,

2 39

5,4

449,

4 33

5,4

353,

7 26

5,4

239,

2 21

0,7

14,6

15

00,0

V

ER02

6 4,

1 42

8,0

0,0

344,

0 21

,2

MR

M

area

14

62,9

32

9,1

254,

9 30

3,2

375,

4 28

3,0

298,

3 22

5,2

205,

2 17

6,0

25,3

10

00,0

V

ER02

7 0,

5 43

9,0

0,0

246,

9 77

,6

MR

M

area

93

2,0

334,

0 50

9,4

902,

6 12

29,1

10

91,0

11

17,7

92

4,3

894,

0 74

6,3

25,1

10

00,0

V

ER02

8 0,

5 43

7,0

0,0

258,

7 71

,7

MR

M

area

10

04,2

37

7,5

659,

2 11

13,7

14

51,8

12

68,2

13

04,6

10

93,5

10

37,7

86

3,9

17,8

15

00,0

V

ER02

9 3,

5 42

7,0

0,0

382,

8 19

,5

MR

M

area

14

27,0

34

4,7

385,

4 41

5,6

458,

4 34

7,6

339,

5 26

7,2

245,

3 20

6,5

20,1

10

00,0

V

ER03

2 0,

2 44

0,0

0,0

220,

8 10

8,3

MR

M

area

28

3,2

126,

3 85

,4

249,

0 44

2,4

463,

8 51

6,6

459,

5 45

0,5

384,

525

,9

1000

,0

VER

038

0,5

434,

0 0,

0 17

2,3

106,

4 M

RM

ar

ea

754,

7 24

6,3

603,

9 84

5,6

1011

,5

893,

9 94

3,0

813,

8 79

9,2

669,

425

,1

1000

,0

MO

R00

5 0,

3 43

4,0

0,0

213,

3 12

3,3

MR

M

area

70

4,4

275,

7 34

7,5

691,

2 95

0,9

859,

1 87

7,2

746,

9 70

8,3

597,

324

,8

1000

,0

VER

044

0,4

438,

0 0,

1 19

2,1

115,

8 M

RM

ar

ea

622,

3 21

3,7

408,

9 69

1,1

894,

9 75

5,6

778,

0 63

9,8

615,

2 51

8,6

21,3

10

00,0

V

ER05

0 0,

6 43

8,0

0,0

220,

3 76

,6

MR

M

area

12

60,3

41

7,4

422,

7 82

3,5

1090

,2

958,

6 96

4,8

820,

6 78

6,8

675,

821

,5

1000

,0

VER

054

0,4

436,

0 0,

0 20

4,7

102,

3 M

RM

ar

ea

686,

4 23

6,1

167,

4 44

5,1

702,

8 68

4,0

723,

7 60

6,9

583,

1 47

9,1

25,0

10

00,0

V

ER06

0 0,

3 43

0,0

0,0

206,

7 11

6,7

MR

M

area

57

0,3

192,

1 24

4,7

493,

9 69

9,1

633,

6 65

1,6

554,

6 53

4,2

442,

422

,9

1000

,0

VER

068

0,5

432,

0 0,

0 18

6,7

108,

9 M

RM

ar

ea

769,

0 24

8,6

286,

3 59

4,0

825,

7 76

8,5

770,

8 65

3,6

616,

9 50

7,2

19,3

10

00,0

V

ER07

6 0,

4 43

0,0

0,0

152,

5 11

2,5

MR

M

area

50

4,9

171,

7 12

3,8

336,

1 57

7,4

557,

8 58

2,1

489,

4 44

8,5

368,

322

,7

1000

,0

VER

084

0,7

437,

0 0,

0 19

0,4

54,8

M

RM

ar

ea

1426

,0

447,

5 52

6,4

800,

1 10

02,1

88

1,9

956,

0 78

2,3

765,

7 64

5,0

20,6

10

00,0

LC

004

0,6

431,

0 0,

0 19

4,9

66,1

M

RM

ar

ea

473,

7 22

7,6

89,7

18

4,6

283,

6 26

0,9

285,

6 21

5,4

208,

7 18

2,4

25,2

50

0,0

LC03

0 0,

5 43

4,0

0,0

183,

0 67

,9

MR

M

area

77

6,2

416,

1 65

,2

272,

8 47

1,0

447,

7 48

4,3

374,

0 37

9,4

316,

825

,1

500,

0 LC

038

0,6

431,

0 0,

0 17

7,6

62,1

M

RM

ar

ea

1095

,3

517,

8 20

2,5

423,

4 67

1,4

602,

2 67

9,0

516,

1 51

7,1

418,

721

,7

500,

0 LC

050

0,4

426,

0 0,

0 15

0,0

72,5

M

RM

ar

ea

471,

4 22

2,0

43,8

14

0,6

267,

5 25

7,3

290,

8 23

5,1

244,

4 20

8,8

24,8

50

0,0

LC08

4 0,

5 42

7,0

0,0

153,

1 77

,6

MR

M

area

82

0,3

378,

3 15

2,1

255,

2 39

6,3

341,

0 39

0,5

290,

4 30

5,6

262,

125

,2

500,

0 LC

092

0,4

428,

0 0,

0 14

5,0

72,5

M

RM

ar

ea

596,

9 30

4,9

39,2

17

2,2

310,

5 30

8,5

335,

3 26

7,8

277,

4 24

7,4

25,4

50

0,0

LC11

0 0,

4 42

5,0

0,0

137,

2 67

,4

MR

M

area

46

0,5

227,

5 23

,0

102,

5 21

5,2

226,

7 27

8,4

222,

7 23

8,0

208,

622

,5

500,

0 LC

122

0,4

427,

0 0,

0 13

9,0

75,6

M

RM

ar

ea

277,

0 15

5,4

10,1

47

,7

131,

6 15

2,6

197,

5 16

6,6

182,

9 15

7,0

21,2

50

0,0

LC13

4 0,

5 43

2,0

0,0

174,

5 83

,0

MR

M

area

38

8,8

207,

1 14

,4

68,1

18

1,4

203,

0 24

6,9

203,

7 21

4,2

191,

725

,6

500,

0 LC

142

0,6

432,

0 0,

0 13

6,1

103,

3 M

RM

ar

ea

761,

2 39

5,3

46,1

15

5,4

338,

4 32

7,9

394,

5 32

2,6

340,

1 29

2,0

19,5

50

0,0

LC14

8 0,

5 42

8,0

0,0

146,

8 78

,7

MR

M

area

35

0,6

183,

2 16

,0

77,0

17

2,7

176,

4 20

8,4

183,

1 19

8,6

176,

827

,2

500,

0 LC

154

0,6

429,

0 0,

0 13

3,9

101,

8 M

RM

ar

ea

762,

7 35

0,2

128,

5 31

1,9

510,

4 47

4,7

505,

7 40

3,2

396,

6 37

3,9

16,4

50

0,0

LC16

2 0,

2 43

6,0

0,0

181,

0 27

1,4

MR

M

area

38

,1

31,8

3,

9 7,

6 33

,7

50,1

65

,3

59,3

62

,6

60,3

25,1

50

0,0

LC17

0 0,

6 43

3,0

0,0

188,

5 62

,3

MR

M

area

90

0,8

563,

5 71

,6

234,

0 48

0,0

450,

3 50

4,3

416,

3 40

2,4

384,

924

,9

500,

0 LC

183

0,6

432,

0 0,

0 29

6,6

44,8

M

RM

ar

ea

1051

,9

639,

1 15

2,7

338,

7 57

5,4

526,

9 61

0,0

494,

2 47

1,3

450,

625

,0

1000

,0

LC18

9 0,

5 43

0,0

0,0

228,

3 56

,6

MR

M

area

43

8,6

303,

4 67

,8

175,

0 31

1,7

297,

3 32

2,6

261,

1 25

2,2

249,

4

N-a

lkan

es V

ocon

tian

Bas

in, F

ranc

e:m

ass

mik

rol.

sam

ple

TOC

%

Tmax

PI

H

I O

I m

easu

rem

ent

g

µl

°C

(1

00*S

2)/T

OC

(1

00*S

3)/T

OC

n-al

kane

s

Ty

p

C23

C

24

C25

C

26

C27

C

28

C29

C

30

Alk

anm

ix 1

0

M

RM

ar

ea

27

1,2

26

3,3

3,

0

MG

S-1-

1 1.

5

M

RM

ar

ea

249,

9 21

5,2

190,

9 18

8,1

143,

8 14

1,7

171,

6 99

,319

,8

1000

,0

VER

004

0,7

437,

0 0,

1 22

1,2

60,6

M

RM

ar

ea

863,

8 74

4,1

726,

7 58

9,0

565,

9 46

1,8

431,

9 28

2,8

22,6

50

0,0

VER

012

0,5

437,

0 0,

0 24

1,7

66,7

M

RM

ar

ea

795,

7 70

2,4

652,

4 56

2,2

506,

9 43

4,4

384,

8 26

3,4

24,1

15

00,0

V

ER02

0 0,

7 43

5,0

0,0

254,

4 64

,7

MR

M

area

66

6,2

579,

0 54

9,4

470,

5 42

7,2

352,

6 34

5,4

223,

225

,5

1000

,0

VER

022

0,3

437,

0 0,

0 21

9,4

119,

4 M

RM

ar

ea

672,

5 59

1,1

563,

2 48

2,2

436,

5 36

0,8

336,

7 22

9,9

16,3

15

00,0

V

ER02

4 2,

9 43

0,0

0,0

239,

8 38

,8

MR

M

area

22

1,4

190,

6 17

2,2

145,

4 13

8,0

124,

6 13

9,5

102,

719

,5

1500

,0

VER

025

2,8

428,

0 0,

0 28

6,2

29,5

M

RM

ar

ea

181,

8 15

6,5

141,

5 11

8,6

109,

6 99

,4

117,

7 72

,714

,6

1500

,0

VER

026

4,1

428,

0 0,

0 34

4,0

21,2

M

RM

ar

ea

154,

7 12

7,6

118,

6 96

,9

91,8

82

,4

101,

0 66

,125

,3

1000

,0

VER

027

0,5

439,

0 0,

0 24

6,9

77,6

M

RM

ar

ea

706,

4 60

3,1

600,

6 49

0,7

452,

1 36

4,5

349,

7 21

7,3

25,1

10

00,0

V

ER02

8 0,

5 43

7,0

0,0

258,

7 71

,7

MR

M

area

83

1,0

720,

4 71

7,0

589,

8 54

3,6

447,

4 42

0,9

279,

017

,8

1500

,0

VER

029

3,5

427,

0 0,

0 38

2,8

19,5

M

RM

ar

ea

180,

6 15

1,6

133,

6 11

0,5

104,

4 91

,2

100,

3 63

,820

,1

1000

,0

VER

032

0,2

440,

0 0,

0 22

0,8

108,

3 M

RM

ar

ea

357,

5 30

7,6

287,

9 24

1,7

223,

6 18

4,0

173,

7 11

5,7

25,9

10

00,0

V

ER03

8 0,

5 43

4,0

0,0

172,

3 10

6,4

MR

M

area

64

8,2

549,

1 54

1,5

431,

6 42

5,0

320,

7 32

4,0

210,

125

,1

1000

,0

MO

R00

5 0,

3 43

4,0

0,0

213,

3 12

3,3

MR

M

area

57

1,4

489,

5 47

7,3

393,

8 34

3,4

289,

8 27

2,0

169,

824

,8

1000

,0

VER

044

0,4

438,

0 0,

1 19

2,1

115,

8 M

RM

ar

ea

508,

2 42

0,6

417,

4 31

9,6

295,

7 23

0,4

221,

2 14

3,9

21,3

10

00,0

V

ER05

0 0,

6 43

8,0

0,0

220,

3 76

,6

MR

M

area

64

1,4

568,

4 52

6,7

448,

6 41

1,6

347,

5 32

9,6

212,

121

,5

1000

,0

VER

054

0,4

436,

0 0,

0 20

4,7

102,

3 M

RM

ar

ea

457,

6 37

8,4

363,

4 28

1,4

256,

5 20

0,9

202,

8 12

4,3

25,0

10

00,0

V

ER06

0 0,

3 43

0,0

0,0

206,

7 11

6,7

MR

M

area

43

1,9

353,

3 34

1,6

265,

1 24

1,2

188,

4 18

6,4

111,

222

,9

1000

,0

VER

068

0,5

432,

0 0,

0 18

6,7

108,

9 M

RM

ar

ea

482,

1 39

5,6

392,

6 29

9,0

293,

3 22

2,5

225,

4 14

1,9

19,3

10

00,0

V

ER07

6 0,

4 43

0,0

0,0

152,

5 11

2,5

MR

M

area

35

3,9

293,

4 29

8,4

227,

4 22

4,9

174,

7 17

7,5

105,

122

,7

1000

,0

VER

084

0,7

437,

0 0,

0 19

0,4

54,8

M

RM

ar

ea

633,

6 51

7,2

511,

2 39

4,5

372,

0 29

7,1

302,

1 17

7,9

20,6

10

00,0

LC

004

0,6

431,

0 0,

0 19

4,9

66,1

M

RM

ar

ea

183,

9 14

7,3

166,

2 11

9,0

129,

1 94

,5

150,

0 25

,2

500,

0 LC

030

0,5

434,

0 0,

0 18

3,0

67,9

M

RM

ar

ea

327,

6 22

7,2

264,

5 19

0,5

205,

8 13

4,1

253,

6 25

,1

500,

0 LC

038

0,6

431,

0 0,

0 17

7,6

62,1

M

RM

ar

ea

430,

6 30

6,3

350,

0 24

6,9

245,

7 20

2,1

310,

0 21

,7

500,

0 LC

050

0,4

426,

0 0,

0 15

0,0

72,5

M

RM

ar

ea

210,

9 14

7,7

171,

6 11

7,7

129,

7 10

5,8

149,

9 24

,8

500,

0 LC

084

0,5

427,

0 0,

0 15

3,1

77,6

M

RM

ar

ea

273,

5 18

8,1

219,

2 14

1,1

168,

6 14

0,2

210,

1 25

,2

500,

0 LC

092

0,4

428,

0 0,

0 14

5,0

72,5

M

RM

ar

ea

247,

6 17

1,7

203,

2 13

8,3

175,

5 12

8,7

184,

2 25

,4

500,

0 LC

110

0,4

425,

0 0,

0 13

7,2

67,4

M

RM

ar

ea

211,

8 15

0,9

172,

6 11

4,6

140,

4 11

0,4

153,

3 22

,5

500,

0 LC

122

0,4

427,

0 0,

0 13

9,0

75,6

M

RM

ar

ea

164,

5 12

2,0

142,

6 97

,0

144,

4 10

8,0

148,

9 21

,2

500,

0 LC

134

0,5

432,

0 0,

0 17

4,5

83,0

M

RM

ar

ea

185,

3 13

2,4

149,

7 93

,1

130,

2 96

,8

161,

4 25

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500,

0 LC

142

0,6

432,

0 0,

0 13

6,1

103,

3 M

RM

ar

ea

288,

4 19

5,9

213,

1 13

8,3

171,

7 13

6,0

319,

0 19

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500,

0 LC

148

0,5

428,

0 0,

0 14

6,8

78,7

M

RM

ar

ea

177,

2 12

5,2

147,

4 91

,6

126,

1 85

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140,

5 27

,2

500,

0 LC

154

0,6

429,

0 0,

0 13

3,9

101,

8 M

RM

ar

ea

371,

3 26

0,9

305,

9 19

8,0

251,

3 20

3,1

344,

5 16

,4

500,

0 LC

162

0,2

436,

0 0,

0 18

1,0

271,

4 M

RM

ar

ea

56,2

41

,9

48,6

34

,2

39,0

32

,9

57,1

25

,1

500,

0 LC

170

0,6

433,

0 0,

0 18

8,5

62,3

M

RM

ar

ea

364,

1 28

5,7

302,

8 21

7,8

253,

9 22

0,3

366,

6 24

,9

500,

0 LC

183

0,6

432,

0 0,

0 29

6,6

44,8

M

RM

ar

ea

414,

7 33

0,9

326,

3 26

0,0

271,

3 24

8,0

389,

6 25

,0

1000

,0

LC18

9 0,

5 43

0,0

0,0

228,

3 56

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MR

M

area

21

1,3

183,

0 18

4,4

142,

7 14

4,3

135,

0 19

4,3

150

151

Appendix

N-a

lkan

es, V

ocon

tian

Bas

in, F

ranc

e:m

ass

mik

rol.

sam

ple

TOC

%

Tmax

PI

H

I O

I m

easu

rem

ent

n-al

kane

g

µl

°C

(100

*S2)

/TO

C (

100*

S3)/T

OC

Typ

C

31

C32

C

33

C34

C

35

C36

C

37

C38

C

39

C40

A

lkan

mix

10

MR

M

area

264,

4

250,

3

240,

43,

0

MG

S-1-

1 1.

5

M

RM

ar

ea

249,

3 57

,6

83,0

57

,4

72,2

26

,9

31,3

18

,2

12,6

14

,019

,8

1000

,0

VER

004

0,7

437,

0 0,

1 22

1,2

60,6

M

RM

ar

ea

596,

9 13

8,4

181,

4 92

,2

89,9

69

,1

66,4

47

,7

49,0

40

,322

,6

500,

0 V

ER01

2 0,

5 43

7,0

0,0

241,

7 66

,7

MR

M

area

57

7,9

169,

4 13

9,9

98,7

84

,5

62,2

61

,8

41,8

46

,2

34,3

24,1

15

00,0

V

ER02

0 0,

7 43

5,0

0,0

254,

4 64

,7

MR

M

area

50

7,6

114,

1 15

6,9

86,7

80

,3

53,7

57

,8

37,7

48

,1

32,4

25,5

10

00,0

V

ER02

2 0,

3 43

7,0

0,0

219,

4 11

9,4

MR

M

area

48

8,2

189,

3 15

7,3

92,2

77

,1

60,7

58

,6

44,0

39

,5

32,7

16,3

15

00,0

V

ER02

4 2,

9 43

0,0

0,0

239,

8 38

,8

MR

M

area

19

3,3

47,1

70

,0

40,5

39

,2

23,4

27

,1

18,4

19

,7

13,6

19,5

15

00,0

V

ER02

5 2,

8 42

8,0

0,0

286,

2 29

,5

MR

M

area

18

3,1

53,4

55

,9

31,0

29

,9

17,1

21

,8

14,9

14

,3

10,5

14,6

15

00,0

V

ER02

6 4,

1 42

8,0

0,0

344,

0 21

,2

MR

M

area

16

1,3

47,4

55

,0

29,2

28

,3

17,2

17

,3

10,4

10

,8

13,7

25,3

10

00,0

V

ER02

7 0,

5 43

9,0

0,0

246,

9 77

,6

MR

M

area

55

2,8

123,

6 16

6,1

86,3

79

,5

65,6

55

,9

45,3

47

,5

40,6

25,1

10

00,0

V

ER02

8 0,

5 43

7,0

0,0

258,

7 71

,7

MR

M

area

62

3,8

177,

4 18

2,2

110,

2 98

,0

67,7

65

,4

51,5

51

,2

46,1

17,8

15

00,0

V

ER02

9 3,

5 42

7,0

0,0

382,

8 19

,5

MR

M

area

17

0,2

50,6

42

,2

23,5

25

,8

16,9

16

,6

14,7

14

,4

11,5

20,1

10

00,0

V

ER03

2 0,

2 44

0,0

0,0

220,

8 10

8,3

MR

M

area

24

1,4

68,3

80

,5

44,1

36

,1

29,2

26

,2

21,0

20

,7

18,7

25,9

10

00,0

V

ER03

8 0,

5 43

4,0

0,0

172,

3 10

6,4

MR

M

area

49

5,0

127,

4 15

2,7

75,3

79

,5

59,4

53

,5

41,8

42

,1

37,8

25,1

10

00,0

M

OR

005

0,3

434,

0 0,

0 21

3,3

123,

3 M

RM

ar

ea

403,

0 10

2,2

111,

6 67

,9

61,8

52

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37,5

31

,7

32,6

28

,424

,8

1000

,0

VER

044

0,4

438,

0 0,

1 19

2,1

115,

8 M

RM

ar

ea

341,

4 10

0,6

96,8

59

,9

57,8

43

,9

46,6

33

,3

37,4

23

,721

,3

1000

,0

VER

050

0,6

438,

0 0,

0 22

0,3

76,6

M

RM

ar

ea

560,

1 12

4,4

135,

8 91

,4

98,5

67

,9

74,5

55

,3

47,2

33

,321

,5

1000

,0

VER

054

0,4

436,

0 0,

0 20

4,7

102,

3 M

RM

ar

ea

355,

1 87

,1

100,

8 54

,4

56,2

42

,6

42,0

30

,8

32,3

23

,125

,0

1000

,0

VER

060

0,3

430,

0 0,

0 20

6,7

116,

7 M

RM

ar

ea

311,

2 98

,8

84,6

49

,0

51,6

39

,1

39,5

30

,7

36,5

22

,822

,9

1000

,0

VER

068

0,5

432,

0 0,

0 18

6,7

108,

9 M

RM

ar

ea

392,

1 10

4,3

110,

9 66

,6

63,3

44

,8

54,8

41

,7

40,9

30

,919

,3

1000

,0

VER

076

0,4

430,

0 0,

0 15

2,5

112,

5 M

RM

ar

ea

287,

7 76

,4

79,3

43

,6

52,8

38

,2

43,4

30

,6

30,6

24

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153

Appendix

Hop

anes

, ste

rane

s, Vo

cont

ian

Bas

in, F

ranc

e.m

ass

mik

rol.

sam

ple

mea

sure

men

t

hopa

nes

Ster

ane

g

µl

peak

: 15

,0

16,0

17

,0

1,0

2,0

3,0

4,0

5,0

6,0

7,0

8,0

9,0

10,0

11

,0

Typ

3b

MC

31R

C

35ab

S C

35ab

R

C26

aaa

C27

13bS

? C

2717

aR?

C27

abS?

C

27ab

R?

C27

aaaS

C

27ab

bS?

C27

abbR

? C

27aa

aR

C28

ba?

C28

ba?

Alk

anm

ix 1

0MR

M

area

3,0

M

GS-

1-1

1.5M

RM

ar

ea

4846

,0

7645

6,0

7893

6,0

19,8

10

00,0

V

ER00

4 M

RM

ar

ea

3426

0,0

1015

92,0

65

601,

0 10

857,

0 58

3759

,0

3842

56,0

10

9012

,0

1498

13,0

30

3071

,0

1176

95,0

10

6343

,0

4416

69,0

33

8946

,0

3839

62,0

22,6

50

0,0

VER

012

MR

M

area

42

565,

0 99

149,

0 73

202,

0 13

565,

0 54

5402

,0

3519

66,0

10

3884

,0

1355

66,0

36

6891

,0

1243

02,0

11

6485

,0

5085

11,0

34

7902

,0

3750

58,0

24,1

15

00,0

V

ER02

0 M

RM

ar

ea

3856

0,0

9222

6,0

6565

2,0

1198

8,0

4666

26,0

30

1759

,0

8769

7,0

1190

70,0

32

9211

,0

1256

60,0

11

0654

,0

4775

64,0

31

7230

,0

3400

32,0

25,5

10

00,0

V

ER02

2 M

RM

ar

ea

3879

5,0

1267

01,0

84

209,

0 13

734,

0 53

4234

,0

3537

95,0

10

6422

,0

1491

93,0

44

0999

,0

1522

25,0

12

2436

,0

6170

69,0

42

9704

,0

4636

10,0

16,3

15

00,0

V

ER02

4 M

RM

ar

ea

8452

,0

4403

6,0

2807

9,0

7525

,0

3270

68,0

21

5962

,0

6630

7,0

8492

8,0

2091

46,0

58

527,

0 56

199,

0 23

6112

,0

1869

33,0

19

4746

,019

,5

1500

,0

VER

025

MR

M

area

69

49,0

38

713,

0 25

476,

0 63

83,0

32

8208

,0

2101

84,0

65

832,

0 88

720,

0 19

5574

,0

5893

9,0

5425

1,0

2309

95,0

18

8795

,0

1977

32,0

14,6

15

00,0

V

ER02

6 M

RM

ar

ea

5885

,0

3754

0,0

2238

5,0

5906

,0

2935

51,0

18

7279

,0

5693

8,0

7638

3,0

1588

62,0

49

176,

0 42

965,

0 17

5763

,0

1685

22,0

18

3939

,025

,3

1000

,0

VER

027

MR

M

area

38

732,

0 12

2645

,0

7531

8,0

1471

2,0

5444

10,0

36

1307

,0

1041

30,0

14

3497

,0

3505

72,0

13

3206

,0

1261

77,0

48

7107

,0

3634

27,0

38

9824

,025

,1

1000

,0

VER

028

MR

M

area

42

910,

0 12

7298

,0

8005

2,0

1544

9,0

6254

30,0

40

2920

,0

1184

54,0

16

4905

,0

4071

33,0

15

5806

,0

1399

38,0

57

4494

,0

4176

05,0

44

0291

,017

,8

1500

,0

VER

029

MR

M

area

74

44,0

40

967,

0 25

453,

0 56

62,0

26

6812

,0

1683

31,0

53

327,

0 68

113,

0 15

8541

,0

5309

3,0

4411

7,0

1853

05,0

16

6121

,0

1755

14,0

20,1

10

00,0

V

ER03

2 M

RM

ar

ea

1896

9,0

5108

2,0

3624

1,0

5676

,0

2247

80,0

14

8187

,0

4528

9,0

6019

4,0

1473

58,0

47

354,

0 41

923,

0 20

2164

,0

1511

14,0

16

8828

,025

,9

1000

,0

VER

038

MR

M

area

31

382,

0 89

415,

0 52

553,

0 92

79,0

44

7118

,0

2908

62,0

85

235,

0 11

5808

,0

2396

40,0

84

122,

0 78

748,

0 33

3539

,0

3046

22,0

31

5285

,025

,1

1000

,0

MO

R00

5 M

RM

ar

ea

2060

4,0

7632

0,0

6289

5,0

1284

8,0

3581

90,0

23

5844

,0

6037

9,0

9747

7,0

3305

95,0

98

315,

0 79

134,

0 50

5833

,0

1917

17,0

19

9854

,024

,8

1000

,0

VER

044

MR

M

area

17

753,

0 68

055,

0 41

735,

0 77

85,0

24

5660

,0

1595

54,0

47

821,

0 63

586,

0 18

1611

,0

6175

6,0

5407

8,0

2889

00,0

96

290,

0 10

0844

,021

,3

1000

,0

VER

050

MR

M

area

28

421,

0 10

6958

,0

6926

8,0

1688

0,0

5328

24,0

34

2580

,0

9783

9,0

1423

08,0

53

6163

,0

1594

14,0

14

2352

,0

8268

94,0

29

4679

,0

3164

98,0

21,5

10

00,0

V

ER05

4 M

RM

ar

ea

1678

1,0

6025

4,0

3581

3,0

9585

,0

2579

44,0

17

3425

,0

4659

0,0

6634

9,0

2427

24,0

79

295,

0 68

007,

0 36

9548

,0

1133

60,0

12

7705

,025

,0

1000

,0

VER

060

MR

M

area

18

416,

0 47

891,

0 33

929,

0 81

39,0

25

8465

,0

1698

71,0

51

334,

0 65

232,

0 21

2519

,0

6963

5,0

6532

4,0

3392

97,0

12

7378

,0

1228

76,0

22,9

10

00,0

V

ER06

8 M

RM

ar

ea

1486

6,0

5368

7,0

3159

4,0

9233

,0

2860

02,0

18

1430

,0

5101

9,0

7724

2,0

1976

33,0

65

749,

0 56

254,

0 32

3705

,0

1176

00,0

12

4934

,019

,3

1000

,0

VER

076

MR

M

area

10

851,

0 45

421,

0 25

948,

0 75

69,0

24

8003

,0

1594

28,0

43

955,

0 64

809,

0 17

9415

,0

6013

8,0

5118

1,0

2875

08,0

11

3875

,0

1203

95,0

22,7

10

00,0

V

ER08

4 M

RM

ar

ea

2326

0,0

8546

4,0

5132

3,0

1934

7,0

5026

65,0

33

1708

,0

9149

4,0

1282

78,0

44

2135

,0

1348

08,0

10

9271

,0

6699

99,0

25

0365

,0

2556

38,0

20,6

10

00,0

LC

004

MR

M

area

20

592,

0 29

359,

0 43

923,

0 23

575,

0 58

8571

,0

5106

20,0

14

7489

,0

1629

89,0

57

6073

,0

8745

9,0

8594

2,0

1721

593,

0 50

7976

,0

6148

13,0

25,2

50

0,0

LC03

0 M

RM

ar

ea

4080

2,0

5992

4,0

7425

7,0

3604

2,0

1327

085,

0 11

5719

3,0

3658

58,0

36

0743

,0

8498

62,0

13

8277

,0

1537

93,0

23

8323

2,0

1188

129,

0 13

6253

9,0

25,1

50

0,0

LC03

8 M

RM

ar

ea

5277

4,0

6220

8,0

1025

63,0

47

698,

0 17

6023

4,0

1575

839,

0 50

3839

,0

5128

70,0

11

9511

8,0

1936

29,0

19

5341

,0

3252

380,

0 14

9948

0,0

1690

798,

021

,7

500,

0 LC

050

MR

M

area

24

627,

0 23

926,

0 33

332,

0 19

643,

0 74

2781

,0

6850

27,0

21

5777

,0

1929

54,0

47

0542

,0

7007

1,0

6616

8,0

1404

260,

0 58

7274

,0

6611

67,0

24,8

50

0,0

LC08

4 M

RM

ar

ea

2973

9,0

3539

0,0

5072

0,0

2655

2,0

1015

834,

0 91

1072

,0

2949

16,0

30

2239

,0

6859

17,0

10

8504

,0

1015

91,0

19

1581

5,0

7581

17,0

87

3168

,025

,2

500,

0 LC

092

MR

M

area

29

079,

0 32

104,

0 46

261,

0 20

011,

0 96

2009

,0

8558

03,0

27

9064

,0

2339

32,0

49

5076

,0

8755

2,0

9677

4,0

1344

856,

0 70

1109

,0

7628

27,0

25,4

50

0,0

LC11

0 M

RM

ar

ea

2302

7,0

2829

9,0

4015

4,0

1518

3,0

9115

62,0

83

1932

,0

2538

86,0

26

4208

,0

3923

52,0

89

301,

0 91

624,

0 10

9859

4,0

6759

87,0

81

4583

,022

,5

500,

0 LC

122

MR

M

area

17

681,

0 19

087,

0 26

509,

0 11

674,

0 62

8008

,0

5787

00,0

17

0779

,0

1676

71,0

29

4704

,0

6207

9,0

6076

2,0

8359

78,0

45

9107

,0

5515

86,0

21,2

50

0,0

LC13

4 M

RM

ar

ea

2136

1,0

2304

0,0

3123

5,0

1727

0,0

6759

18,0

61

4890

,0

1930

65,0

20

6039

,0

5036

87,0

94

266,

0 97

282,

0 14

3829

5,0

6284

21,0

73

1603

,025

,6

500,

0 LC

142

MR

M

area

41

275,

0 51

667,

0 64

659,

0 31

947,

0 11

7697

1,0

1051

042,

0 33

5200

,0

3455

50,0

93

1953

,0

1511

96,0

14

8550

,0

2580

798,

0 98

8852

,0

1141

219,

019

,5

500,

0 LC

148

MR

M

area

21

637,

0 24

190,

0 29

941,

0 13

376,

0 68

1462

,0

6165

30,0

19

7779

,0

1771

44,0

35

0800

,0

7361

8,0

6668

9,0

9993

49,0

52

2960

,0

6586

01,0

27,2

50

0,0

LC15

4 M

RM

ar

ea

4813

0,0

3221

4,0

4531

3,0

4371

4,0

2122

764,

0 19

3709

7,0

6506

70,0

58

5147

,0

1079

509,

0 23

8747

,0

2343

32,0

30

3654

5,0

1616

292,

0 20

3217

6,0

16,4

50

0,0

LC16

2 M

RM

ar

ea

8973

,0

3547

,0

7258

,0

7685

,0

2910

43,0

27

3700

,0

9670

4,0

8364

3,0

2437

51,0

32

228,

0 35

912,

0 76

8208

,0

2112

35,0

27

8920

,025

,1

500,

0 LC

170

MR

M

area

49

068,

0 37

377,

0 64

632,

0 55

559,

0 18

8516

2,0

1764

725,

0 66

5370

,0

5778

19,0

20

9595

9,0

2861

27,0

31

4442

,0

5844

193,

0 11

3912

9,0

1334

636,

024

,9

500,

0 LC

183

MR

M

area

61

277,

0 63

049,

0 10

1912

,0

7745

6,0

2048

498,

0 18

9935

0,0

7130

32,0

64

2599

,0

2979

727,

0 27

8233

,0

3291

38,0

76

8214

8,0

1103

778,

0 12

7733

1,0

25,0

10

00,0

LC

189

MR

M

area

29

806,

0 26

349,

0 46

360,

0 41

773,

0 13

0734

2,0

1236

377,

0 41

7540

,0

3878

44,0

13

1983

9,0

1630

12,0

19

3493

,0

3519

139,

0 88

1562

,0

1116

840,

0

Ster

anes

, Voc

ontia

n B

asin

, Fra

nce:

mas

s m

ikro

l. sa

mpl

e m

easu

rem

ent

ster

anes

g

µl

peak

: 12

,0

13,0

14

,0

15,0

16

,0

17,0

18

,0

19,0

20

,0

21,0

22

,0

23,0

24

,0

Typ

C

28ba

? C

28ba

? C

28ab

S?

C28

abR

? C

28aa

aS

C28

abbR

C

28ab

bS

C28

aaaR

C

29ba

S?

C29

baR

? C

29ab

S?

C29

abR

? C

29aa

aS

Alk

anm

ix 1

0MR

M

area

3,

0

MG

S-1-

1 1.

5MR

M

area

19

,8

1000

,0

VER

004

MR

M

area

17

5232

,0

2531

19,0

12

6473

,0

2043

24,0

15

4560

,0

2008

73,0

10

6222

,0

3270

39,0

72

9875

,0

4932

76,0

10

5272

,0

2044

56,0

53

9172

,022

,6

500,

0 V

ER01

2 M

RM

ar

ea

2325

71,0

25

1159

,0

1320

58,0

19

4824

,0

2145

85,0

23

1516

,0

1234

86,0

40

3378

,0

6998

86,0

49

9371

,0

1268

73,0

20

9091

,0

6896

41,0

24,1

15

00,0

V

ER02

0 M

RM

ar

ea

1919

46,0

21

1309

,0

1213

49,0

18

0737

,0

1983

97,0

23

3226

,0

1277

62,0

40

0534

,0

5953

89,0

43

0005

,0

1081

26,0

17

6405

,0

6099

63,0

25,5

10

00,0

V

ER02

2 M

RM

ar

ea

2584

89,0

31

4301

,0

1706

25,0

24

7966

,0

2584

28,0

26

4606

,0

1436

24,0

50

3318

,0

7070

08,0

47

9119

,0

1282

00,0

22

2213

,0

8277

70,0

16,3

15

00,0

V

ER02

4 M

RM

ar

ea

9829

6,0

1313

91,0

79

073,

0 10

5135

,0

1228

32,0

11

1318

,0

6367

3,0

1820

81,0

60

9559

,0

4244

42,0

11

0700

,0

1599

29,0

34

9083

,019

,5

1500

,0

VER

025

MR

M

area

10

0987

,0

1271

63,0

79

259,

0 10

3410

,0

1080

22,0

10

8359

,0

6277

4,0

1845

50,0

69

0685

,0

4614

63,0

12

7199

,0

2031

40,0

38

3008

,014

,6

1500

,0

VER

026

MR

M

area

93

720,

0 11

1896

,0

7335

1,0

1044

54,0

92

466,

0 94

322,

0 56

074,

0 14

7850

,0

6227

88,0

42

6025

,0

1181

50,0

18

9019

,0

3183

62,0

25,3

10

00,0

V

ER02

7 M

RM

ar

ea

1903

54,0

25

5138

,0

1453

83,0

21

8023

,0

2114

92,0

25

4836

,0

1410

05,0

41

2795

,0

6794

83,0

49

7182

,0

1437

04,0

21

0929

,0

6822

17,0

25,1

10

00,0

V

ER02

8 M

RM

ar

ea

2355

98,0

29

5019

,0

1673

54,0

24

4288

,0

2676

61,0

29

3466

,0

1604

44,0

47

9549

,0

8135

63,0

56

9382

,0

1572

64,0

24

7948

,0

7856

17,0

17,8

15

00,0

V

ER02

9 M

RM

ar

ea

9652

3,0

1212

67,0

62

842,

0 10

0051

,0

9809

6,0

9184

8,0

5284

2,0

1545

00,0

54

0915

,0

3575

74,0

99

596,

0 15

4993

,0

2757

26,0

20,1

10

00,0

V

ER03

2 M

RM

ar

ea

8377

7,0

1194

11,0

65

336,

0 87

128,

0 84

149,

0 91

074,

0 51

518,

0 15

5021

,0

2967

76,0

21

1718

,0

6121

1,0

8660

0,0

2753

08,0

25,9

10

00,0

V

ER03

8 M

RM

ar

ea

1657

71,0

21

6432

,0

1184

06,0

17

9958

,0

1272

80,0

14

2200

,0

7852

3,0

2363

44,0

56

6279

,0

3950

83,0

10

5239

,0

1663

03,0

41

3997

,025

,1

1000

,0

MO

R00

5 M

RM

ar

ea

1230

45,0

13

5337

,0

7599

5,0

1117

82,0

20

1503

,0

2165

58,0

11

0816

,0

4231

88,0

48

1711

,0

3338

15,0

96

212,

0 13

9180

,0

5648

45,0

24,8

10

00,0

V

ER04

4 M

RM

ar

ea

5200

3,0

6700

6,0

4449

1,0

5690

8,0

9326

6,0

1001

66,0

50

696,

0 20

0541

,0

3141

88,0

21

1937

,0

5497

8,0

8585

7,0

3054

36,0

21,3

10

00,0

V

ER05

0 M

RM

ar

ea

1655

95,0

24

7227

,0

1164

27,0

15

1271

,0

3823

86,0

36

7743

,0

1883

01,0

78

1622

,0

7629

47,0

56

0250

,0

1602

84,0

24

4994

,0

9814

97,0

21,5

10

00,0

V

ER05

4 M

RM

ar

ea

6511

7,0

9210

1,0

4585

6,0

6972

6,0

1381

49,0

14

2158

,0

7421

2,0

3008

43,0

37

6739

,0

2534

84,0

63

466,

0 10

2496

,0

4027

84,0

25,0

10

00,0

V

ER06

0 M

RM

ar

ea

7113

5,0

9258

0,0

5073

7,0

7470

3,0

1204

01,0

13

0338

,0

6794

2,0

2629

79,0

35

2312

,0

2470

79,0

63

181,

0 10

0507

,0

3688

92,0

22,9

10

00,0

V

ER06

8 M

RM

ar

ea

6553

7,0

9193

8,0

4826

2,0

6685

4,0

1073

66,0

12

1671

,0

5876

1,0

2343

14,0

39

1923

,0

2577

59,0

76

125,

0 11

5918

,0

3256

40,0

19,3

10

00,0

V

ER07

6 M

RM

ar

ea

7158

0,0

7722

9,0

4849

0,0

7052

3,0

1151

76,0

12

0503

,0

5904

2,0

2470

39,0

35

9867

,0

2418

88,0

61

242,

0 99

348,

0 29

3203

,022

,7

1000

,0

VER

084

MR

M

area

12

6764

,0

1992

85,0

10

8461

,0

1517

35,0

28

9786

,0

2952

95,0

14

3142

,0

6204

66,0

84

2521

,0

5637

29,0

13

5781

,0

2339

77,0

72

7778

,020

,6

1000

,0

LC00

4 M

RM

ar

ea

4575

35,0

53

5827

,0

2809

19,0

34

9003

,0

2732

73,0

78

1074

,0

2071

63,0

19

9282

4,0

7848

53,0

71

3103

,0

2008

18,0

28

8236

,0

3676

83,0

25,2

50

0,0

LC03

0 M

RM

ar

ea

1050

159,

0 10

6835

8,0

6641

60,0

76

1986

,0

3940

80,0

12

0701

4,0

4104

07,0

26

0148

1,0

1874

688,

0 17

1491

8,0

4803

44,0

61

1463

,0

5332

20,0

25,1

50

0,0

LC03

8 M

RM

ar

ea

1403

003,

0 13

5119

1,0

8973

11,0

10

0738

3,0

5865

56,0

15

7183

8,0

5056

41,0

35

3492

2,0

2612

810,

0 24

9318

5,0

6764

17,0

97

4305

,0

7821

17,0

21,7

50

0,0

LC05

0 M

RM

ar

ea

5192

36,0

59

3729

,0

3356

86,0

35

8172

,0

1973

04,0

60

6547

,0

1866

65,0

13

1668

0,0

1096

615,

0 96

7683

,0

2421

72,0

36

4203

,0

2841

65,0

24,8

50

0,0

LC08

4 M

RM

ar

ea

6008

64,0

80

5464

,0

4410

77,0

49

9516

,0

2151

84,0

80

7033

,0

2731

04,0

18

0269

9,0

1497

068,

0 14

2448

4,0

3844

80,0

46

5134

,0

3652

43,0

25,2

50

0,0

LC09

2 M

RM

ar

ea

4861

44,0

74

5998

,0

3849

82,0

44

1869

,0

1794

63,0

59

4518

,0

2064

72,0

12

0184

5,0

1482

427,

0 13

8520

4,0

3612

21,0

43

9196

,0

2897

57,0

25,4

50

0,0

LC11

0 M

RM

ar

ea

5984

85,0

70

3933

,0

3925

87,0

42

5788

,0

1349

87,0

46

8191

,0

1675

06,0

90

0170

,0

1397

866,

0 12

3912

4,0

3266

33,0

45

9507

,0

2220

76,0

22,5

50

0,0

LC12

2 M

RM

ar

ea

4226

67,0

48

4742

,0

2747

36,0

28

9102

,0

9276

8,0

3562

05,0

13

1702

,0

7079

95,0

97

3341

,0

9174

99,0

23

3875

,0

2997

18,0

14

7243

,021

,2

500,

0 LC

134

MR

M

area

57

1106

,0

5696

84,0

36

4234

,0

21

5726

,0

7187

73,0

21

2635

,0

1593

541,

0 10

8419

3,0

9982

12,0

24

8363

,0

3297

29,0

29

9296

,025

,6

500,

0 LC

142

MR

M

area

89

3830

,0

9620

89,0

59

5838

,0

6376

02,0

39

2113

,0

1361

466,

0 43

5183

,0

3037

752,

0 15

2151

8,0

1414

655,

0 38

4988

,0

5109

29,0

44

8414

,019

,5

500,

0 LC

148

MR

M

area

45

5763

,0

5688

80,0

31

7579

,0

3575

54,0

11

8734

,0

4732

60,0

15

5707

,0

1047

892,

0 94

7011

,0

9032

62,0

25

1422

,0

2903

20,0

18

8603

,027

,2

500,

0 LC

154

MR

M

area

15

0717

6,0

1446

810,

0 10

1762

1,0

1149

661,

0 43

0544

,0

1422

902,

0 54

9108

,0

2820

251,

0 33

2767

1,0

3175

999,

0 86

4446

,0

1090

067,

0 64

6733

,016

,4

500,

0 LC

162

MR

M

area

20

0408

,0

2184

00,0

13

7641

,0

1433

12,0

82

964,

0 26

1439

,0

7947

4,0

6411

90,0

41

9858

,0

4028

96,0

96

704,

0 13

0114

,0

1209

46,0

25,1

50

0,0

LC17

0 M

RM

ar

ea

1021

574,

0 14

1398

0,0

7589

72,0

84

3770

,0

7875

35,0

21

2118

3,0

4793

94,0

54

8245

2,0

2032

828,

0 21

3115

3,0

74

0566

,0

1031

688,

024

,9

500,

0 LC

183

MR

M

area

97

1096

,0

1197

781,

0 71

2535

,0

8653

78,0

10

9784

8,0

3043

466,

0 66

7595

,0

7875

151,

0 19

7888

4,0

1847

972,

0 49

0183

,0

7162

84,0

14

6297

7,0

25,0

10

00,0

LC

189

MR

M

area

86

3116

,0

8892

91,0

59

3958

,0

6549

13,0

48

0453

,0

1427

123,

0 33

3055

,0

3581

914,

0 15

3298

2,0

1480

633,

0 42

8691

,0

5266

73,0

59

1951

,0

154

155

Appendix

Ster

anes

, Voc

ontia

n B

asin

, Fra

nce:

mas

s m

ikro

l. sa

mpl

e m

easu

rem

ent

ster

anes

g µ

l

pe

ak:

25,0

26

,0

27,0

28

,0

29,0

30

,0

31,0

32

,0

33,0

?

34,0

35

,0

Typ

C

29ab

bS?

C29

abbR

? C

29aa

aR

C30

S?

C30

R?

MC

29?

MC

29?

MC

29?

MC

29?

? M

C29

? M

C29

?

A

lkan

mix

10M

RM

ar

ea

3,

0

MG

S-1-

1 1.

5MR

M

area

19,8

10

00,0

V

ER00

4 M

RM

ar

ea

2155

12,0

40

8987

,0

8938

37,0

40

421,

0 84

139,

0 76

840,

0 53

999,

0 20

63,0

61

95,0

4767

4,0

3920

,022

,6

500,

0 V

ER01

2 M

RM

ar

ea

3243

28,0

58

3571

,0

1128

145,

0 50

490,

0 97

890,

0 81

467,

0 78

044,

0 18

08,0

60

72,0

5520

3,0

5205

,024

,1

1500

,0

VER

020

MR

M

area

31

9321

,0

4527

04,0

10

2153

7,0

4987

0,0

1062

62,0

71

346,

0 73

369,

0 20

84,0

52

73,0

5349

3,0

5939

,025

,5

1000

,0

VER

022

MR

M

area

35

0335

,0

5702

93,0

13

2091

7,0

6443

5,0

1274

17,0

79

688,

0 82

856,

0 24

23,0

75

08,0

7067

1,0

8612

,016

,3

1500

,0

VER

024

MR

M

area

15

9130

,0

2139

59,0

52

9092

,0

1864

1,0

3845

2,0

4509

4,0

3521

2,0

1530

,0

3030

,0

24

949,

0 20

83,0

19,5

15

00,0

V

ER02

5 M

RM

ar

ea

1711

69,0

23

1931

,0

5469

45,0

20

937,

0 36

281,

0 47

638,

0 35

745,

0 10

55,0

26

49,0

2359

8,0

1851

,014

,6

1500

,0

VER

026

MR

M

area

12

2809

,0

1861

22,0

41

9827

,0

1466

2,0

2713

0,0

4243

2,0

2927

3,0

1159

,0

2089

,0

19

225,

0 19

68,0

25,3

10

00,0

V

ER02

7 M

RM

ar

ea

3565

22,0

47

9249

,0

1079

691,

0 58

233,

0 11

6963

,0

8091

1,0

6844

3,0

3569

,0

5530

,0

59

339,

0 81

54,0

25,1

10

00,0

V

ER02

8 M

RM

ar

ea

4793

10,0

57

8932

,0

1257

201,

0 62

841,

0 13

1767

,0

9571

7,0

8736

2,0

3637

,0

6491

,0

73

313,

0 97

52,0

17,8

15

00,0

V

ER02

9 M

RM

ar

ea

1205

67,0

17

9910

,0

4029

43,0

14

836,

0 26

657,

0 36

937,

0 26

188,

0 98

7,0

1971

,0

17

329,

0 20

37,0

20,1

10

00,0

V

ER03

2 M

RM

ar

ea

1076

43,0

20

2800

,0

4475

41,0

21

763,

0 39

375,

0 32

963,

0 33

841,

0 13

26,0

29

35,0

2556

5,0

3610

,025

,9

1000

,0

VER

038

MR

M

area

25

0492

,0

2857

95,0

68

2743

,0

2952

8,0

5920

3,0

5417

6,0

5291

6,0

1885

,0

5363

,0

43

064,

0 53

49,0

25,1

10

00,0

M

OR

005

MR

M

area

24

7200

,0

4445

70,0

10

6066

7,0

4308

3,0

9835

1,0

6790

6,0

8211

7,0

4743

,0

6533

,0

59

376,

0 90

00,0

24,8

10

00,0

V

ER04

4 M

RM

ar

ea

1190

58,0

24

3257

,0

5898

59,0

24

678,

0 51

851,

0 44

060,

0 44

563,

0 13

05,0

42

52,0

3137

7,0

4131

,021

,3

1000

,0

VER

050

MR

M

area

39

8438

,0

7207

89,0

17

4130

8,0

7277

2,0

1588

52,0

98

329,

0 12

0009

,0

4892

,0

8557

,0

88

476,

0 13

667,

021

,5

1000

,0

VER

054

MR

M

area

15

4298

,0

3497

57,0

76

3590

,0

3590

1,0

7429

2,0

5390

3,0

5871

6,0

2377

,0

4832

,0

43

437,

0 53

90,0

25,0

10

00,0

V

ER06

0 M

RM

ar

ea

1639

34,0

28

2643

,0

6999

40,0

26

523,

0 56

648,

0 51

467,

0 54

071,

0 10

82,0

49

62,0

3740

0,0

5556

,022

,9

1000

,0

VER

068

MR

M

area

13

8952

,0

2679

08,0

62

2366

,0

2533

7,0

5606

1,0

5101

6,0

5575

7,0

2576

,0

6494

,0

38

576,

0 52

96,0

19,3

10

00,0

V

ER07

6 M

RM

ar

ea

1462

21,0

23

8236

,0

6012

32,0

29

618,

0 66

401,

0 45

454,

0 52

696,

0 20

39,0

42

60,0

3895

8,0

5962

,022

,7

1000

,0

VER

084

MR

M

area

31

7675

,0

5906

28,0

13

8076

1,0

6877

1,0

1542

24,0

10

2058

,0

1090

69,0

41

82,0

90

42,0

8297

3,0

1052

7,0

20,6

10

00,0

LC

004

MR

M

area

32

7377

,0

1361

142,

0 39

9731

9,0

4059

5,0

4950

03,0

65

050,

0 24

1012

,0

4415

3,0

9877

,0

6148

2,0

4215

86,0

67

057,

025

,2

500,

0 LC

030

MR

M

area

61

8226

,0

1844

746,

0 52

4913

8,0

4870

8,0

5960

36,0

12

8635

,0

4056

81,0

63

736,

0 18

726,

0 71

008,

0 60

9619

,0

9066

7,0

25,1

50

0,0

LC03

8 M

RM

ar

ea

4766

85,0

28

9028

8,0

7481

029,

0 10

0104

,0

9092

65,0

16

7371

,0

5313

63,0

84

666,

0 47

957,

0 96

382,

0 77

5858

,0

8999

0,0

21,7

50

0,0

LC05

0 M

RM

ar

ea

3060

91,0

11

9272

9,0

3169

035,

0 25

626,

0 31

7734

,0

8308

6,0

2608

29,0

33

999,

0 13

202,

0 37

337,

0 36

4774

,0

3902

6,0

24,8

50

0,0

LC08

4 M

RM

ar

ea

4701

54,0

15

5535

8,0

4318

295,

0 26

882,

0 41

4608

,0

1105

69,0

40

7407

,0

4506

5,0

1326

0,0

5332

5,0

5381

11,0

63

457,

025

,2

500,

0 LC

092

MR

M

area

34

8403

,0

1193

061,

0 30

1021

1,0

2336

5,0

2983

53,0

94

383,

0 27

9999

,0

2598

9,0

1005

4,0

3935

5,0

3778

78,0

31

283,

025

,4

500,

0 LC

110

MR

M

area

34

8703

,0

9788

94,0

24

1942

6,0

1534

5,0

2387

99,0

79

449,

0 25

8320

,0

2535

4,0

1354

5,0

3670

5,0

3214

18,0

23

876,

022

,5

500,

0 LC

122

MR

M

area

19

1140

,0

7623

33,0

18

1885

6,0

1420

6,0

1494

71,0

64

189,

0 20

4234

,0

1746

2,0

8035

,0

2508

2,0

2287

55,0

21

327,

021

,2

500,

0 LC

134

MR

M

area

28

6691

,0

1300

362,

0 34

0764

6,0

2726

0,0

2930

56,0

85

157,

0 30

9790

,0

2935

5,0

1751

8,0

3540

6,0

3718

34,0

39

068,

025

,6

500,

0 LC

142

MR

M

area

53

3558

,0

1821

272,

0 49

3448

5,0

5284

2,0

5785

46,0

14

0375

,0

4628

68,0

99

947,

0 27

161,

0 11

2796

,0

8392

78,0

13

5165

,019

,5

500,

0 LC

148

MR

M

area

15

3323

,0

8323

61,0

20

9799

7,0

1374

0,0

2247

75,0

58

904,

0 18

2566

,0

2286

5,0

8481

,0

3297

1,0

2767

78,0

29

523,

027

,2

500,

0 LC

154

MR

M

area

57

3768

,0

2660

767,

0 63

8003

2,0

6931

3,0

7209

82,0

19

7773

,0

5188

00,0

84

203,

0 26

703,

0 15

0433

,0

9194

61,0

10

1691

,016

,4

500,

0 LC

162

MR

M

area

11

7998

,0

4955

95,0

15

1410

7,0

8941

,0

1256

17,0

28

430,

0 85

714,

0 12

754,

0 41

21,0

21

962,

0 15

6234

,0

1359

6,0

25,1

50

0,0

LC17

0 M

RM

ar

ea

6698

34,0

39

2811

9,0

1045

4594

,0 1

0650

4,0

1275

934,

0 17

5747

,0

5303

93,0

15

1685

,0

3930

3,0

2560

26,0

10

5099

0,0

1740

82,0

24,9

50

0,0

LC18

3 M

RM

ar

ea

1045

448,

0 41

0914

5,0

1274

8099

,0 1

2716

9,0

1569

666,

0 26

8614

,0

5306

41,0

31

4726

,0

7386

5,0

3599

22,0

18

2464

6,0

2742

61,0

25,0

10

00,0

LC

189

MR

M

area

34

3248

,0

1997

751,

0 59

4649

1,0

6378

6,0

7348

95,0

97

018,

0 26

3504

,0

6055

3,0

2152

3,0

1159

66,0

51

3959

,0

6163

3,0

156

Selected chromatograms France:VER32

VER84

LC50

Appendix

157

LC122

LC170

Selected chromatograms Poland

PIG1.28

158

PIG1.45

PIG1.91

PIG1.114

Appendix

159

PIG1.159

Cou

nts s

pore

-pol

len

Mid

-Pol

ish

Trou

gh, P

olan

d:

dept

h:

30

50

76

102

143

183

220

250

282

320

360

420

460

500

530

560

600

taxa

: sa

mpl

e:

PIG

1_5

PIG

1_8

PIG

1_11

PI

G1_

13

PIG

1_17

PI

G1_

21

PIG

1_25

PI

G1_

28

PIG

1_31

PI

G1_

35

PIG

1_39

PI

G1_

45

PIG

1_49

PI

G1_

53

PIG

1_56

PI

G1_

59

PIG

1_63

Parv

isac

cite

s 0

0 2

0 0

0 0

4 1

0 0

0 2

6 3

0 2

%

0,

0 0,

0 0,

7 0,

0 0,

0 0,

0 0,

0 1,

4 0,

3 0,

0 0,

0 0,

0 0,

7 2,

0 1,

0 0,

0 0,

7Po

doca

rpid

ites

0 0

1 0

1 0

0 0

0 0

0 0

1 0

0 0

0%

0,0

0,0

0,3

0,0

0,3

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,3

0,0

0,0

0,0

0,0

Pinu

spol

leni

tes

1 0

0 2

0 5

0 4

0 1

0 4

0 1

1 0

2%

0,3

0,0

0,0

0,7

0,0

1,6

0,0

1,4

0,0

0,3

0,0

1,3

0,0

0,3

0,3

0,0

0,7

bisa

ccat

e in

diff.

32

30

71

16

79

25

61

54

50

32

48

32

41

37

26

55

33

%

10

,7

10,0

23

,6

5,3

26,3

8,

1 20

,3

18,4

16

,7

10,6

15

,9

10,5

13

,7

12,3

8,

7 18

,3

11,0

Ara

ucar

iaci

tes

16

3 10

40

8

4 8

14

9 7

15

10

12

12

0 11

14

%

5,

3 1,

0 3,

3 13

,2

2,7

1,3

2,7

4,8

3,0

2,3

5,0

3,3

4,0

4,0

0,0

3,7

4,7

Cal

lialla

spor

ites

20

1 14

21

10

18

18

48

9

24

9 34

6

6 0

13

32%

6,7

0,3

4,7

7,0

3,3

5,8

6,0

16,4

3,

0 7,

9 3,

0 11

,1

2,0

2,0

0,0

4,3

10,7

Inap

ertu

ropo

lleni

tes

14

3 18

14

15

14

39

18

34

16

20

15

21

11

1

11

20%

4,7

1,0

6,0

4,6

5,0

4,5

13,0

6,

1 11

,3

5,3

6,6

4,9

7,0

3,7

0,3

3,7

6,7

Perin

opol

leni

tes

8 1

22

19

18

29

33

21

17

16

10

41

36

44

0 43

7

%

2,

7 0,

3 7,

3 6,

3 6,

0 9,

4 11

,0

7,2

5,7

5,3

3,3

13,4

12

,0

14,6

0,

0 14

,3

2,3

Cer

ebro

polle

nite

s 15

12

10

8

10

24

12

7 28

14

23

18

12

24

5

18

14%

5,0

4,0

3,3

2,6

3,3

7,7

4,0

2,4

9,3

4,6

7,6

5,9

4,0

8,0

1,7

6,0

4,7

Cla

ssop

ollis

16

22

0 40

3

28

8 13

2

6 6

6 7

14

14

165

7 7

%

5,

3 73

,1

13,3

1,

0 9,

3 2,

6 4,

3 0,

7 2,

0 2,

0 2,

0 2,

3 4,

7 4,

7 55

,0

2,3

2,3

Exes

ipol

leni

tes

2 1

0 1

1 1

1 0

0 0

0 0

1 0

1 0

2%

0,7

0,3

0,0

0,3

0,3

0,3

0,3

0,0

0,0

0,0

0,0

0,0

0,3

0,0

0,3

0,0

0,7

Cyc

adop

ites

13

1 3

21

3 20

3

5 0

8 3

6 0

1 8

5 7

%

4,

3 0,

3 1,

0 7,

0 1,

0 6,

5 1,

0 1,

7 0,

0 2,

6 1,

0 2,

0 0,

0 0,

3 2,

7 1,

7 2,

3Eu

com

miid

ites

6 0

0 1

0 2

0 1

0 0

0 1

0 4

17

0 1

%

2,

0 0,

0 0,

0 0,

3 0,

0 0,

6 0,

0 0,

3 0,

0 0,

0 0,

0 0,

3 0,

0 1,

3 5,

7 0,

0 0,

3A

lispo

rites

6 0

1 10

4

3 0

1 1

0 0

2 3

4 17

0

0%

2,0

0,0

0,3

3,3

1,3

1,0

0,0

0,3

0,3

0,0

0,0

0,7

1,0

1,3

5,7

0,0

0,0

Vitr

eisp

orite

s pal

lidus

1

1 5

4 4

1 5

0 2

4 1

0 2

5 0

1 5

%

0,

3 0,

3 1,

7 1,

3 1,

3 0,

3 1,

7 0,

0 0,

7 1,

3 0,

3 0,

0 0,

7 1,

7 0,

0 0,

3 1,

7Ep

hedr

epite

s 3

0 0

1 0

2 1

0 0

1 1

2 0

0 0

0 6

%

1,

0 0,

0 0,

0 0,

3 0,

0 0,

6 0,

3 0,

0 0,

0 0,

3 0,

3 0,

7 0,

0 0,

0 0,

0 0,

0 2,

0C

yath

idite

s

15

3 26

11

51

23

49

16

54

38

65

21

33

22

1

18

49%

5,0

1,0

8,6

3,6

17,0

7,

4 16

,3

5,5

18,0

12

,5

21,6

6,

9 11

,0

7,3

0,3

6,0

16,3

Leio

trile

tes

24

1

15

14

12

24

14

3 17

10

22

15

13

11

0

22

19%

8,0

0,3

5,0

4,6

4,0

7,7

4,7

1,0

5,7

3,3

7,3

4,9

4,3

3,7

0,0

7,3

6,3

Con

cavi

spor

ites

6 0

0 2

1 1

0 0

0 2

0 2

1 1

0 1

3%

2,0

0,0

0,0

0,7

0,3

0,3

0,0

0,0

0,0

0,7

0,0

0,7

0,3

0,3

0,0

0,3

1,0

Con

cavi

ssim

ispo

rites

0

0 1

0 0

1 0

0 0

0 0

0 0

0 0

0 0

%

0,

0 0,

0 0,

3 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0Tr

ilobo

spor

ites

0 0

0 1

1 2

0 3

0 1

0 1

3 0

0 2

2%

0,0

0,0

0,0

0,3

0,3

0,6

0,0

1,0

0,0

0,3

0,0

0,3

1,0

0,0

0,0

0,7

0,7

Cic

atric

osis

porit

es

9 0

1 8

2 4

0 6

2 5

4 6

4 4

0 4

0%

3,0

0,0

0,3

2,6

0,7

1,3

0,0

2,0

0,7

1,7

1,3

2,0

1,3

1,3

0,0

1,3

0,0

Klu

kisp

orite

s 0

0 0

0 0

0 0

0 0

0 1

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,

3 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0Is

chyo

spor

ites

0 0

0 0

5 0

1 0

0 2

1 1

0 0

0 2

1%

0,0

0,0

0,0

0,0

1,7

0,0

0,3

0,0

0,0

0,7

0,3

0,3

0,0

0,0

0,0

0,7

0,3

Osm

unda

cidi

tes

4 0

0 4

1 10

1

4 1

7 4

2 0

3 0

1 2

%

1,

3 0,

0 0,

0 1,

3 0,

3 3,

2 0,

3 1,

4 0,

3 2,

3 1,

3 0,

7 0,

0 1,

0 0,

0 0,

3 0,

7

160

Cou

nts s

pore

-pol

len

Mid

-Pol

ish

Trou

gh, P

olan

d:

dept

h:

30

50

76

102

143

183

220

250

282

320

360

420

460

500

530

560

600

taxa

: sa

mpl

e:

PIG

1_5

PIG

1_8

PIG

1_11

PI

G1_

13

PIG

1_17

PI

G1_

21

PIG

1_25

PI

G1_

28

PIG

1_31

PI

G1_

35

PIG

1_39

PI

G1_

45

PIG

1_49

PI

G1_

53

PIG

1_56

PI

G1_

59

PIG

1_63

Gle

iche

niid

ites

14

0 12

11

17

15

11

15

27

20

29

16

12

14

3

11

29%

4,7

0,0

4,0

3,6

5,7

4,8

3,7

5,1

9,0

6,6

9,6

5,2

4,0

4,7

1,0

3,7

9,7

Cla

vatis

porit

es

0 0

3 0

0 0

0 0

0 1

1 0

0 1

0 0

0%

0,0

0,0

1,0

0,0

0,0

0,0

0,0

0,0

0,0

0,3

0,3

0,0

0,0

0,3

0,0

0,0

0,0

delto

ide

Spor

e in

diff.

60

9

22

58

19

55

15

42

34

71

26

50

55

66

0 65

31

%

20

,0

3,0

7,3

19,2

6,

3 17

,7

5,0

14,3

11

,3

23,4

8,

6 16

,3

18,4

21

,9

0,0

21,7

10

,3To

disp

orite

s 0

0 0

0 0

1 0

0 0

0 0

1 0

0 0

0 0

%

0,

0 0,

0 0,

0 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

0 0,

0Ve

rruc

osis

porit

es

0 0

0 0

0 0

0 0

0 0

0 0

1 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

0,3

0,0

0,0

0,0

0,0

Gem

mat

rilet

es

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

0,0

0,0

0,0

0,0

0,0

0,0

0,0

Bire

tispo

rites

0

0 7

0 3

1 1

1 2

0 2

0 2

1 0

1 0

%

0,

0 0,

0 2,

3 0,

0 1,

0 0,

3 0,

3 0,

3 0,

7 0,

0 0,

7 0,

0 0,

7 0,

3 0,

0 0,

3 0,

0N

eora

istri

ckia

1

0 1

0 1

3 5

2 2

1 4

0 3

1 0

2 0

%

0,

3 0,

0 0,

3 0,

0 0,

3 1,

0 1,

7 0,

7 0,

7 0,

3 1,

3 0,

0 1,

0 0,

3 0,

0 0,

7 0,

0R

etril

etes

0 0

0 0

0 1

1 0

0 0

0 0

0 0

0 0

0%

0,0

0,0

0,0

0,0

0,0

0,3

0,3

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

Aeq

uitri

radi

tes

0 0

2 0

0 0

0 0

0 0

0 0

2 0

0 0

0%

0,0

0,0

0,7

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,7

0,0

0,0

0,0

0,0

Lept

olep

idite

s 1

0 3

4 1

2 0

3 0

1 1

0 4

3 0

2 1

%

0,

3 0,

0 1,

0 1,

3 0,

3 0,

6 0,

0 1,

0 0,

0 0,

3 0,

3 0,

0 1,

3 1,

0 0,

0 0,

7 0,

3C

amar

ozon

ospo

rites

0

0 0

1 0

0 1

0 0

0 0

0 1

0 0

0 0

%

0,

0 0,

0 0,

0 0,

3 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

0D

enso

ipor

ites

0 0

0 0

0 1

0 1

0 0

0 0

0 0

0 0

0%

0,0

0,0

0,0

0,0

0,0

0,3

0,0

0,3

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

Fove

ospo

rites

0

0 0

1 0

0 0

0 0

0 0

0 0

0 0

0 0

%

0,

0 0,

0 0,

0 0,

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,

0Ly

copo

diac

idite

s 0

0 0

0 0

0 1

0 0

0 0

1 0

0 0

0 0

%

0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

0 0,

0A

ratri

spor

ites

0 5

0 0

0 0

0 0

0 0

0 0

0 0

21

0 0

%

0,

0 1,

7 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 7,

0 0,

0 0,

0Ly

copo

dium

spor

ites

0 0

0 0

0 0

0 0

0 0

1 0

0 1

0 0

0%

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,3

0,0

0,0

0,3

0,0

0,0

0,0

Echi

natis

poris

0

0 3

2 0

0 0

0 1

0 3

1 7

2 0

2 3

%

0,

0 0,

0 1,

0 0,

7 0,

0 0,

0 0,

0 0,

0 0,

3 0,

0 1,

0 0,

3 2,

3 0,

7 0,

0 0,

7 1,

0Fo

ram

inis

poris

0

0 1

0 1

0 3

0 1

0 0

1 6

0 0

0 0

%

0,

0 0,

0 0,

3 0,

0 0,

3 0,

0 1,

0 0,

0 0,

3 0,

0 0,

0 0,

3 2,

0 0,

0 0,

0 0,

0 0,

0Sp

ore

mon

olet

e 2

0 0

6 0

1 0

0 0

0 0

0 0

0 0

0 4

%

0,

7 0,

0 0,

0 2,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 1,

3St

aplin

ispo

rites

0

0 3

0 0

0 0

0 0

1 0

0 0

0 0

0 0

%

0,

0 0,

0 1,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0M

icro

retic

ulat

ispo

ris

0 0

0 0

0 1

0 0

0 0

0 2

0 0

0 0

0%

0,0

0,0

0,0

0,0

0,0

0,3

0,0

0,0

0,0

0,0

0,0

0,7

0,0

0,0

0,0

0,0

0,0

Ster

eisp

orite

s 0

1 1

1 0

0 0

2 0

2 1

2 0

1 0

0 2

%

0,

0 0,

3 0,

3 0,

3 0,

0 0,

0 0,

0 0,

7 0,

0 0,

7 0,

3 0,

7 0,

0 0,

3 0,

0 0,

0 0,

7O

void

ites

9

9 1

14

4 8

3 14

2

12

0 10

0

0 31

3

2%

3,0

3,0

0,3

4,6

1,3

2,6

1,0

4,8

0,7

4,0

0,0

3,3

0,0

0,0

10,3

1,

0 0,

7Le

iosp

haer

idia

2

0 2

3 0

0 0

2 0

0 0

2 1

1 0

0 0

%

0,

7 0,

0 0,

7 1,

0 0,

0 0,

0 0,

0 0,

7 0,

0 0,

0 0,

0 0,

7 0,

3 0,

3 0,

0 0,

0 0,

0

161

Appendix

Cou

nts s

pore

-pol

len

Mid

-Pol

ish

Trou

gh, P

olan

d:

dept

h:

640

680

710

752

790

830

880

910

950

990

1030

10

70

1110

11

50

1190

12

20

1250

taxa

: sa

mpl

e:

PIG

1_67

PI

G1_

71

PIG

1_74

PI

G1_

78

PIG

1_82

PI

G1_

86

PIG

1_91

PI

G1_

94

PIG

1_98

PI

G1_

102

PIG

1_10

6 PI

G1_

110

PIG

1_11

4 PI

G1_

118

PIG

1_12

2 PI

G1_

127

PIG

1_13

1Pa

rvis

acci

tes

0 1

0 0

0 0

0 0

2 0

1 4

0 0

1 0

0%

0,0

0,3

0,0

0,0

0,0

0,0

0,0

0,0

0,7

0,0

0,3

1,3

0,0

0,0

0,3

0,0

0,0

Podo

carp

idite

s 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 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0Pi

nusp

olle

nite

s 0

5 2

0 0

2 2

1 0

0 5

0 6

1 4

1 2

%

0,

0 1,

6 0,

7 0,

0 0,

0 0,

7 0,

7 0,

3 0,

0 0,

0 1,

7 0,

0 2,

0 0,

3 1,

3 0,

3 0,

7bi

sacc

ate

indi

ff.

23

41

35

31

27

23

46

20

34

32

45

28

33

24

6 22

28

%

7,

6 13

,4

11,6

10

,3

9,0

7,7

15,5

6,

6 11

,3

10,7

14

,9

9,3

10,8

8,

0 2,

0 7,

3 9,

4A

rauc

aria

cite

s 16

8

25

5 11

8

22

18

8 22

11

17

10

14

13

27

14

%

5,

3 2,

6 8,

3 1,

7 3,

7 2,

7 7,

4 6,

0 2,

7 7,

3 3,

6 5,

7 3,

3 4,

7 4,

3 9,

0 4,

7C

allia

llasp

orite

s 5

23

6 31

12

28

21

22

39

12

26

20

30

17

42

30

29

%

1,

7 7,

5 2,

0 10

,3

4,0

9,4

7,1

7,3

13,0

4,

0 8,

6 6,

7 9,

8 5,

7 14

,0

10,0

9,

7In

aper

turo

polle

nite

s 17

19

19

33

26

23

20

16

5

27

18

30

9 22

6

8 6

%

5,

6 6,

2 6,

3 11

,0

8,6

7,7

6,7

5,3

1,7

9,0

6,0

10,0

3,

0 7,

3 2,

0 2,

7 2,

0Pe

rinop

olle

nite

s 50

14

38

23

13

7

26

12

1 5

14

32

12

21

6 5

4%

16,6

4,

6 12

,6

7,7

4,3

2,3

8,8

4,0

0,3

1,7

4,6

10,7

3,

9 7,

0 2,

0 1,

7 1,

3C

ereb

ropo

lleni

tes

29

15

22

5 10

12

5

14

7 6

11

8 16

8

5 7

2%

9,6

4,9

7,3

1,7

3,3

4,0

1,7

4,7

2,3

2,0

3,6

2,7

5,2

2,7

1,7

2,3

0,7

Cla

ssop

ollis

9

7 13

12

7

12

9 26

9

23

6 11

6

34

1 8

4%

3,0

2,3

4,3

4,0

2,3

4,0

3,0

8,6

3,0

7,7

2,0

3,7

2,0

11,3

0,

3 2,

7 1,

3Ex

esip

olle

nite

s 2

0 2

0 0

0 0

2 2

0 1

0 5

0 0

0 1

%

0,

7 0,

0 0,

7 0,

0 0,

0 0,

0 0,

0 0,

7 0,

7 0,

0 0,

3 0,

0 1,

6 0,

0 0,

0 0,

0 0,

3%

50,0

43

,5

53,8

46

,7

35,2

38

,6

50,8

43

,5

35,7

42

,3

45,7

50

,0

41,6

47

,0

28,0

36

,0

30,1

Cyc

adop

ites

2 6

3 11

9

8 7

2 9

5 9

8 7

10

3 4

6%

0,7

2,0

1,0

3,7

3,0

2,7

2,4

0,7

3,0

1,7

3,0

2,7

2,3

3,3

1,0

1,3

2,0

Euco

mm

iidite

s 0

0 0

2 0

1 0

2 0

0 2

4 0

5 1

0 2

%

0,

0 0,

0 0,

0 0,

7 0,

0 0,

3 0,

0 0,

7 0,

0 0,

0 0,

7 1,

3 0,

0 1,

7 0,

3 0,

0 0,

7A

lispo

rites

0 1

1 1

0 0

0 2

5 2

4 0

2 3

0 2

3%

0,0

0,3

0,3

0,3

0,0

0,0

0,0

0,7

1,7

0,7

1,3

0,0

0,7

1,0

0,0

0,7

1,0

Vitr

eisp

orite

s pal

lidus

9

2 0

6 5

5 2

7 4

3 3

6 4

7 4

9 4

%

3,

0 0,

7 0,

0 2,

0 1,

7 1,

7 0,

7 2,

3 1,

3 1,

0 1,

0 2,

0 1,

3 2,

3 1,

3 3,

0 1,

3Ep

hedr

epite

s 0

7 0

0 0

8 10

2

3 0

3 0

3 1

4 0

0%

0,0

2,3

0,0

0,0

0,0

2,7

3,4

0,7

1,0

0,0

1,0

0,0

1,0

0,3

1,3

0,0

0,0

Cya

thid

ites

19

30

28

38

40

24

34

17

29

26

12

19

18

17

29

31

25

%

6,

3 9,

8 9,

3 12

,7

13,3

8,

1 11

,4

5,6

9,7

8,7

4,0

6,3

5,9

5,7

9,7

10,3

8,

4Le

iotri

lete

s

18

15

11

15

29

11

8 18

18

21

13

16

14

7

9 15

42

%

6,

0 4,

9 3,

7 5,

0 9,

6 3,

7 2,

7 6,

0 6,

0 7,

0 4,

3 5,

3 4,

6 2,

3 3,

0 5,

0 14

,0C

onca

visp

orite

s 0

0 0

3 2

4 1

1 1

4 1

2 2

1 4

6 1

%

0,

0 0,

0 0,

0 1,

0 0,

7 1,

3 0,

3 0,

3 0,

3 1,

3 0,

3 0,

7 0,

7 0,

3 1,

3 2,

0 0,

3C

onca

viss

imis

porit

es

1 0

0 0

0 1

0 0

1 0

0 0

3 0

0 0

1%

0,3

0,0

0,0

0,0

0,0

0,3

0,0

0,0

0,3

0,0

0,0

0,0

1,0

0,0

0,0

0,0

0,3

Trilo

bosp

orite

s 6

1 8

1 0

2 0

0 0

0 0

5 0

0 2

1 0

%

2,

0 0,

3 2,

7 0,

3 0,

0 0,

7 0,

0 0,

0 0,

0 0,

0 0,

0 1,

7 0,

0 0,

0 0,

7 0,

3 0,

0C

icat

ricos

ispo

rites

5

6 6

2 1

7 6

4 4

7 5

2 3

0 7

6 8

%

1,

7 2,

0 2,

0 0,

7 0,

3 2,

3 2,

0 1,

3 1,

3 2,

3 1,

7 0,

7 1,

0 0,

0 2,

3 2,

0 2,

7K

luki

spor

ites

0 0

0 0

0 0

0 0

0 0

1 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

0,3

0,0

0,0

0,0

0,0

0,0

0,0

Isch

yosp

orite

s 3

0 0

0 0

1 1

1 1

1 0

0 1

0 0

0 1

%

1,

0 0,

0 0,

0 0,

0 0,

0 0,

3 0,

3 0,

3 0,

3 0,

3 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

3O

smun

daci

dite

s 4

8 2

2 7

3 1

5 0

6 2

5 6

5 4

5 0

%

1,

3 2,

6 0,

7 0,

7 2,

3 1,

0 0,

3 1,

7 0,

0 2,

0 0,

7 1,

7 2,

0 1,

7 1,

3 1,

7 0,

0

162

Cou

nts s

pore

-pol

len

Mid

-Pol

ish

Trou

gh, P

olan

d:

dept

h:

640

680

710

752

790

830

880

910

950

990

1030

10

70

1110

11

50

1190

12

20

1250

taxa

: sa

mpl

e:

PIG

1_67

PI

G1_

71

PIG

1_74

PI

G1_

78

PIG

1_82

PI

G1_

86

PIG

1_91

PI

G1_

94

PIG

1_98

PI

G1_

102

PIG

1_10

6 PI

G1_

110

PIG

1_11

4 PI

G1_

118

PIG

1_12

2 PI

G1_

127

PIG

1_13

1G

leic

heni

idite

s 7

13

12

22

5 21

26

7

24

9 14

6

16

6 30

15

22

%

2,

3 4,

2 4,

0 7,

3 1,

7 7,

0 8,

8 2,

3 8,

0 3,

0 4,

6 2,

0 5,

2 2,

0 10

,0

5,0

7,4

Cla

vatis

porit

es

0 0

1 0

0 0

1 0

0 2

0 0

0 0

0 0

0%

0,0

0,0

0,3

0,0

0,0

0,0

0,3

0,0

0,0

0,7

0,0

0,0

0,0

0,0

0,0

0,0

0,0

delto

ide

Spor

e in

diff.

71

62

54

39

76

68

15

88

65

71

85

64

81

81

90

91

74

%

23

,5

20,3

17

,9

13,0

25

,2

22,8

5,

1 29

,2

21,7

23

,7

28,1

21

,3

26,6

27

,0

30,0

30

,3

24,7

Todi

spor

ites

0 0

1 0

0 0

0 0

0 1

0 0

0 0

0 0

0%

0,0

0,0

0,3

0,0

0,0

0,0

0,0

0,0

0,0

0,3

0,0

0,0

0,0

0,0

0,0

0,0

0,0

Verr

ucos

ispo

rites

0

0 0

0 0

0 0

1 0

0 0

1 0

0 0

0 0

%

0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

0 0,

0G

emm

atril

etes

0

0 0

1 0

0 0

0 0

0 0

0 0

0 0

0 0

%

0,

0 0,

0 0,

0 0,

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,

0B

iretis

porit

es

0 2

2 0

3 0

1 0

1 4

0 1

0 5

0 3

0%

0,0

0,7

0,7

0,0

1,0

0,0

0,3

0,0

0,3

1,3

0,0

0,3

0,0

1,7

0,0

1,0

0,0

Neo

rais

trick

ia

0 1

2 0

5 0

1 0

0 3

0 0

0 5

0 0

2%

0,0

0,3

0,7

0,0

1,7

0,0

0,3

0,0

0,0

1,0

0,0

0,0

0,0

1,7

0,0

0,0

0,7

Ret

rilet

es

0

1 0

1 0

0 0

1 0

0 0

0 0

0 0

1 1

%

0,

0 0,

3 0,

0 0,

3 0,

0 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

3 0,

3A

equi

trira

dite

s 0

0 1

0 1

0 0

1 0

0 0

0 0

0 0

0 0

%

0,

0 0,

0 0,

3 0,

0 0,

3 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0Le

ptol

epid

ites

2 0

0 1

3 0

1 1

1 0

0 3

2 0

3 0

1%

0,7

0,0

0,0

0,3

1,0

0,0

0,3

0,3

0,3

0,0

0,0

1,0

0,7

0,0

1,0

0,0

0,3

Cam

aroz

onos

porit

es

0 0

0 0

0 0

0 0

0 0

0 1

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

0,3

0,0

0,0

0,0

0,0

0,0

Den

soip

orite

s 0

0 0

0 1

0 0

0 0

0 0

0 0

0 0

0 0

%

0,

0 0,

0 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0Fo

veos

porit

es

0 0

2 0

0 0

0 3

1 1

0 0

0 0

0 1

0%

0,0

0,0

0,7

0,0

0,0

0,0

0,0

1,0

0,3

0,3

0,0

0,0

0,0

0,0

0,0

0,3

0,0

Lyco

podi

acid

ites

0 0

0 1

0 0

2 0

0 0

0 1

0 0

0 1

0%

0,0

0,0

0,0

0,3

0,0

0,0

0,7

0,0

0,0

0,0

0,0

0,3

0,0

0,0

0,0

0,3

0,0

Ara

trisp

orite

s 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 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0Ly

copo

dium

spor

ites

0 0

0 0

0 0

0 0

10

0 1

0 0

0 13

0

0%

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

3,3

0,0

0,3

0,0

0,0

0,0

4,3

0,0

0,0

Echi

natis

poris

0

2 3

3 2

0 0

3 0

5 0

3 1

2 3

0 0

%

0,

0 0,

7 1,

0 1,

0 0,

7 0,

0 0,

0 1,

0 0,

0 1,

7 0,

0 1,

0 0,

3 0,

7 1,

0 0,

0 0,

0Fo

ram

inis

poris

2

0 0

0 0

0 0

0 0

0 1

0 0

2 0

0 0

%

0,

7 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

3 0,

0 0,

0 0,

7 0,

0 0,

0 0,

0Sp

ore

mon

olet

e 0

0 0

0 0

4 4

0 6

0 0

0 5

0 3

0 3

%

0,

0 0,

0 0,

0 0,

0 0,

0 1,

3 1,

3 0,

0 2,

0 0,

0 0,

0 0,

0 1,

6 0,

0 1,

0 0,

0 1,

0St

aplin

ispo

rites

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 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0M

icro

retic

ulat

ispo

ris

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

0,0

0,0

0,0

0,0

0,0

0,0

0,0

Ster

eisp

orite

s 0

0 0

3 3

0 0

1 0

0 0

0 0

1 3

0 1

%

0,

0 0,

0 0,

0 1,

0 1,

0 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

0 0,

0 0,

3 1,

0 0,

0 0,

3O

void

ites

1

13

2 4

2 13

25

0

9 2

7 2

10

0 4

1 10

%

0,

3 4,

2 0,

7 1,

3 0,

7 4,

4 8,

4 0,

0 3,

0 0,

7 2,

3 0,

7 3,

3 0,

0 1,

3 0,

3 3,

3Le

iosp

haer

idia

1

3 0

4 1

2 0

3 1

0 1

1 0

1 0

0 2

%

0,

3 1,

0 0,

0 1,

3 0,

3 0,

7 0,

0 1,

0 0,

3 0,

0 0,

3 0,

3 0,

0 0,

3 0,

0 0,

0 0,

7

163

Appendix

Cou

nts s

pore

-pol

len

Mid

-Pol

ish

Trou

gh, P

olan

d:

dept

h:

1290

13

30

1380

14

30

1480

15

20

1590

16

30

1670

17

10

1750

17

90

taxa

: sa

mpl

e:

PIG

1_13

6 PI

G1_

140

PIG

1_14

5 PI

G1_

150

PIG

1_15

5 PI

G1_

159

PIG

1_16

6 PI

G1_

170

PIG

1_17

4 PI

G1_

178

PIG

1_18

2 PI

G1_

186

MIN

M

AX

AV

RPa

rvis

acci

tes

0 0

0 0

2 0

0 0

0 0

0 1

0 6

1%

0,0

0,0

0,0

0,0

0,7

0,0

0,0

0,0

0,0

0,0

0,0

0,3

0,0

2,0

0,2

Podo

carp

idite

s 0

0 0

2 0

3 0

0 2

0 0

0 0

3 0

%

0,

0 0,

0 0,

0 0,

7 0,

0 1,

0 0,

0 0,

0 0,

7 0,

0 0,

0 0,

0 0,

0 1,

0 0,

1Pi

nusp

olle

nite

s 0

0 0

6 5

3 2

1 2

0 0

1 0

6 2

%

0,

0 0,

0 0,

0 2,

0 1,

7 1,

0 0,

7 0,

3 0,

7 0,

0 0,

0 0,

3 0,

0 2,

0 0,

5bi

sacc

ate

indi

ff.

21

29

20

73

42

26

31

61

37

31

12

18

6 79

35

%

7,

0 9,

5 6,

7 24

,2

14,0

8,

7 10

,3

20,2

12

,3

10,3

4,

0 6,

0 2,

0 26

,3

11,7

Ara

ucar

iaci

tes

15

8 20

20

21

12

23

42

34

13

31

5

0 42

15

%

5,

0 2,

6 6,

7 6,

6 7,

0 4,

0 7,

7 13

,9

11,3

4,

3 10

,4

1,7

0,0

13,9

5,

0C

allia

llasp

orite

s 19

34

22

36

18

26

42

13

20

22

37

25

0

48

22%

6,3

11,1

7,

4 11

,9

6,0

8,7

14,0

4,

3 6,

6 7,

3 12

,5

8,3

0,0

16,4

7,

2In

aper

turo

polle

nite

s 12

12

21

18

16

8

19

24

12

11

16

16

1 39

17

%

4,

0 3,

9 7,

0 6,

0 5,

4 2,

7 6,

3 7,

9 4,

0 3,

7 5,

4 5,

3 0,

3 13

,0

5,6

Perin

opol

leni

tes

11

13

19

8 36

9

14

10

9 17

24

56

0

56

19%

3,7

4,3

6,4

2,6

12,0

3,

0 4,

7 3,

3 3,

0 5,

6 8,

1 18

,7

0,0

18,7

6,

3C

ereb

ropo

lleni

tes

4 4

2 5

4 1

4 6

19

6 3

9 1

29

11%

1,3

1,3

0,7

1,7

1,3

0,3

1,3

2,0

6,3

2,0

1,0

3,0

0,3

9,6

3,6

Cla

ssop

ollis

1

2 6

17

7 9

2 8

20

15

7 15

1

220

19%

0,3

0,7

2,0

5,6

2,3

3,0

0,7

2,6

6,6

5,0

2,4

5,0

0,3

73,1

6,

3Ex

esip

olle

nite

s 1

0 0

1 0

1 0

0 0

1 0

0 0

5 1

%

0,

3 0,

0 0,

0 0,

3 0,

0 0,

3 0,

0 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 1,

6 0,

2C

ycad

opite

s 6

12

6 13

5

8 7

2 7

3 12

4

0 21

7

%

2,

0 3,

9 2,

0 4,

3 1,

7 2,

7 2,

3 0,

7 2,

3 1,

0 4,

0 1,

3 0,

0 7,

0 2,

2Eu

com

miid

ites

3 0

0 3

1 0

1 0

0 0

2 3

0 17

1

%

1,

0 0,

0 0,

0 1,

0 0,

3 0,

0 0,

3 0,

0 0,

0 0,

0 0,

7 1,

0 0,

0 5,

7 0,

5A

lispo

rites

4 1

0 6

3 4

4 6

1 2

0 1

0 17

2

%

1,

3 0,

3 0,

0 2,

0 1,

0 1,

3 1,

3 2,

0 0,

3 0,

7 0,

0 0,

3 0,

0 5,

7 0,

8V

itrei

spor

ites p

allid

us

19

1 15

7

10

3 3

5 7

13

0 9

0 19

5

%

6,

3 0,

3 5,

0 2,

3 3,

3 1,

0 1,

0 1,

7 2,

3 4,

3 0,

0 3,

0 0,

0 6,

3 1,

5Ep

hedr

epite

s 0

1 0

0 0

1 0

0 0

0 1

1 0

10

1%

0,0

0,3

0,0

0,0

0,0

0,3

0,0

0,0

0,0

0,0

0,3

0,3

0,0

3,4

0,4

Cya

thid

ites

32

32

41

15

15

36

25

21

12

24

19

22

1

65

27%

10,7

10

,5

13,7

5,

0 5,

0 12

,0

8,3

7,0

4,0

8,0

6,4

7,3

0,3

21,6

8,

9Le

iotri

lete

s

8 25

11

5

15

24

7 8

16

9 7

7 0

42

14%

2,7

8,2

3,7

1,7

5,0

8,0

2,3

2,6

5,3

3,0

2,4

2,3

0,0

14,0

4,

8C

onca

visp

orite

s 4

6 1

1 2

1 5

0 3

5 3

1 0

6 2

%

1,

3 2,

0 0,

3 0,

3 0,

7 0,

3 1,

7 0,

0 1,

0 1,

7 1,

0 0,

3 0,

0 2,

0 0,

6C

onca

viss

imis

porit

es

0 0

0 0

0 0

2 0

0 0

0 0

0 3

0%

0,0

0,0

0,0

0,0

0,0

0,0

0,7

0,0

0,0

0,0

0,0

0,0

0,0

1,0

0,1

Trilo

bosp

orite

s 0

0 2

0 0

0 0

1 2

1 0

0 0

8 1

%

0,

0 0,

0 0,

7 0,

0 0,

0 0,

0 0,

0 0,

3 0,

7 0,

3 0,

0 0,

0 0,

0 2,

7 0,

3C

icat

ricos

ispo

rites

4

3 0

1 1

9 5

1 6

1 5

3 0

9 4

%

1,

3 1,

0 0,

0 0,

3 0,

3 3,

0 1,

7 0,

3 2,

0 0,

3 1,

7 1,

0 0,

0 3,

0 1,

3K

luki

spor

ites

1 1

1 0

0 0

0 0

0 0

0 0

0 1

0%

0,3

0,3

0,3

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,3

0,0

Isch

yosp

orite

s 2

0 0

0 0

0 1

0 0

0 0

2 0

5 1

%

0,

7 0,

0 0,

0 0,

0 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

0 0,

7 0,

0 1,

7 0,

2O

smun

daci

dite

s 1

1 5

1 3

0 5

3 4

1 1

4 0

10

3%

0,3

0,3

1,7

0,3

1,0

0,0

1,7

1,0

1,3

0,3

0,3

1,3

0,0

3,2

1,0

164

Cou

nts s

pore

-pol

len

Mid

-Pol

ish

Trou

gh, P

olan

d:

dept

h:

1290

13

30

1380

14

30

1480

15

20

1590

16

30

1670

17

10

1750

17

90

taxa

: sa

mpl

e:

PIG

1_13

6 PI

G1_

140

PIG

1_14

5 PI

G1_

150

PIG

1_15

5 PI

G1_

159

PIG

1_16

6 PI

G1_

170

PIG

1_17

4 PI

G1_

178

PIG

1_18

2 PI

G1_

186

MIN

M

AX

AV

RG

leic

heni

idite

s 8

18

22

19

9 26

15

13

13

27

19

25

0

30

16%

2,7

5,9

7,4

6,3

3,0

8,7

5,0

4,3

4,3

9,0

6,4

8,3

0,0

10,0

5,

2C

lava

tispo

rites

0

0 0

0 2

0 0

0 0

0 0

1 0

3 0

%

0,

0 0,

0 0,

0 0,

0 0,

7 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

3 0,

0 1,

0 0,

1de

ltoid

e Sp

ore

indi

ff.

115

96

76

20

72

80

50

75

66

87

78

62

0 11

5 59

%

38

,3

31,5

25

,4

6,6

24,1

26

,7

16,7

24

,8

21,9

28

,9

26,3

20

,7

0,0

38,3

19

,7To

disp

orite

s 0

0 0

0 0

0 0

1 0

0 0

0 0

1 0

%

0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

3 0,

0 0,

0 0,

0 0,

0 0,

0 0,

3 0,

0Ve

rruc

osis

porit

es

0 0

0 0

0 0

0 0

0 0

0 0

0 1

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

0,3

0,0

Gem

mat

rilet

es

0 0

0 0

0 0

0 0

0 0

0 0

0 1

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

0,3

0,0

Bire

tispo

rites

3

0 2

1 2

0 0

0 0

3 0

2 0

7 1

%

1,

0 0,

0 0,

7 0,

3 0,

7 0,

0 0,

0 0,

0 0,

0 1,

0 0,

0 0,

7 0,

0 2,

3 0,

4N

eora

istri

ckia

1

1 0

3 0

0 5

0 0

0 0

0 0

5 1

%

0,

3 0,

3 0,

0 1,

0 0,

0 0,

0 1,

7 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 1,

7 0,

4R

etril

etes

0 0

0 0

0 0

1 0

0 0

0 0

0 1

0%

0,0

0,0

0,0

0,0

0,0

0,0

0,3

0,0

0,0

0,0

0,0

0,0

0,0

0,3

0,1

Aeq

uitri

radi

tes

1 0

0 0

2 0

0 0

0 2

0 1

0 2

0%

0,3

0,0

0,0

0,0

0,7

0,0

0,0

0,0

0,0

0,7

0,0

0,3

0,0

0,7

0,1

Lept

olep

idite

s 0

0 0

2 1

0 2

0 1

0 1

1 0

4 1

%

0,

0 0,

0 0,

0 0,

7 0,

3 0,

0 0,

7 0,

0 0,

3 0,

0 0,

3 0,

3 0,

0 1,

3 0,

4C

amar

ozon

ospo

rites

0

0 0

0 0

0 0

0 0

0 0

0 0

1 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,

3 0,

0D

enso

ipor

ites

0 0

0 0

0 1

0 0

0 1

0 0

0 1

0%

0,0

0,0

0,0

0,0

0,0

0,3

0,0

0,0

0,0

0,3

0,0

0,0

0,0

0,3

0,0

Fove

ospo

rites

0

0 0

0 0

0 0

0 0

0 0

1 0

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,

3 0,

0 1,

0 0,

1Ly

copo

diac

idite

s 0

0 0

0 0

0 0

0 0

0 0

0 0

2 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,

7 0,

1A

ratri

spor

ites

0 0

0 0

0 0

0 0

0 0

0 0

0 21

1

%

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 7,

0 0,

2Ly

copo

dium

spor

ites

0 0

3 0

0 0

0 0

0 2

2 1

0 13

1

%

0,

0 0,

0 1,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

7 0,

7 0,

3 0,

0 4,

3 0,

2Ec

hina

tispo

ris

3 0

1 0

2 0

0 1

0 3

1 1

0 7

1%

1,0

0,0

0,3

0,0

0,7

0,0

0,0

0,3

0,0

1,0

0,3

0,3

0,0

2,3

0,5

Fora

min

ispo

ris

0 0

0 0

0 0

0 0

0 0

0 0

0 6

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

2,0

0,1

Spor

e m

onol

ete

0 0

0 1

0 0

0 0

2 0

1 0

0 6

1%

0,0

0,0

0,0

0,3

0,0

0,0

0,0

0,0

0,7

0,0

0,3

0,0

0,0

2,0

0,3

Stap

linis

porit

es

0 0

0 0

0 1

0 0

0 0

0 0

0 3

0%

0,0

0,0

0,0

0,0

0,0

0,3

0,0

0,0

0,0

0,0

0,0

0,0

0,0

1,0

0,0

Mic

rore

ticul

atis

poris

0

0 0

0 0

0 0

0 0

0 0

0 0

2 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,

7 0,

0St

erei

spor

ites

1 0

0 3

1 0

3 0

0 0

1 0

0 3

1%

0,3

0,0

0,0

1,0

0,3

0,0

1,0

0,0

0,0

0,0

0,3

0,0

0,0

1,0

0,2

Ovo

idite

s

0 5

2 15

1

5 20

0

5 1

11

2 0

31

6%

0,0

1,6

0,7

5,0

0,3

1,7

6,7

0,0

1,7

0,3

3,7

0,7

0,0

10,3

2,

1Le

iosp

haer

idia

0

0 1

0 1

3 2

0 2

0 3

0 0

4 1

%

0,

0 0,

0 0,

3 0,

0 0,

3 1,

0 0,

7 0,

0 0,

7 0,

0 1,

0 0,

0 0,

0 1,

3 0,

3

165

Appendix

Cou

nts s

pore

-pol

len,

Voc

ontia

n B

asin

, Fra

nce.

de

pth:

4,

1 6,

3 9,

3 12

,1

16,4

18

,2

20,3

23

,1

25,2

26

,8

32,3

34

,0

36,4

38

,8

40,6

45

,6

51,9

taxa

: sa

mpl

e:

VER

2 V

ER6

VER

11

VER

16

VER

22

VER

28

VER

32

VER

36

VER

38

VER

40

VER

41

VER

44

VER

50

VER

53

VER

56

VER

64

VER

72Pa

rvis

acci

tes

7,0

0,0

1,0

0,0

3,0

0,0

2,0

0,0

3,0

0,0

0,0

4,0

1,0

1,0

0,0

2,0

2,0

%

2,

3 0,

0 0,

3 0,

0 1,

0 0,

0 0,

7 0,

0 1,

0 0,

0 0,

0 1,

3 0,

3 0,

3 0,

0 0,

7 0,

7Po

doca

rpid

ites

1,0

0,0

0,0

1,0

0,0

3,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,0

1,0

%

0,

3 0,

0 0,

0 0,

3 0,

0 1,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

3Pi

nusp

olle

nite

s 0,

0 0,

0 0,

0 0,

0 0,

0 2,

0 4,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 0,

0 1,

0 0,

0 1,

0%

0,0

0,0

0,0

0,0

0,0

0,7

1,3

0,0

0,0

0,0

0,0

0,0

0,0

0,0

0,3

0,0

0,3

bisa

ccat

e in

diff.

27

,0

24,0

16

,0

29,0

9,

0 26

,0

35,0

20

,0

18,0

15

,0

11,0

25

,0

12,0

4,

0 37

,0

29,0

22

,0%

9,0

8,0

5,3

9,7

3,0

8,7

11,7

6,

7 6,

0 5,

0 3,

6 8,

3 4,

0 1,

3 12

,3

9,7

7,3

Ara

ucar

iaci

tes

51,0

20

,0

37,0

16

,0

30,0

16

,0

37,0

57

,0

38,0

40

,0

43,0

43

,0

44,0

4,

0 22

,0

54,0

13

,0%

17,0

6,

7 12

,3

5,3

10,0

5,

3 12

,3

19,0

12

,7

13,3

14

,2

14,3

14

,7

1,3

7,3

18,0

4,

3C

allia

llasp

orite

s 18

,0

28,0

20

,0

10,0

17

,0

25,0

31

,0

15,0

16

,0

18,0

11

,0

12,0

18

,0

3,0

40,0

45

,0

37,0

%

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166

167

Appendix

Cou

nts s

pore

-pol

len,

Voc

ontia

n B

asin

, Fra

nce.

de

pth:

4,

1 6,

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: sa

mpl

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2 V

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11

VER

16

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22

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28

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32

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36

VER

38

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40

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41

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44

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50

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53

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56

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64

VER

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leic

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vatis

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mat

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168

169

Appendix

Cou

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Fora

min

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len,

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min

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void

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2,

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%

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7

170

171

Appendix

Paly

nofa

cies

cou

nts,

Mid

-Pol

ish

troug

h, P

olan

dde

pth

(m)

sam

ple

coun

ts

AO

M

Acr

itarc

hs

Din

ocys

ts

For

am li

n. n

on-c

ut. m

emb.

Cut

icle

s P

hyto

cl. o

p. P

hyto

cl. t

rans

l. po

llen

Spo

res

char

coal

no

t ide

ntifi

able

0,3

PIG

1.5

41

7,0

17,0

15

,0

23,0

3,

0 15

,0

0,0

122,

0 58

,0

19,0

25

,0

3,0

117,

00,

5 PI

G 1

.8

405,

0 26

,0

1,0

2,0

0,0

8,0

0,0

125,

0 15

8,0

36,0

10

,0

25,0

14

,00,

8 PI

G 1

.11

411,

0 2,

0 5,

0 29

,0

4,0

19,0

0,

0 8,

0 28

6,0

36,0

13

,0

0,0

10,0

1,0

PIG

1.1

3 40

5,0

5,0

5,0

46,0

1,

0 13

,0

0,0

102,

0 77

,0

16,0

19

,0

2,0

119,

01,

4 PI

G1.

17

401,

0 15

,0

9,0

52,0

2,

0 17

,0

1,0

22,0

22

5,0

27,0

19

,0

1,0

11,0

1,8

PIG

1.2

1 40

2,0

8,0

7,0

23,0

2,

0 9,

0 0,

0 17

7,0

65,0

15

,0

12,0

20

,0

65,0

2,2

PIG

1.2

5 40

4,0

4,0

1,0

19,0

0,

0 11

,0

0,0

6,0

341,

0 7,

0 8,

0 1,

0 6,

02,

5 PI

G 1

.28

407,

0 21

,0

11,0

19

,0

1,0

12,0

0,

0 96

,0

69,0

27

,0

30,0

3,

0 11

8,0

2,8

PIG

1.3

1 41

2,0

18,0

2,

0 19

,0

1,0

7,0

0,0

18,0

32

9,0

8,0

4,0

0,0

6,0

3,2

PIG

1.3

5 41

8,0

10,0

7,

0 21

,0

1,0

11,0

0,

0 15

0,0

83,0

24

,0

23,0

0,

0 88

,03,

6 PI

G 1

.39

420,

0 9,

0 4,

0 24

,0

2,0

12,0

0,

0 7,

0 32

8,0

14,0

6,

0 2,

0 12

,04,

2 PI

G 1

.45

407,

0 22

,0

3,0

16,0

2,

0 7,

0 0,

0 11

8,0

97,0

31

,0

23,0

0,

0 88

,04,

6 PI

G 1

.49

406,

0 18

,0

5,0

8,0

1,0

9,0

3,0

113,

0 15

3,0

44,0

32

,0

2,0

18,0

5,0

PIG

1.5

3 42

3,0

21,0

9,

0 32

,0

5,0

23,0

1,

0 11

8,0

132,

0 39

,0

30,0

0,

0 13

,05,

3 PI

G 1

.56

436,

0 2,

0 0,

0 1,

0 0,

0 2,

0 0,

0 12

6,0

122,

0 10

3,0

10,0

28

,0

42,0

5,6

PIG

1.5

9 42

1,0

19,0

30

,0

28,0

1,

0 26

,0

0,0

143,

0 89

,0

32,0

34

,0

0,0

19,0

6,0

PIG

1.6

3 44

7,0

15,0

3,

0 6,

0 2,

0 6,

0 0,

0 14

,0

370,

0 3,

0 5,

0 0,

0 23

,06,

4 PI

G 1

.67

398,

0 12

,0

7,0

17,0

0,

0 26

,0

0,0

116,

0 10

9,0

49,0

35

,0

2,0

25,0

6,8

PIG

1.7

1 40

0,0

10,0

14

,0

34,0

0,

0 14

,0

0,0

128,

0 69

,0

20,0

24

,0

1,0

86,0

7,1

PIG

1.7

4 42

0,0

19,0

16

,0

51,0

0,

0 36

,0

0,0

111,

0 11

3,0

36,0

26

,0

4,0

8,0

7,5

PIG

1.7

8 41

8,0

9,0

0,0

7,0

0,0

3,0

0,0

6,0

384,

0 1,

0 2,

0 0,

0 7,

07,

9 PI

G1.

82

413,

0 29

,0

21,0

47

,0

1,0

28,0

0,

0 96

,0

100,

0 33

,0

34,0

4,

0 20

,08,

3 PI

G 1

.86

409,

0 27

,0

10,0

20

,0

3,0

20,0

0,

0 11

9,0

68,0

27

,0

25,0

5,

0 85

,08,

8 PI

G 1

.91

411,

0 8,

0 7,

0 5,

0 0,

0 2,

0 0,

0 23

,0

355,

0 4,

0 0,

0 5,

0 2,

09,

1 PI

G 1

.94

412,

0 27

,0

20,0

46

,0

6,0

16,0

0,

0 92

,0

102,

0 30

,0

26,0

5,

0 43

,09,

5 PI

G 1

.98

423,

0 16

,0

10,0

12

9,0

6,0

17,0

0,

0 75

,0

54,0

23

,0

16,0

10

,0

67,0

9,9

PIG

1.1

02

408,

0 20

,0

15,0

24

,0

0,0

37,0

0,

0 11

0,0

117,

0 27

,0

20,0

1,

0 38

,010

,3

PIG

1.1

06

406,

0 11

,0

23,0

62

,0

1,0

8,0

1,0

78,0

54

,0

15,0

17

,0

0,0

137,

010

,7

PIG

1.1

10

417,

0 45

,0

6,0

38,0

4,

0 12

,0

0,0

72,0

12

6,0

49,0

32

,0

1,0

32,0

11,1

PI

G 1

.114

41

8,0

19,0

15

,0

80,0

3,

0 15

,0

1,0

97,0

52

,0

17,0

19

,0

5,0

95,0

11,5

PI

G 1

.118

40

6,0

25,0

13

,0

32,0

0,

0 29

,0

3,0

74,0

14

9,0

39,0

24

,0

1,0

17,0

11,9

PI

G 1

.122

43

6,0

7,0

13,0

12

7,0

8,0

5,0

0,0

76,0

54

,0

12,0

32

,0

5,0

97,0

12,2

PI

G 1

.127

40

9,0

15,0

13

,0

66,0

6,

0 22

,0

1,0

122,

0 78

,0

25,0

33

,0

2,0

26,0

12,5

PI

G 1

.131

40

7,0

4,0

25,0

83

,0

5,0

7,0

1,0

107,

0 51

,0

18,0

21

,0

4,0

81,0

12,9

PI

G 1

.136

40

8,0

16,0

21

,0

89,0

7,

0 25

,0

0,0

80,0

83

,0

19,0

37

,0

0,0

31,0

13,4

PI

G 1

.140

46

1,0

3,0

28,0

90

,0

9,0

16,0

1,

0 15

9,0

55,0

20

,0

17,0

5,

0 58

,013

,8

PIG

1.1

45

405,

0 27

,0

17,0

11

8,0

11,0

27

,0

0,0

45,0

77

,0

25,0

20

,0

3,0

35,0

14,3

PI

G 1

.150

43

3,0

10,0

6,

0 43

,0

7,0

9,0

0,0

63,0

17

1,0

23,0

6,

0 7,

0 88

,014

,8

PIG

1.1

55

413,

0 21

,0

18,0

69

,0

11,0

28

,0

0,0

91,0

95

,0

22,0

19

,0

2,0

37,0

15,2

PI

G 1

.159

42

7,0

2,0

29,0

77

,0

1,0

15,0

1,

0 95

,0

53,0

16

,0

22,0

6,

0 11

0,0

15,9

PI

G 1

.166

42

8,0

1,0

9,0

33,0

1,

0 30

,0

1,0

86,0

14

3,0

7,0

10,0

16

,0

91,0

16,3

PI

G 1

.170

40

8,0

20,0

18

,0

89,0

6,

0 29

,0

0,0

95,0

77

,0

33,0

19

,0

0,0

22,0

16,7

PI

G 1

.174

43

2,0

18,0

13

,0

49,0

0,

0 41

,0

1,0

120,

0 67

,0

19,0

21

,0

4,0

79,0

17,1

PI

G 1

.178

42

5,0

17,0

16

,0

117,

0 7,

0 25

,0

0,0

97,0

74

,0

24,0

28

,0

2,0

18,0

17,5

PI

G 1

.182

41

0,0

24,0

12

,0

79,0

0,

0 15

,0

0,0

121,

0 43

,0

8,0

18,0

10

,0

80,0

17,9

PI

G 1

.186

41

2,0

53,0

17

,0

62,0

9,

0 21

,0

0,0

67,0

89

,0

30,0

31

,0

0,0

33,0

Paly

nofa

cies

cou

nts V

ocon

tian

Bas

in, F

ranc

e:de

pth

(m)

sam

ple

coun

ts

A

OM

Acr

itarc

hs

Din

ocys

ts

Fora

m li

n. n

on-c

ut. m

emb.

Cut

icle

s P

hyto

cl. o

p.

Phyt

ocl.

trans

l. p

olle

n

sp

ores

ch

arco

al n

ot id

entifi

able

4 V

ER2

412

182

17

34

13

30

0 8

24

47

22

0 35

6 V

ER6

415

109

19

38

3 61

0

53

21

7 10

1

939

VER

11

402

126

20

54

8 16

0

31

35

35

22

0 55

12

VER

16

402

90

19

36

7 70

0

36

28

12

11

2 91

16

VER

22

403

102

37

48

14

38

0 20

21

42

14

1

6618

V

ER24

42

3 38

9 0

1 1

3 0

3 16

1

4 0

518

V

ER26

40

7 39

9 1

0 0

0 0

0 5

1 0

0 1

18

VER

28

421

109

16

39

13

47

0 57

29

16

11

2

8219

V

ER29

39

9 38

5 0

0 0

0 0

1 7

0 0

0 6

20

VER

32

419

89

40

64

17

46

0 43

26

9

10

3 72

23

VER

36

407

93

38

46

12

32

0 35

35

55

9

0 52

25

VER

38

406

59

39

48

19

25

0 41

44

63

25

1

4227

V

ER40

40

6 81

49

51

12

26

0

36

38

44

13

1 55

30

MO

R5

412

371

1 6

0 0

0 2

2 0

1 0

2932

V

ER41

42

1 92

61

31

22

25

0

43

40

62

10

0 35

34

VER

44

403

131

53

36

16

63

0 31

20

8

4 0

4136

V

ER50

41

9 17

1 23

20

12

26

0

19

30

64

17

0 37

39

VER

53

411

122

30

31

14

19

0 20

30

75

36

1

3341

V

ER56

40

7 21

9 20

19

8

29

0 13

24

13

9

0 53

46

VER

64

408

145

39

23

21

29

0 19

33

50

20

0

2952

V

ER72

41

4 97

37

43

4

40

1 51

40

29

11

0

6157

V

ER78

40

2 14

3 20

26

8

24

0 44

35

42

36

2

2262

V

ER86

41

1 14

2 13

44

1

32

0 48

34

18

20

0

5965

LC

4 41

5 14

9 47

18

4

21

0 36

47

47

19

0

2769

LC

10

411

262

10

11

1 26

0

33

9 19

9

0 31

75

LC26

41

5 19

4 8

28

6 27

0

39

36

31

10

0 36

82

LC38

41

3 16

8 8

43

7 24

0

20

46

21

9 1

6688

LC

50

429

110

4 41

6

23

0 80

75

9

9 5

6795

LC

66

404

42

14

48

9 23

0

92

66

31

16

3 60

103

LC81

40

6 11

0 11

34

8

21

0 46

56

27

24

1

6810

7 LC

87

425

230

18

32

4 22

0

52

35

11

3 0

1811

1 LC

98

423

61

22

35

11

31

0 95

83

13

17

2

5311

5 LC

113

417

40

14

57

7 30

0

97

84

22

6 1

5912

1 LC

132

411

21

16

69

2 8

0 97

98

6

14

0 80

131

LC14

8 40

8 66

28

47

5

24

2 68

83

20

12

2

5114

1 LC

158

410

56

26

51

3 33

2

54

95

25

15

0 50

147

LC16

6 40

9 29

10

10

1 2

18

21

94

75

14

12

1 32

152

LC17

3 41

1 12

9 12

54

4

28

0 63

42

32

20

0

2716

0 LC

179

407

69

7 11

1 5

22

0 52

66

6

3 0

6616

9 LC

189

414

241

7 50

1

13

0 27

41

13

7

0 14

178

LC19

8 41

4 99

3

59

4 11

0

105

78

6 5

1 43

172

173

Appendix

174

175

Curriculum vitae

Curriculum vitae

Ariane Kujau * 2nd April1983 in Nordhorn, Germany unmarried

03/2009 - 07/2012 PhD student and research assistant at the Ruhr-University Bochum, Germany, Faculty for Geosciences, Institute for Geology, Mineralogy and Geophysics

Dissertation: “Climatic and environmental dynamics during the Valanginian carbon isotope event - Evidence from geochemistry and palynology“.

Supervisors: Prof. Dr. Ulrich Heimhofer (University of Hannover, Germany), Prof. Dr. Jörg Mutterlose (Ruhr-University Bochum, Germany).

11/2008 - 02/2009 Participation in a playground renovation project (evaluation), City of Osnabrück, Department of Environment.

10/2002 - 10/2008 Study of Physical Geography (minor subjects: Biology, Social Sciences) at the University of Osnabrück, Germany. Diploma thesis at the GEOMAR - Helmholtz-Centre for Ocean Research, Kiel, Germany: “The terrigenous Mississippi discharge into the northeaastern Gulf of Mexico over the last 560,000 years - Indications from x-ray fluorescence core-scanning data“. Supervisors: Prof. Dr. Christoph Zielhofer (University of Leipzig, Germany), Prof. Dr. Dirk Nürnberg (GEOMAR - Helmholtz- Centre for Ocean Research, Kiel, Germany).

09/2006 - 03/2007 Visiting student at the Università degli Studi di Roma, Rome, Italy.

1995 - 2002 Grammar School, Windthorst Gymnasium Meppen, Germany.1993 - 1995 5th and 6th grade Orientierungsstufe, Johannesschule Meppen, Germany.1989 - 1993 Primary School, Johannes Gutenberg Grundschule Meppen, Germany.