Research Collection

Doctoral Thesis

Response of terrestrial paleoenvironments to past changes in climate and carbon-cycling: Insights from palynology and stable isotope geochemistry

Author(s): Heimhofer, Ulrich

Publication Date: 2004

Permanent Link: https://doi.org/10.3929/ethz-a-004741183

Rights / License: In Copyright - Non-Commercial Use Permitted

This page was generated automatically upon download from the ETH Zurich Research Collection. For more information please consult the Terms of use.

ETH Library DISS ETH No. 15463

Response of terrestrial palaeoenvironments to past changes in climate and carbon-cycling: Insights from palynology and stable isotope geochemistry

A dissertation submitted to the SWISS FEDERAL INSTITUTE OF TECHNOLOGY ZURICH for the degree of DOCTOR OF SCIENCES

Presented by

Ulrich Heimhofer Dipl. Geol. Univ. Erlangen-Nürnberg born October 19, 1971 Sonthofen i. Allgäu / Germany

Accepted on the recommendation of

Prof. Dr. Helmut Weissert, ETH Zurich, examiner Dr. Peter A. Hochuli, University of Zurich, co-examiner Prof. Dr. Judith A. McKenzie, ETH Zurich, co-examiner Dr. Stephen P. Hesselbo, University of Oxford, co-examiner

2004

Table of contents

Table of contents

Abstract ……………………………………………………………………………………..3 Zusammenfassung ……………………………………………………………………………..5

Chapter 1 Introduction …...………………………………………………………………………………...7

Chapter 2

Absence of major vegetation and palaeoatmospheric pCO2 changes associated with Oceanic Anoxic Event 1a (Early Aptian, SE France) ……………….…...……………………………..17

Chapter 3 Palynological and calcareous nannofossil records across the late Early Aptian OAE 1a: Implications for palaeoclimate, palaeofertility and detrital input .………………….……...43

Chapter 4 Terrestrial carbon-isotope records from coastal deposits (Algarve, Portugal): A tool for chemostratigraphic correlation on an intrabasinal and global scale …………………73

Chapter 5 A well-dated and continuous early angiosperm pollen record from mid-Cretaceous coastal deposits (Lusitanian and Algarve Basins, Portugal): Implications for the timing of the early angiosperm radiation ………………………………….85

Chapter 6 Conclusions …………………………………………………………………………………..147

Appendix A 1 to A 7 ………………………………………………………..…………………….…...149

Acknowledgements ………………..…………………………………………………………165

Curriculum vitae……………………………………………...………………………..…..…...167

Abstract 3

Abstract

The mid-Cretaceous (Aptian to Turonian, 120-90 Ma) was characterised by globally averaged surface temperatures of up to 10ºC and is considered as one of the best examples of greenhouse-type climate conditions in the Phanerozoic Earth history. Evidence for this exceptional climate mode includes low latitudinal thermal gradients, increased surface and bathyal ocean water temperatures, the occurrence of thermophilic plant assemblages in high latitude regions and the absence of expanded polar ice-sheets. However, during this period of global warmth, climatic conditions were far from stable. Short-term perturbations of the global carbon-cycle and climates are reflected in the deposition of organic carbon-rich black shales, shifts in the carbon isotope record and dramatic growth-crisis of biocalcifying organisms. In order to investigate terrestrial environments and their response to mid- Cretaceous global change, Late Barremian to Albian deposits are studied with a combined approach, including palynology, carbon isotopes and organic geochemistry. The late Early Aptian oceanic anoxic event (OAE) 1a interval in the Vocontian Basin, SE France has been chosen to serve as a high-resolution environmental archive, covering a time of short-term palaeo-climatic and oceanographic change. Based on the δ13C composition of marine carbonates and individual biomarkers, palaeoatmospheric CO2 partial pressure during and after black shale formation has been estimated. To address possible vegetation changes in the hinterland of the Vocontian Basin, the occurring spore-pollen assemblages were determined. Furthermore, dinoflagellate cyst and calcareous nannofossil assemblages were analysed and Corg accumulation rates were estimated to identify changes in palaeoceanographic conditions. Our results indicate that intensified Corg burial in black shales during the late Early Aptian was accompanied by an only moderate drop in CO2 partial pressure. The pollen spectrum indicates relatively stable vegetation patterns during and after times of OAE 1a formation. Likewise, the organic-walled and calcareous plankton display no significant changes in the prevailing palaeoceanographic conditions across the black shale interval. In contrast to previous studies, our results exhibit no indication of enhanced humidity and nutrient-input, which probably triggered oceanic surface water productivity and resulted in the deposition of Corg-rich sediments. In the Vocontian Basin, the late Early Aptian OAE 1a black shales are associated with times of low detrital input, probably due to sea-level fluctuations and/or a shift towards more arid climate conditions. In order to investigate the causes and consequences of long-term climatic and floral change during the mid-Cretaceous, coastal sediments from Southern and Western Portugal (Algarve Abstract 4 and Lusitanian Basin) serve as environmental archives. The studied sections are Late Barremian to Middle Albian in age. A revised stratigraphic framework has been established for both sections using dinoflagellate cyst biostratigraphy. In order to obtain a terrestrial carbon isotope record for the Algarve section, the δ13C signature of fossil wood, cuticles, 13 charcoal and bulk Corg was measured. The distinct δ C pattern of the resulting record allows for chemostratigraphic correlation with existing carbon isotope curves, resulting in a significant enhancement of the stratigraphic resolution. Subsequently, the accurately dated successions are studied from a palynological perspective, with special emphasis on the qualtitative and quantitative analysis of the occurring angiosperms (flowering plants) pollen. A distinct increase in diversity and relative abundance of angiosperm pollen in the Barremian to Albian interval is observed in both studied sections, reflecting the incipient radiation of angiosperms on a resolution not obtained so far. Our results shed new light on the age interpretation of the well-known angiosperm mesofossil floras from the Portuguese Estremadura region, which have been assigned to a Barremian or possibly Aptian age. Several lines of evidence, including sequence- and biostratigraphy as well as palynology indicate an Albian or younger age for the mesofossil assemblages, hence indicating a major radiation phase during the Early Albian. Zusammenfassung 5

Zusammenfassung

Die mittlere Kreidezeit (Apt bis Turon, 120-90 Ma) war durch höhere globale Durchschnittstemperaturen von bis zu 10ºC gekennzeichnet und wird als eines der besten Beispiele für erdgeschichtliche Treibhausklima-Perioden betrachtet. Dies zeigt sich sowohl in einem geringen latitudinalen Temperatur-Gradienten und erhöhten ozeanischen Tiefen- und Oberflächenwasser-Temperaturen und als auch im Auftreten thermophiler Pflanzenver- gesellschaftungen in hohen Breiten und weitgehend eisfreien Polen. Doch auch während dieser globalen Warm-Phase waren die klimatischen Bedingungen keineswegs durchwegs stabil und ausgeglichen. Kurzzeitige Störungen des globalen Kohlenstoff-Kreislaufs sowie damit einhergehende klimatische Schwankungen sind in der Ablagerung organisch-reicher Schwarzschiefer, dem Kohlenstoff-Isotopensignal sowie in dramatischen Wachstumskrisen biokalzifizierender Organismen dokumentiert. Palynologische sowie Isotopen- und organisch- geochemische Untersuchungen an sedimentären Abfolgen aus dem Zeitraum Spät-Barrême bis Alb erlauben es, die Auswirkungen dieser globalen Veränderungen auf terrestrische Ökosysteme im Detail zu studieren. Um kurzfristige paläo-klimatische und -ozeanographische Veränderungen während einer Schwarzschiefer-Phase im späten Unter-Apt zu untersuchen, wurde der OAE 1a Horizont (oceanic anoxic event 1a) im Vocontischen Becken, SE Frankreich als hoch-auflösendes Umweltarchiv ausgewählt. Gestützt auf δ13C Analysen von marinen Karbonaten sowie von einzelnen organischen Verbindungen wurde eine Abschätzung des CO2 Partialdrucks während und nach der Schwarzschiefer-Phase durchgeführt. Zusätzlich wurde eine Analyse der auftretenden Pollen und Sporen Vergesellschaftung in den hemipelagischen Sedimenten durchgeführt, um Rückschlüsse auf mögliche Änderungen der Vegetation im Hinterland des Vocontischen Beckens zu erhalten. Darüber hinaus wurden Dinoflagellaten Zysten und

Nannoplankton Assoziationen bestimmt sowie Corg-Akkumulationsraten abgeschätzt, um Veränderungen der paläozeanographischen Bedingungen zu identifizieren. Die Resultate dieser Untersuchungen zeigen, dass trotz der verstärkten Ablagerung Corg-reicher

Schwarzschiefer nur ein sehr moderater Abfall des atmosphärischen CO2 Partialdrucks stattfand. Desweiteren weisen die analysierten Pollen-Spektren auf ein relativ stabiles Vegetationsmuster hin - sowohl in Zeiten verstärkter Schwarzschiefer Bildung als auch in den darüber folgenden normal-marinen Sedimenten. Die Auswertung des marinen organischen und kalkigen Planktons deutet ebenfalls auf relativ stabile paläozeanographische Bedingungen während und nach der OAE 1a Schwarzschiefer-Phase hin. Im Gegensatz zu früheren Studien Zusammenfassung 6 des OAE 1a fanden sich in der untersuchten Abfolge keine Hinweise auf ein direkte Verknüpfung von erhöhter Humidität, verstärktem Nährstoff-Eintrag sowie einer daraus resultierenden Produktivitätszunahme im ozeanischen Oberflächenwasser, welche sich wiederum in der Ablagerung von Corg-reichen Sedimenten äusserte. Vielmehr konnte festgestellt werden, dass die Bildung der OAE 1a Schwarzschiefer im Vocontischen Becken mit Phasen geringen detritischen Eintrags einherging, möglicherweise ausgelöst durch Meeresspiegel Schwankungen und/oder aride Klimabedingungen. Um die langerfristigen Ursachen und Auswirkungen des mittel-kretazischen Klima- und Florenwechsels zu untersuchen, wurden fossile Küsten-Sedimente in Süd- und West-Portugal (Algarve und Lusitanisches Becken) untersucht. Die beiden Profile reichen vom Spät- Barrême bis ins Mittlere Alb, die genaue zeitliche Einordung wurde mittels Dinoflagellaten Zysten-Biostratigraphie wesentlich verbessert. Messungen der δ13C Signatur von fossilen 13 Holzresten, Blatt-Kutikulen und Gesamt-Corg ermöglichen die Erstellung einer δ C Kurve für das Algarve Profil, welche eine chemostratigraphische Korrelation mit existierenden Isotopen-Kurven erlaubt und zu einer deutlich verbesserten stratigraphischen Auflösung der sedimentären Abfolge führt. Daran anschliessend werden palynologische Analysen der neu- datierten Sedimente durchgeführt, wobei der Schwerpunkt hierbei auf der qualitativen und quantitativen Auswertung der auftretenden Angiospermen (Blütenpflanzen) Pollen liegt. Eine deutlich Zunahme sowohl in der Diversität als auch in der relativen Häufigkeit der Angiospermen Pollen vom Barrême bis ins Alb zeigt sich in beiden untersuchten Profilen und dokumentiert in bisher nicht vorhandener zeitlicher Auflösung die Radiation der frühen Angiospermen. Aus kontinentalen Serien (Barrême bis Apt) der west-portugiesischen Estremadura Region wurde bereits früher ein Anzahl mesoskopischer Angiospermen Fossilien beschrieben, welche einen wesentlichen Beitrag zur Klärung der frühen Phylogenese dieser Gruppe leisteten. Im Vergleich mit unseren gut-datierten palynologischen Befunden zeigt sich jedoch, dass jene Angiospermen Reste ein wesentlich jüngeres stratigraphisches Alter besitzen als bisher angenommen. Chapter 1 7

Chapter 1

Introduction

1. The mid-Cretaceous: a time of global change

The mid-Cretaceous period (Aptian to Turonian, 120 to 90 Ma) offers the opportunity to study earth’s climate and its variability during times of exceptional warmth. The early Late Cretaceous is considered to reflect the warmest conditions during the last 145 Ma and represents one of the best examples of “greenhouse” climate conditions in the Phanerozoic Earth history (Barron, 1983). Substantial evidence for this exceptional climate mode includes increased surface and bathyal ocean water temperatures (Huber et al., 1999; Norris and Wilson, 1998), low equator-to-pole temperature gradients (Huber et al., 1995) as well as extensive forests in polar regions (Francis and Frakes, 1993; Spicer and Parrish, 1986) and the absence of expanded polar ice sheets. According to the Cretaceous Climate Ocean Dynamics (CCOD) workshop report (Bice et al., 2003) globally averaged surface temperatures in the mid-Cretaceous were more than 10ºC higher than today. Apart from a long-term rise in global mean temperatures, the study of mid-Cretaceous sediments provides evidence for several transient events of climatic and oceanographic perturbations. A multitude of sedimentological, geochemical and palaeotological data provides evidence for prominent fluctuations of the thermal and chemical state of the Cretaceous oceans and continents. Episodes of climatic cooling are reflected in the occurrence of ice-rafted debris and glacial deposits (Frakes and Francis, 1988; Price, 1999) as well as in shifts of the stable isotope records (Stoll and Schrag, 1996; Weissert and Lini, 1991). The episodic and widespread deposition of organic carbon-rich black shales (Oceanic Anoxic Events (OAEs) of Schlanger and Jenkyns, 1976), the drowning of carbonate platforms (Weissert et al., 1998) and concomitant shifts in the carbon isotope record (e.g. Scholle and Arthur, 1980) display sustained disturbances of the global carbon cycle. These observations question the long-held view of an equable and stable mid-Cretaceous climate mode and suggest the occurrence of severe short-term perturbations of the entire ocean-atmosphere system. These changes are superimposed on a gradual warming trend, resulting in exceptional greenhouse conditions of the early Late Cretaceous. Chapter 1 8

2. Causes and consequences of the environmental changes

The ultimate cause for the mid-Cretaceous environmental perturbations is still a matter of debate. Variations in the atmospheric composition are suggested to have played a key role for the observed climatic changes. Results from geochemical modelling (Berner, 1994) are in agreement with geochemical and stomatal-derived pCO2 estimates (Beerling and Royer,

2002) indicating strongly increased pCO2 levels (4 to 10 times preindustrial levels) during the Aptian to Turonian greenhouse period. Accoding to several authors (e.g. Arthur et al., 1985; Larson and Erba, 1999), the increase in greenhouse gases was triggered by extensive submarine volcanic activity, including enhanced spreading along mid-ocean ridges and the formation of Large Igneous Provinces (LIP) oceanic plateaus (e.g. the Ontong Java and Manihiki Plateaus). Additional greenhouse forcing could have been triggered by the concomitant and rapid dissociation of methane gas hydrates trapped in marine sediments (e.g. Beerling et al., 2002).

-75ºE -60ºE-45ºE -30ºE -15ºE 0ºE

Southern Laurasian province (subtropical to warm-temperate)

30ºN 3 proto North Atlantic 1 2 western Tethys transitional zone

15ºN

Northern Gondwana province (arid to semi-arid) 0ºN

EAG

-15ºN

Fig. 1: Palaeogeographic reconstruction of the North Atlantic and Tethyan realm during the mid- Cretceous at ~115 Ma (modified after Geomar map generator; www.ods.de). Asterisks mark the location of the study sites: (1), Lusitanian Basin; (2) Algarve Basin; (3) Vocontian Basin. Major floral belts and corresponding climates after Brenner (1976) and Chumakov et al. (1995). EAG: Equatorial Atlantic Gateway. Chapter 1 9

On longer time-scales, plate-tectonic forcing has been invoked as an important trigger mechanism for the observed perturbations (Fig. 1). The mid-Cretaceous rifting of South America and Africa and the concomitant development of the Equatorial Atlantic Gateway (EAG) is supposed to have caused a major reorganisation in oceanographic circulation and climatic patterns (Kuypers et al., 2002; Wagner and Pletsch, 1999). Based on coupled ocean- atmosphere model simulations, Poulsen et al. (2003) demonstrated that the onset of the Cretaceous thermal maximum was directly related to the tectonic evolution of the proto- Atlantic. This major tectonic rearrangement presumably resulted in significant long-term climatic changes and had a strong impact on temperature and precipitation patterns, and consequently on weathering and erosion processes as well as on the distribution of vegetation (Hallam, 1985; Weissert et al., 1998). According to Chumakov et al. (1995), the establishment of an equatorial humid belt during the Albian was probably triggered by the opening of the South Atlantic Ocean.

The response to the above mentioned processes is reflected in different short- and long-term perturbations of various parts of the mid-Cretaceous ocean-atmosphere system. One of the best-studied intervals of past oceanographic and climatic change is the late Early Aptian OAE 1a, lasting for about 0.5 to 1.0 Ma. The OAE 1a represents the first globally distributed black shale in the Cretaceous and is therefore regarded as a turning point in mid-Cretaceous palaeoceanography. Shifts in the δ13C signature of marine sediments deposited during and after the black shale event have been interpreted in terms of increased burial of organic carbon in marine sediments or reflecting changes in partitioning of carbonate and organic carbon (Arthur et al., 1988; Weissert et al., 1998). Disturbances of the Early Aptian carbon-cycle are furthermore displayed in major growth crises of carbonate-producing organisms, reflected in the demise of carbonate platforms (Wissler et al., 2003) and a pronounced decline in calcareous nannoplankton (nannoconnid crisis of Erba, 1994). According to Erbacher et al. (1996) and Leckie et al. (2002), the OAE 1a is accompanied by dramatic turnovers in siliceous and calcareous plankton due to changes in palaeofertility during episodes of black shale deposition. On a longer timescale - in the order of several millions of years - the establishment of greenhouse climate conditions during the mid-Cretaceous, with peak warmth in the Turonian (e.g. Wilson et al., 2002), probably reflects the combined effects of tectonic rearrangement Chapter 1 10

and concomitant CO2 forcing. Long-term changes of the prevalent weathering and erosion processes on the continents are reflected in varying clay-mineral compositions and sedimentation patterns (Ruffell and Batten, 1990, Wortmann et al., in press). According to Haq et al. (1987), the increase in global mean temperatures was accompanied by a stepwise rise in sea-level during the Aptian to Cenomanian interval.

3. Response of terrestrial environments to short- and long-term perturbations

Detailed information on the mid-Cretaceous climatic and carbon-cycle perturbations are mainly based on marine records from DSDP/ODP cores and from on-land sections. The available studies comprise a multitude of geochemical and micropalaeontological data addressing the thermal state, the palaeofertility and the circulation patterns of mid-Cretaceous oceans. In contrast, only few studies have been carried out with focus on the response of terrestrial ecosystems to short- and long-term changes. Land plant communities are sensitive recorders of changes in the physical environment. Their composition and spatial distribution is strongly influenced by variations in regional precipitation and temperature patterns. Hence, the study of palynofloral associations (pollen and spores) represents an important proxy for the investigation of past climate and environmental conditions on different time scales. Whereas the palynological approach is widely applied for the reconstruction of Quaternary and Neogene climates (Bradley, 1999 and references therein), only very few high-resolution data- sets exist for the Mesozoic (e.g. Hochuli et al., 1999; Looy et al., 2001).

Via the consumption of atmospheric CO2, terrestrial vascular plants are directly connected to the global carbon cycle. Prominent changes in the carbon isotopic composition of the carbon pool are not only reflected in marine-derived carbon but also in organic carbon of land-plant origin. Consequently, the δ13C composition of land-plant remains can be used to trace major shifts of the global carbon isotope record allowing for correlation of marine and terrestrial strata (Gröcke et al., 1999; Hesselbo et al., 2002). On longer time scales, the burial of terrestrial biomass along continental margins represents a major carbon sink and is therefore considered to play a key role in the global carbon cycle. Due to the different δ13C composition of terrestrial and marine organic carbon, intensified burial or oxidation of continental biomass can result in major short- and long-term shifts of the global carbon isotope record, e.g. during the Palaeocene/Eocene (Kurtz et al., 2004) or the Carboniferous (Beerling and Royer, 2002). Chapter 1 11

A similar mechanism has been suggested by Wissler (2001) to account for the Aptian δ13C anomaly. From a palaeobotancial perspective, the mid-Cretaceous period is characterised by the evolution and rapid diversifications of the flowering plants (Fig. 2). Early evidence for the occurrence of flowering plants (angiosperms) has been reported by Brenner (1996) who documented angiosperm-type pollen grains from supposedly Valanginian to Hauterivian deposits of Israel. The palaeogeographic dispersal of early angiosperm pollen suggests a latitudinally diachronous pattern. Angiosperms probably occurred first in the palaeoequatorial regions of Northern Gondwana (Fig. 1) and subsequently migrated towards northern and southern high-latitudes, were they appeared some 20 to 30 Ma later (Brenner, 1976; Crane and Lidgard, 1989). By the end of the Cenomanian, angiosperms dominated the diversity of low-latitude floras, accounting for ~70 % of species (Crane et al., 1995; Lidgard and Crane, 1988). Palaeo-botanical and -ecological interpretations of fossil angiosperm remains indicate that early angiosperm plants were of low stature, perhaps herbs or woody shrubs, which flourished predominantly in unstable environments (Crane et al., 1995; Friis et al., 1999; Wing and Boucher, 1998).

600

500

400

Cycadales 300 Angiosperms

number of species 200 Pteridophytes Fig. 2: Absolute species diversity of 100 Conifers Cretaceous macrofossil plant assemblages

Ginkgoales (redrawn from Lidgard and Crane 1988). 0 U Jur Neocom Ba-Ap Alb Ce T-S Cmp Ma Pal Note the dramatic increase in the number of 160 140 120 100 80 60 angiosperm taxa from the Albian onwards. time (Ma)

Many aspects of the early angiosperm radiation during the mid-Cretaceous are still ambiguous. In particular, the Barremian to Albian phase of the diversification is poorly Chapter 1 12 documented with regard to timing, diversity and relative abundance. The problematic age assignment of many records hampers detailed comparison and correlation with other assemblages as well as with major climatic and/or tectonic changes. According to several authors (e.g. Crane et al., 1995; Lupia et al., 2000) unstable environmental conditions during the mid-Cretaceous might have had significant influence on the evolution of flowering plants.

4. Main objectives and general outline

The purpose of this study is to investigate the response of terrestrial ecosystems to short- and long-term environmental changes during the mid-Cretaceous. Sedimentary deposits from SE France and Portugal are chosen as archives for the past perturbations, which are studied with palynological and geochemical methods. The presented thesis is closely connected to ongoing research on the impact of mid-Cretaceous carbon-cycle perturbations on shallow water carbonate systems, currently carried out by Stefan Burla at the ETH Zürich. The following two main objectives are addressed in this thesis.

(i) Tracing environmental change during times of late Early Aptian black shale formation The first two chapters focus on the climatic and oceanographic perturbations which are accompanied by the formation of the late Early Aptian OAE 1a. The Niveau Goguel interval of the Serre Chaitieu section (Vocontian Basin, SE France) represents a well-documented equivalent of the OAE 1a black shale (Bréhéret, 1997; Herrle and Mutterlose, 2003) and has been sampled on a high resolution. The hemipelagic deposits of this section provide well- preserved organic matter and palynomorphs, allowing for detailed analysis of the organic geochemistry and palynological assemblages.

(ii) Tracing patterns of early angiosperm radiation during the Barremian-Albian interval The second part of the study addresses long-term changes of the mid-Cretaceous carbon-cycle and vegetation patterns with special focus on the diversifying angiosperms. Coastal marine deposits from the Portuguese Algarve and Estremadura regions, covering Barremian to Albian strata have been chosen as environmental archives (Rey, 1972; Rey, 1986). Both successions provide well-preserved land plant-derived organic matter, including cuticles, fossil wood and excellent pollen assemblages. The chosen study sites are located close to a number of well- known and intensely studied angiosperm mesofossil sites in the Estremadura region (Friis et al., 1994; Friis et al., 1999). Chapter 1 13

Chapters 2 to 5 represent discrete manuscripts, which are either published, in review or in preparation for publication.

In Chapter 2 a combined geochemical and palynological approach is applied to study variations in palaeoatmospheric CO2 concentrations and concomitant floral changes across the OAE 1a interval (Vocontian Basin, SE France). The δ13C composition of carbonate and organic carbon as well as of individual biomarkers is used to estimate past changes in pCO2. A detailed chemostratigraphic correlation with an existing, more pelagic record (Cismon, N Italy) allows for comparison of pollen assemblages from two different sites. Possible consequences for the palaeoceanographic and palaeoatmospheric conditions are discussed.

Chapter 3 assesses variations of the pollen assemblage, the organic-walled plankton and the calcareous nannofossils across the OAE 1a interval (Vocontian Basin, SE France). The marine and terrestrial-derived microfossils serve as proxies for past climatic and oceanographic change during times of black shale formation. In combination with tentative estimations of sedimentation rates and organic carbon fluxes, the palynological and nannofossil results contrast to previously proposed scenarios for the formation of late Early Aptian black shales.

In Chapter 4 the carbon isotopic composition of land plant-derived organic material from two coastal marine records (Algarve Basin, S Portugal) is analysed. The obtained δ13C records display several distinct shifts, which allow for correlation on an intrabasinal as well as on a global scale with existing Aptian carbon isotope curves. In combination with biostratigraphic data, the applied method results in a significant enhancement of the stratigraphic resolution of the studied records.

Chapter 5 addresses the radiation of early angiosperms within the Barremian to Albian interval from a palynological perspective. The angiosperm pollen records of two well-dated sections (Lusitanian and Algarve Basins, Portugal) are analysed with respect to composition, abundance and diversity and compared with previously published records. The implications for the timing of the early angiosperm diversification are discussed including a revised age assignment for several angiosperm mesofossil floras from the Portuguese Estremadura region.

Chapter 1 14

References

Arthur, M.A., Dean, W.E. and Pratt, L.M., 1988. Geochemical and climatic effects of increased marine organic carbon burial at the Cenomanian/Turonian boundary. Nature, 335, 714-717. Arthur, M.A., Dean, W.E. and Schlanger, S.O., 1985. Variations in the global carbon cycle during the Cretaceous related to climate, volcanism, and changes in atmospheric CO2. In: E.T. Sundquist and W.S. Broecker (Editors), The carbon cycle and atmospheric CO2 :Natural variations Archean to present. Geophysical Monograph. American Geophysical Union, Washington, pp. 504-529. Barron, E.J., 1983. A warm, equable Cretaceous: The nature of the problem. Earth Science Reviews, 19, 305-338. Beerling, D.J., Lomas, M.R. and Gröcke, D.R., 2002. On the nature of methane gas-hydrate dissociation during the Toarcian and Aptian oceanic anoxic events. American Journal of Science, 302, 28-49. Beerling, D.J. and Royer, D.L., 2002. Fossil plants as indicators of the phanerozoic global carbon cycle. Annual Review of Earth and Planetary Sciences, 30, 527-556. Berner, R.A., 1994. GEOCARB II: A revised model of atmospheric CO2 over Phanerozoic time. American Journal of Science, 294, 56-91. Bice, K.l., Bralower, T.J., Duncan, R.A., Huber, B.T., Leckie, R.M. and Sageman, B.B., 2003. http://www.whoi.edu/ccod/CCOD_report.html. Bradley, R.S., 1999. Paleoclimatology: reconstructing climates of the Quaternary. Internationl Geophysics Series, 64. Harcourt Academic Press, San Diego, 613 pp. Bréhéret, J.G., 1997. L'Aptian et l'Albian de la fosse vocontienne (bordures au bassin): Évolution de la sédimentation et enseignements sur les événements anoxiques. Publications de la Société Géologique du Nord, 25, 614. Brenner, G., 1976. Middle Cretaceous floral provinces and early migration of Angiosperms. In: C.B. Beck (Editor), Origin and early evolution of angiosperms. Columbia University Press, New York, pp. 23-44. Brenner, G.J., 1996. Evidence for the earliest stage of angiosperm pollen evolution: a paleoequatorial section from Israel. In: D.W. Taylor and L.J. Hickey (Editors), Flowering plant origin, evolution and phylogeny. Chapman & Hall, pp. 91-115. Chumakov, N.M., Zharkov, M.A., Herman, A.B., Doludenko, M.P., Kalandadze, N.M., Lebedev, E.L., Ponomarenko, A.G. and Rautian, A.S., 1995. Climatic belts of the mid-Cretaceous time. Stratigraphy and Geological Correlation, 3, 241-260. Crane, P.R., Friis, E.M. and Pedersen, K.R., 1995. The origin and early diversification of angiosperms. Nature, 374, 27-33. Crane, P.R. and Lidgard, S., 1989. Angiosperm diversification and paleolatiudinal gradients in Cretaceous floristic diversity. Science, 246, 675-678. Erba, E., 1994. Nannofossils and superplumes: The early Aptian "nannoconid crisis". Paleoceanography, 9, 483-501. Erbacher, J., Thurow, J. and Littke, R., 1996. Evolution patterns of radiolaria and organic matter variations: A new approach to identify sea-level changes in Mid-Cretaceous pelagic environments. Geology, 24, 499-502. Frakes, L.A. and Francis, J.E., 1988. A guide to Phanerozoic cold polar climates from high-latitude ice-rafting in the Cretaceous. Nature, 333, 547-549. Francis, J.E. and Frakes, L.A., 1993. Cretaceous climates. In: V.P. Wright (Editor), Sedimentology Review. Blackwell Scientific Publications, Oxford, pp. 17-30. Friis, E.M., Pedersen, K.R. and Crane, P.R., 1994. Angiosperm floral structures from the Early Cretaceous of Portugal. Plant Systematics and Evolution, 8 (supplement), 31-49. Friis, E.M., Pedersen, K.R. and Crane, P.R., 1999. Early angiosperm diversification: the diversity of pollen associated with angiosperm reproductive structures in Early Cretaceous floras from Portugal. Annals of the Missouri Botanical Garden, 86, 259-296. Chapter 1 15

Gröcke, D.R., Hesselbo, S.P. and Jenkyns, H.C., 1999. Carbon-isotope composition of Lower Cretaceous fossil wood: Ocean-atmosphere chemistry and relation to sea-level change. Geology, 27, 155-158. Hallam, A., 1985. A review of Mesozoic climate. Journal of the Geological Society of London, 142, 433-445. Haq, B.U., Hardenbol, J. and Vail, P.R., 1987. Chronology of fluctuating sea levels since the Triassic. Science, 235, 1156-1167. Herrle, J.O. and Mutterlose, J., 2003. Calcareous nannofossils from the Aptian-Lower Albian of southeast France: Palaeoecological and biostratigraphic implications. Cretaceous Research, 24, 1-22. Hesselbo, S.P., Robinson, S.A., Surlyk, F. and Piasecki, S., 2002. Terrestrial and marine extinction at the Triassic-Jurassic boundary synchronized with major carbon-cycle perturbation: a link to initiation of massive volcanism? Geology (Boulder), 30, 251-254. Hochuli, P.A., Menegatti, A.P., Weissert, H., Riva, A., Erba, E. and Premoli Silva, I., 1999. Episodes of high productivity and cooling in the early Aptian Alpine Tethys. Geology, 27, 657-660. Huber, B.T., Hodell, D.A. and Hamilton, C.P., 1995. Middle-Late Cretaceous climate of the southern high latitudes: stable isotopic evidence for minimal equator-to-pole thermal gradients. Geological Society of America Bulletin, 107, 1164-1191. Huber, B.T., Leckie, R.M., Norris, R.D., Bralower, T.J. and CoBabe, E., 1999. Foraminiferal assemblage and stable isotopic change across the Cenomanian-Turonian boundary in the subtropical North Atlantic. Journal of Foraminiferal Research, 29, 392-417. Kurtz, A.C., Kump, L.R., Arthur, M.A., Zachos, J.C. and Paytan, A., 2004. Early Cenozoic decoupling of the global carbon and sulfur cycles. Paleoceanography, 18, 14 (1-13). Kuypers, M.M.M., Pancost, R.D., Nijenhuis, I.A. and Sinninghe Damste, J.S., 2002. Enhanced productivity led to increased organic carbon burial in the euxinic North Atlantic basin during the late Cenomanian oceanic anoxic event. Paleoceanography, 17, 1-13. Larson, R.L. and Erba, E., 1999. Onset of the mid-Cretaceous greenhouse in the Barremian-Aptian igneous events and the biological, sedimentary and geochemical responses. Paleoceanography, 14, 663-678. Leckie, R.M., Bralower, T.J. and Cashman, R., 2002. Oceanic anoxic events and plankton evolution: Biotic response to tectonic forcing during the mid-Cretaceous. Paleoceanography, 17, 13-1 - 13-29. Lidgard, S. and Crane, P.R., 1988. Quantitative analyses of the early angiosperm radiation. Nature, 331, 344-346. Looy, C.V., Twitchett, R.J., Dilcher, D.L., Konijnenburg-Van Cittert, J.H.A. and Visscher, H., 2001. Life in the end-Permian dead zone. Proceedings of the National Academy of Sciences of the United States of America, 98, 7879-7883. Lupia, R., Crane, P.R. and Lidgard, S., 2000. Angiosperm diversification and mid-Cretaceous environmental change. In: S.J. Culver and P.F. Rawson (Editors), Biotic response to global change: the last 245 million years. Cambridge University Press, Cambridge, United Kingdom, pp. 207-222. Norris, R.D. and Wilson, P.A., 1998. Low-latitude sea-surface temperatures for the mid-Cretaceous and the evolution of planktic foraminifera. Geology (Boulder), 26, 823-826. Poulsen, C.J., Gendaszek, A.S. and Jacob, R.L., 2003. Did the rifting of the Atlantic Ocean cause the Cretaceous thermal maximum? Geology (Boulder), 31, 115-118. Price, G.D., 1999. The evidence and implications of polar ice during the Mesozoic. Earth-Science Reviews, 48, 183-210. Rey, J., 1972. Recherches géologiques sur le crétacé inférieur de l'Estremadura (Portugal). Memoria (N.S.), 21. Serviços Geológicos de Portugal, Lisbon, 1-477 pp. Rey, J., 1986. Micropaleontological assemblages, paleoenvironments and sedimentary evolution of Cretaceous deposits in the Algarve (southern Portugal). Palaeogeography, Palaeoclimatology, Palaeoecology, 55, 233-246. Ruffell, A.H. and Batten, D.J., 1990. The Barremian-Aptian arid phase in Western Europe. Palaeogeography, Palaeoclimatology, Palaeoecology, 80, 197-212. Chapter 1 16

Schlanger, S.O. and Jenkyns, H.C., 1976. Cretaceous oceanic anoxic events: Causes and consequences. Geologie en Mijnbouw, 55, 179-184. Scholle, P.A. and Arthur, M.A., 1980. Carbon isotope fluctuations in Cretaceous pelagic limestones: Potential stratigraphic and petroleum exploration tool. American Association of Petroleum Geologists Bulletin, 64, 67-87. Spicer, R.A. and Parrish, J.T., 1986. Paleobotanical evidence for cool North Polar climates in Middle Cretaceous (Albian-Cenomanian) time. Geology, 14, 703-706. Stoll, H.M. and Schrag, D.P., 1996. Evidence for glacial control of rapid sea level changes in the Early Cretaceous. Science, 272, 1771-1774. Wagner, T. and Pletsch, T., 1999. Tectono-sedimentary controls on Cretaceous black shale deposition along the opening Equatorial Atlantic gateway (ODP Leg 159), The oil and gas habitats of the South Atlantic. Geological Society of London, London, United Kingdom, pp. 241-265. Weissert, H. and Lini, A., 1991. Ice age interludes during the time of Cretaceous greenhouse climate? In: D.W. Müller, J.A. McKenzie and H. Weissert (Editors), Controversies in modern geology: Evolution of geological theories in sedimentology, earth history and tectonics. Academic Press, London, UK, pp. 173-191. Weissert, H., Lini, A., Foellmi, K.B. and Kuhn, O., 1998. Correlation of Early Cretaceous carbon isotope stratigraphy and platform drowning events: A possible link? Palaeogeography, Palaeoclimatology, Palaeoecology, 137, 189-203. Wilson, P.A., Norris, R.D. and Cooper, M.J., 2002. Testing the Cretaceous greenhouse hypothesis using "glassy" foraminiferal calcite from the core of the Turonian tropics on Demerara Rise. Geology, 30, 607-610. Wing, S.L. and Boucher, L.D., 1998. Ecological aspects of the Cretaceous flowering plant radiation. Annual Review of Earth and Planetary Sciences, 26, 379-421. Wissler, L., 2001. Response of Early Cretaceous sedimentary systems to perturbations in global carbon cycling: insights from stratigraphy, sedimentology and geochemical modeling, Eidgenössische Technische Hochschule Zürich, Zürich, 109 pp. Wissler, L., Funk, H. and Weissert, H., 2003. Response of Early Cretaceous carbonate platforms to changes in atmospheric carbon dioxide levels. Palaeogeography, Palaeoclimatology, Palaeoecology, 200, 187-205.

Chapter 2 17

Chapter 2

Absence of major vegetation and palaeoatmospheric pCO2 changes associated with Oceanic Anoxic Event 1a (Early Aptian, SE France)

Abstract

The deposition of organic-rich sediments during the late Early Aptian Oceanic Anoxic Event

(OAE) 1a has been interpreted to result in a major decrease of palaeoatmospheric CO2 concentrations, accompanied by significant changes in the terrestrial flora. In order to test this hypothesis, the OAE 1a interval in the Vocontian Basin (SE France) has been studied with a combined approach including stable carbon isotopes, organic geochemistry and palynology. To 13 estimate changes in palaeoatmospheric CO2 levels across the OAE 1a, the δ C composition of presumed algal biomarkers (low-molecular-weight n-alkanes, steranes) and of bulk carbonate carbon are used. Our results yield estimated Early Aptian pCO2 values 3 to 4 times the preindustrial level and only a moderate drop across the black shale event. This moderate drop in pCO2 is supported by palynological results. The frequency patterns of climate-sensitive sporomorphs (incl. pteridophyte spores, bisaccate pollen and Classopollis spp.) display only minor fluctuations throughout the studied section and indicate relatively stable patterns of terrestrial vegetation during formation of the OAE 1a black shale. The occurrence of a characteristic Early Aptian carbon isotope pattern across the OAE 1a interval permits accurate chemostratigraphic correlation with the well-studied Livello Selli interval of the Cismon record (N Italy). The contemporaneous formation of individual black shale layers at both sites indicates that transient episodes of dysoxic-anoxic bottom waters prevailed over large areas in the W Tethys Ocean independent of depositional setting. Comparison of the palynological data from the two locations displays significant differences in the frequency patterns of bisaccate pollen. The contrasting pollen spectra are interpreted to reflect prominent changes in the palaeoceanographic current patterns and/or selective sorting due to sea level rise rather than latitudinal shifts of the major floral belts.

Keywords: Early Cretaceous; Aptian; black shales; OAE; carbon isotopes; palynology; organic geochemistry; palaeoatmosphere

Chapter 2 18

1. Introduction

1.1. Palaeoclimatic and palaeoceanographic background conditions

The mid-Cretaceous is generally referred to as a greenhouse period characterized by exceptionally warm climates (Hallam, 1985; Wilson and Norris, 2001), a weak meridional temperature gradient (Huber et al., 1995) and considerably high levels of atmospheric carbon dioxide (Berner, 1994; Freeman and Hayes, 1992). These extraordinary climatic conditions are also reflected in the composition and spatial distribution of terrestrial plant assemblages. Mid-Cretaceous fossil floras typically include ferns, conifers and cycadophytes which grew throughout low to high latitudes, indicating tropical/subtropical to warm temperate conditions (Hallam, 1985). The occurrence of extensive forests dominated by podocarpian and araucarian conifers and other thermophilic taxa in polar regions in combination with the absence of expanded polar ice sheets points to more equable and warmer climates during the mid-Cretaceous in comparison to the present-day situation (Francis and Frakes, 1993). During this period of greenhouse conditions, sedimentation in the world oceans was characterized by the episodic deposition of organic carbon-rich sediments, informally called “black shales”. The relative short-lived episodes (~ 50 to 500 ka) of organic carbon (OC) accumulation in pelagic and hemipelagic environments were of regional to global extent and have been termed Oceanic Anoxic Events (OAE) by Schlanger and Jenkyns (1976). The Early Aptian OAE 1a represents the first globally distributed black shale event and therefore is regarded as a major turning point of mid-Cretaceous palaeoceanography. The OAE 1a is accompanied by dramatic turnovers in calcareous nannoplankton (Erba, 1994) and by high extinction and origination rates of siliceous and calcareous plankton (Erbacher and Thurow, 1997; Leckie et al., 2002). In addition, a phase of carbonate platform demise has been documented from the northern Tethys margin as well as from circum-Atlantic regions (Weissert et al., 1998) which slightly predates the OAE 1a. Prominent changes in the global carbon budget during and after times of black shale formation are reflected in the 13C/12C ratio 13 of organic (Corg) and carbonate carbon (Ccarb). The resulting δ C pattern is characteristic for Early Aptian times and has been documented worldwide from marine successions (Bralower et al., 1999; Herrle et al., 2004; Menegatti et al., 1998) as well as from terrestrial environments (Gröcke et al., 1999; Heimhofer et al., 2003). According to Larson and Erba (1999) the mid-Cretaceous period of global warmth was triggered by extensive submarine volcanic activity, including increased spreading rates along

Chapter 2 19 mid-ocean ridges and the formation of extensive oceanic plateaus termed “Large Igneous Provinces”. The Early Aptian OAE 1a slightly postdates a period of intensive volcanic activity on the Ontong Java Plateau and the Manihiki Plateau in the western Pacific between 125 and 120.5 Ma (Larson and Erba, 1999). The accompanying volcanic degassing may have resulted in exceptionally high atmospheric carbon dioxide levels and related mid-Cretaceous greenhouse warming (Arthur et al., 1985). Additional greenhouse forcing could have been triggered by the rapid dissociation of isotopically-light methane gas-hydrates as indicated by a negative δ13C anomaly at the onset of OAE 1a (Beerling et al., 2002). The accumulation and burial of large quantities of OC in sediments during OAEs is assumed to result in a significant drop in atmospheric carbon dioxide partial pressure (pCO2) and consequent climate cooling in the aftermath of these events (Arthur et al., 1988; Kuypers et al., 1999). A prominent decline in mean annual temperatures during or after OAE 1a formation is expected to affect terrestrial vegetation patterns significantly. A first detailed spore-pollen record across OAE 1a (Hochuli et al., 1999) shows a significant increase in boreal floral elements following the interval of black shale formation and has been interpreted to reflect a major cooling episode and/or prominent changes in oceanographic circulation patterns of the SW Tethys.

1.2. Aim of the study

To address possible changes in pCO2 across the OAE 1a and its potential impact on terrestrial ecosystems we chose a two-fold approach. Organic and inorganic carbon isotope geochemistry is used to obtain estimates of pCO2 based on variations in the photosynthetic fractionation factor of marine phytoplankton. In addition, palynological analysis of climate- sensitive floral elements offers the opportunity to study climatically induced variations in past vegetation patterns. The distribution and abundance of pteridophyte spores, Classopollis spp. and bisaccate pollen are strongly controlled by the prevailing palaeoclimatic conditions (Batten, 1984; Vakhrameyev, 1982) and therefore can serve as indicator for palaeoclimatic variations. However, in distal sedimentary facies, selective sorting of spores and pollen can have a strong effect on the palynological composition and therefore has to be considered carefully (Traverse, 1988; Tyson, 1995). In order to study the onset of the positive carbon isotope excursion and the presumed palaeoclimatic changes from OAE to post-OAE conditions, the upper part of the OAE 1a interval and the overlying sediments were analysed from the Serre Chaitieu section (Vocontian Basin, SE France). Based on chemostratigraphic

Chapter 2 20 correlation, our geochemical and palynological results were compared in great detail with data from the Cismon section (Belluno Trough, N Italy).

2. Palaeogeographic setting

The Serre Chaitieu section is situated in the north-eastern part of the Vocontian Basin (SE France), which formed part of the northern continental margin of the Alpine Tethys Ocean (Fig. 1). A palaeolatitude of 25-30°N has been inferred for the mid-Cretaceous position of the basin (e.g. Hay et al., 1999). The section studied represents the lowest part of the “Marnes Bleues” formation, a monotonous, up to 750 m thick succession of Aptian to Albian age, which is composed mainly of grey to dark-grey marls intercalated with calcareous marls and limestones. Numerous organic-rich shale horizons occur throughout the succession, some of which can be correlated on a Tethys-wide or even global scale (Bréhéret, 1988).

Drowned platform facies B Shallow open marine environments C 40° Deep open marine environments pre-Triassic basement

Massif Central Grenoble Internal zones Valence of the Alps Southern Laurasian province Die (subtropical to warm-temperate) 30°

Serre Chaitieu Roter Sattel section Serre Chaitieu

Digne Luz Cismon Castellane 20° Provence Platform Nice A N Inte rmed Sea iate flo Northern Gondwana province ral France be terranean (arid to semi-arid) lt Medi 0 50 km N 10° B Spain 0 500 km

Fig. 1. A: Location of the Vocontian Basin in SE France. B: Spatial distribution of different depositional settings within the Vocontian Basin during the mid-Cretaceous. The location of the Serre Chaitieu section is marked with an asterisk. Map modified after Arnaud and Lemoine (1993). C: Plate tectonic reconstruction of the W Tethys realm during the mid-Cretaceous (~115 Ma). Floral provinces and inferred climates after Brenne (1967). Positions of the study site (black asterisk) and of the sections used for comparison (white asterisks) are marked. Map modified after Geomar map generator (www.odsn.de).

Chapter 2 21

During late Early Aptian times, the study area was situated in the north-western part of the Vocontian Basin, several tens of km south of the northern palaeo-margin (Arnaud and Lemoine, 1993). To the east the basin opened towards the Tethys Ocean facilitating exchange with Tethyan water masses. Although deposition within the basin was largely pelagic, the study area received a considerable amount of terrigenous detrital material from the nearby continental areas (Bréhéret, 1994). The studied basin was located in the southern part of the Southern Laurasian floral province of Brenner (1976), which was restricted to the mid-latitudes of the Northern Hemisphere during Aptian to Albian times (Fig. 1). Its microfloral assemblage is characterized by a high diversity and abundance of pteridophyte spores - namely by spores of gleicheniaceous and schizaeaceous affinity - and bisaccate pollen of pinaceous and podocarpaceous origin. Other gymnosperm pollen such as Classopollis (Cheirolepidiaceae) and Araucariacites (Araucariaceae) represent common elements of its floral assemblages (Batten, 1984; Brenner, 1976). During the mid-Cretaceous the major floral belts were broadly latitudinally arranged and exhibited progressive compositional changes with increasing latitude, despite low equator-to-pole thermal gradients (Batten, 1984). According to Brenner (1976) and Vakhrameyev (1978) the Southern Laurasian floral province was characterized by a subtropical to warm-temperate climate, whereas the climate of the Northern Gondwana province adjacent to the south is regarded as tropical to semi-arid. In contrast, the Northern Laurasian floral province (situated north of 60° N) is dominated by bisaccate pollen of pinaceous origin indicating temperate and humid conditions.

3. Lithology and stratigraphy of the studied Serre Chaitieu section (SE France)

The Serre Chaitieu section, located about 1 km south of the village Lesches-en-Diois has been studied in detail by several authors (Bréhéret, 1988; 1997; Herrle and Mutterlose, 2003). The interval included in this study encompasses 12 m and represents the lowermost part of the Serre Chaitieu section. The interval below the base of the sampled succession was not accessible. The studied section is mainly composed of dark-grey marls with low to moderate carbonate (12.0 to 36.0 %) and organic carbon (0.5 to 1.8 %) content, respectively. The lower part of the section (0 to 6.5 m) shows elevated total organic carbon (TOC) values. The marls are highly bioturbated with Chondrites and Planolites as the most common trace fossils (Bréhéret, 1997). Intercalated within the homogenous marly succession, 6 distinctly laminated paper shale horizons (PS-1 to 6) ranging in thickness from 20 cm to 35 cm can be observed.

Chapter 2 22

These finely laminated horizons yield higher contents of TOC (up to 2.3 %) and are essentially devoid of bioturbation. According to Bréhéret (1994) the four lowermost paper shale horizons are referred to as the Niveau Goguel interval, which corresponds to the OAE 1a. Considering the elevated TOC values, sediment thickness and chemostratigraphic results we assign the entire lower part of the section (0 to 6.5 m) to the OAE 1a interval. The studied interval lies within the Deshayesites deshayesi ammonite Zone (Bréhéret, 1997 and references therein) and within the middle part of the Leupoldina cabri planktic foraminiferal zone. It comprises the first occurrence of Eprolithus floralis which marks the onset of the NC7A (Rhagodiscus angustus) calcareous nannofossil subzone and the end of the NC6 (Chiastozygus litterarius) nannofossil zone in the Vocontian Basin (Herrle and Mutterlose, 2003). Based on time series analysis and biostratigraphic data, an average sedimentation rate of 3.0 to 3.5 cm/ka was calculated for the lowermost Late Aptian part of the Serre Chaitieu section by Kössler et al. (2001) and Herrle et al. (2003). Due to the condensed character of the intercalated paper shale horizons, mean sedimentation rates in the lowermost part of the section including the OAE 1a interval are probably even lower (2.5 to 2.0 cm/ka). According to Bréhéret (1994) the OAE 1a interval can be interpreted to reflect a major transgressive pulse or a maximum flooding (2nd order sequence). This interpretation correspond well with the comprehensive sequence-stratigraphic framework established by Hardenbol et al. (1998) for the major European basins. The Serre Chaitieu record from the Vocontian Basin is compared in detail with the Cismon section (N Italy), from which detailed bio-, magneto- and chemostratigraphic data are available (Erba et al., 1999 and references therein). A record of palynofacies and palynology has been published from the same section by Hochuli et al. (1999). The pelagic sediments of the Cismon section have been deposited at a palaeolatitude of ~ 20° N in the Belluno Trough. During the mid-Cretaceous this basin was situated on the northern continental margin of Apulia, approximately ~ 1000 km SE of the Vocontian Basin. Based on cyclostratigraphic studies of the Cismon section, the duration of the OAE 1a (Livello Selli equivalent) has been estimated between 500 ka to 1 Ma (Erba et al., 1999; Herbert, 1992) resulting in a mean sedimentation rate of 1.0 to 0.5 cm/ka.

Chapter 2 23

4. Material and Methods

4.1. Carbon isotope analysis and total organic carbon contents

To analyse the δ13C composition of bulk carbonate carbon, 36 powdered samples were treated with phosphoric acid at 90°C. Subsequently, the liberated CO2 gas was analysed with a VG PRISM mass spectrometer. For determination of δ13C of bulk organic carbon, samples were treated with 3 N HCl for 24h to remove the inorganic carbonates. About 40 mg of the residue were analysed via combustion in a CNS Elemental Analyser connected to an isotope ratio mass spectrometer (Optima/Micromass). All carbon-isotope ratios are expressed in the standard δ notation in per mil (‰) relative to the international VPDB isotope standard. The δ13C values of the carbonate carbon were calibrated against a laboratory internal standard (Carrara marble; δ13C = 2.14 ‰); analytical reproducibility was ± 0.05 ‰. For bulk organic carbon measurements, a laboratory internal standard (Atropina; δ13C = -28.48 ‰) and an international standard (NBS 22; δ13C = -29.74 ‰) were used; analytical reproducibility was better than ± 0.2 ‰. Inorganic carbon contents of 36 samples were determined using a UIC CM 5012 Coulomat; total carbon contents were measured on a CNS Elemental Analyser (Carlo Erba Instruments). Total organic carbon (TOC) contents were calculated from the difference between total and inorganic carbon contents.

4.2. Biomarker analysis and compound-specific carbon isotope analysis

An aliquot (10 g) of 15 powdered samples were extracted using an UP 200s ultrasonic disrupter probe (amplitude 50; cycle 0.5) and three successively less polar mixtures of methanol and dichloromethane, each for 3 min. After sulphur removal and desalination, the extracts were concentrated by rotary evaporation and evaporated under N2. Compound class fractions were separated by column chromatography using 7 g silica gel; the apolar fraction was eluted with 30 ml n-hexane; the polar fraction with a 1:1 mixture of dichloromethane and methanol. Samples were then analysed by gas chromatography-mass spectrometry for compound identification using a HP 6890 GC fitted with a HP-5MS column (30 m × 0.25 mm, df = 0.25 µm) and interfaced to a mass selective detector HP 5973. Carbon isotopic compositions of individual n-alkanes were determined using a TRACE GC fitted with a HP-1 column (50 m × 0.32 mm, df = 0.17 µm) and coupled to a Thermo Finnigan Plus Delta XL mass spectrometer. A series of 9 different n-alkanes (n-C19 to n-C40) was used as

Chapter 2 24 an internal working standard. Reported δ13C values represent the means of multiple analysis (n = 3) expressed versus VPDB. Except for one sample, standard errors of the mean were better than ± 0.7 ‰.

4.3. Palynological preparation

A total of 16 cleaned and weighed (10 – 12 g) samples were treated with hydrochloric and hydrofluoric acid following standard palynological preparation techniques (Traverse, 1988).

The residue was sieved over a 11 µm mesh-sieve and a short oxidation with HNO3 was performed on all residues. A minimum of 250 sporomorphs per sample (mean 254) were counted from strew mounts. Only three sporomorph categories were distinguished and the estimated standard deviation is expected to be better than ± 3% (Traverse, 1988).

5. Preservation of stable carbon isotope signals and organic matter

Strong diagenetic overprint of the carbonate carbon isotope signature can be excluded for the following reasons: (1) The shallow burial depth of the sedimentary succession (< 700m) according to Levert and Frey (1988). (2) The well preserved calcareous nannofossils with only minor contribution of cements and micrite observed in nannofossil samples (Herrle and 13 18 2 Mutterlose, 2003). (3) The lack of covariance between δ Ccarb and δ Ocarb (r = 0.04; n = 36).

Thermally unaltered conditions of the organic matter (OM) are inferred from Tmax values of 420 to 435°C (Bréhéret, 1994), the unaltered colour of the palynomorphs (TAI < 2) and the moderate to strong UV fluorescence of the amorphous OM fraction. Visually, the preservation of all palynomorphs is good to excellent throughout the interval studied. In addition, several biomarker maturity indices including the 22S/(22R + 22S)-hopane (C31) index, the Mor/(Mor + Hop) index and the Ts/(Ts + Tm) index confirm the immature stage of the organic matter with respect to hydrocarbon generation (Peters and Moldowan, 1993).

6. Results and discussion

13 6.1. Chemostratigraphic correlation of the δ Ccarb records from SE France and N Italy

The lower part of the Serre Chaitieu section (0 to 4.5 m) is characterized by an interval of 13 stable δ Ccarb values of ~ 3.0 ‰ with the lowest value of 2.4 ‰ occurring at the base of the 13 studied interval. Within the first paper-shale horizon (PS-1), the δ Ccarb signature shifts

Chapter 2 25

13 towards higher values of ~ 3.5 ‰. At 7.8 m the δ Ccarb curve displays a second positive shift and peaks at values of ~ 4.5 ‰ at 9.4 m within PS-5 (Fig. 2). 13 The distinct shifts in the δ Ccarb record in combination with accurate biostratigraphic data allow a detailed chemostratigraphic correlation between the two OAE 1a intervals from the Serre Chaitieu section (SE France) and the Cismon section (N Italy). In the lower part of the 13 Serre Chaitieu record the stable δ Ccarb values correspond well with the C5 segment of Menegatti et al. (1998) in the Cismon record (Fig. 2). Both records show a subsequent shift towards more positive values (C6) reaching the peak values of the Early Aptian δ13C positive excursion (C7). In contrast to Menegatti et al. (1998), the entire shift towards more positive values is included here in the C6 segment. Both carbon isotope curves display not only similar patterns, but also show essentially the 13 same absolute δ Ccarb values and a comparable positive shift of ~ 1.5 ‰ included in the C6 segment. The resulting correlations are in good agreement with the biostratigraphic data. In both records the first occurrence of Eprolithus floralis is situated in the uppermost part of the black shale interval (Fig. 2). The chemostratigraphic correlation clearly indicates the absence of a negative δ13C spike and the corresponding segments C2, C3 and, in part C4 in the Serre Chaitieu section. Sedimentary evidence for an incomplete transition between the uppermost Early Aptian limestone beds and the onset of the “Marnes Bleues” formation has been reported from other localities within the Vocontian Basin (e.g. Les Sauziere section) by Bréhéret (1997). Except from this basal hiatus there is no sedimentological or stratigraphical evidence for further gaps within the studied interval.

6.2. Organic matter composition and origin

The extractable hydrocarbons are dominated primarily by short-chain n-alkanes, acyclic isoprenoids and abundant steroidal and hopanoid hydrocarbons. In all samples studied, peak maxima are represented by pristane and phytane, followed by short-chain n-alkanes (n-C15 to n-C19) and steranes (C27 and C29). In contrast, long-chain n-alkanes (n-C27 to n-C33) with a relatively low odd-over-even predominance (OEP) of 1.4 – 1.7 form only a minor constituent. A significant algal contribution is suggested from the high abundance of short-chain n-alkanes (Farrimond et al., 1990; Gelpi et al., 1970) and steroidal components (Volkman, 1986).

Chapter 2 26

Cismon section (Menegatti et al. 1998) meters Formation Stage Foram.-zone Nanno.-zone Lithology

(NC7) C7 -258 Serre Chaitieu section (this study) Lithology Nanno.-zone Foram.-zone Stage Formation meters

R. angustus PS-6 12

C7

C6 (NC7A) 10

PS-5 Leupoldina cabri -263

8 R. angustus C5 C6 PS-4 6 C4 PS-3 Livello Selli Interval PS-2 Lower Aptian Lower Aptian PS-1 Leupoldina cabri C3 Marnes Bleues

-268 (NC6) 4

(NC6) C5 2

C2 C. litterarius Niveau Goguel Interval C4 C. litterarius 0 12345 section not δ13 -273 Maiolica Scaglia Variegata Ccarb accessible

Globigerinelloides blowi lamination "black shale"

marl C1 -278 marly limestone 12345 limestone δ13Ccarb Fig. 2. Chemo- and biostratigraphic correlation of the OAE 1a black shale interval from the Serre Chaitieu section (Vocontian Basin, SE France) and the Cismon section (Belluno Basin, N Italy). Note that the differences in thickness of the individual chemostratigraphic segments corresponds well with the inferred sedimentation rates for the two different depositional settings. δ13C of bulk carbonate carbon reported in per mil versus VPDB. Chemostratigraphy and lithology of the Cismon section after Menegatti et al. (1998), biostratigraphy after Erba et al. (1999). Biostratigraphy of the Serre Chaitieu section after Herrle and Mutterlose (2003). Labels C1 to C7 indicate chemostratigraphic segments [14]. Dark grey bars refer to black shale horizons. Solid lines indicate the chemostratigraphic correlation, stippled line reflects the boundary between NC6 and NC7. PS-1 to PS-6 represent individual laminated black shale horizons in the Serre Chaitieu section.

Furthermore, high quantities of phytane point to a phytoplanktonic source (Didyk et al., 1978; Kohnen et al., 1992). Bacterial contributions are recorded in the high abundance of hopanes (Rohmer et al., 1992). There is no evidence for an important cyanobacterial contribution in the studied interval. The low quantities of long-chain n-alkanes in the sediments indicate only minor inputs of continent-derived vascular plant waxes (Eglinton and Hamilton, 1967).

Chapter 2 27

Within the OAE 1a interval the biomarker distribution shows no significant variation between the laminated facies and the bioturbated marls whereas in the upper part of the section, the decreasing TOC values are paralleled by a continuous decline in hopane and sterane abundances. These geochemical results are supported by optical studies of the amorphous organic matter (AOM). AOM forms the main constituent of the bulk kerogen, up to ~ 95 % within the OAE 1a and ~ 70 % to 90 % in the interval above. Two different types of AOM can be distinguished. Type A is composed of glossy, inclusion-rich, orange-brown floccules with moderate to strong fluorescence and dominates within the OAE 1a interval and within the laminated black shales. Type B has a matt, grey to grey-brown appearance with weak to moderate fluorescence and represents the major constituent in the upper part of the section. According to different authors (Tyson, 1995 and references therein) fluorescent AOM is considered to be derived from phytoplankton and/or bacteria and their decompositional products and dominates in dysoxic-anoxic environments. Although degraded terrestrial material can have a similar appearance to marine-derived AOM (Gorin and Feist-Burkhardt, 1990), in the Serre Chaitieu section, the absence of any woody or cuticular structures and the present fluorescence clearly suggests a marine origin for both AOM types. In summary, the results of extractable hydrocarbon analysis and optical AOM studies consistently indicate a marine phytoplankton and/or bacterial origin for most of the OM in the studied section. A similar, predominantly marine OM composition with only minor terrestrial contribution has been reported by Bréhéret (1994) for the same section based on Rock-Eval data and by Baudin et al. (1998) for the time-equivalent Livello Selli interval in the Umbria- Marche Basin (Italy).

6.3. Organic carbon isotope geochemistry

13 The carbon isotopic composition of the bulk OM displays a significant shift in δ Corg from mean values of ~ -25.5 ‰ in the lower part to values of ~ -23.8 ‰ prevailing in the upper part 13 of the section (Fig. 3). The increase in δ Corg towards higher values has several superimposed 13 smaller-scale fluctuations (up to ~ 0.5 ‰). In comparison to the δ Ccarb record, the bulk 13 δ Corg record shows a similar positive excursion of ~1.7 ‰ with a stepwise shift towards higher values in the C6 segment.

Chapter 2 28

In order to minimize secondary processes affecting the δ13C signature of bulk OM, the isotopic composition of biomarkers derived predominantly from marine primary producers was determined (Hayes et al., 1989; Kuypers et al., 2002; Sinninghe Damsté et al., 1998).

Short-chain n-alkanes (n-C17, n-C18) are interpreted to derive from algal precursor compounds 13 (Gelpi et al., 1970) and display a similar carbon isotopic shift as the bulk δ Corg signal, although the n-alkanes are depleted by ~ 5.0 ‰. Within the OAE 1a interval the carbon isotopic composition of C28 steranes (24-methyl-5α-cholestane) parallels the short-chain n- alkane record almost perfectly. This sterane derives from C28 sterol, a compound which is biosynthesized predominantly by marine algae (Volkman, 1986). The congruence in δ13C of

C28 steranes and short-chain n-alkanes strongly supports the interpretation of an algal origin 13 for the latter. Due to the low abundance of C28 steranes in the upper part, δ C values could not be determined. The intermediate n-alkanes (n-C23, n-C24) cannot be assigned to a specific 13 13 marine source but again parallel the bulk δ Corg pattern with a depletion in C of ~ 4.0 ‰. 13 In summary, the δ C composition of biomarkers (short-chain n-alkanes, C28 sterane) and bulk OM reveal a similar pattern during and after deposition of the OAE 1a interval characterized by a stepwise shift towards higher δ13C values. Individual biomarkers show more pronounced small-scale carbon isotope fluctuations and a stronger all-over shift in δ13C than bulk OM. In 13 comparison to δ Ccarb, the biomarker record displays an increased overall shift of ~ 2.5 to 3.0 13 ‰ (δ Ccarb = ~ 1.5 ‰) within the C6 segment.

6.4. Estimation of pCO2 change in the course of OAE 1a formation

In general, positive carbon isotope excursions have been interpreted in terms of increased organic carbon burial, resulting from preferential removal of 12C into the sediments and the accompanying enrichment of 13C in the oceanic DIC reservoir (Arthur et al., 1985; Scholle and Arthur, 1980). Intense OC burial is expected to result in a lowering of oceanic [CO2 (aq)] and consequently in a reduction of atmospheric pCO2 (Arthur et al., 1988; Freeman and Hayes, 1992). The carbon isotopic composition of inorganic carbon and primary organic carbon can be used to estimate changes in ancient pCO2 and/or in palaeoproductivity (Andersen et al., 1999; Freeman and Hayes, 1992; Hayes et al., 1989; Joachimski et al., 2002; Pagani et al., 1999). The δ13C composition of marine primary organic matter is determined by the isotopic 13 composition of the carbon source (δ C of oceanic dissolved CO2) and by the photosynthetic

Chapter 2 29

fractionation factor (εp) of the carbon-consuming primary producers. The isotopic fractionation is in turn a function of the concentration of dissolved CO2 ([CO2 (aq)]) (Freeman and Hayes, 1992; Rau and Takahashi, 1989) as well as of various physiological factors including growth rate and cell geometry (Bidigare et al., 1997; Popp et al., 1998). A decrease in atmospheric pCO2 is expected to be paralled by a decrease in oceanic [CO2 (aq)], leading to 13 a reduction in εp values. This εp decrease should result in a shift towards higher δ C values of primary organic carbon due to reduced photosynthetic fractionation. However, a similar signal is expected to result from an increase in palaeoproductivity, being accompanied by increasing growth rates. 13 The ~1.5 ‰ positive shift in δ Ccarb represents a well documented and characteristic feature of the Aptian isotope curve. It has been reproduced from many sites independent of facies or latitudinal variations (Erbacher and Thurow, 1997; Menegatti et al., 1998; Strasser et al., 13 2001). Hence, the δ Ccarb pattern measured in the Vocontian Basin is interpreted to reflect ocean-wide variations in the carbon isotopic composition of the oceanic DIC reservoir and can be used to determine changes in εp values.

To estimate palaeoatmospheric CO2 concentrations, the fractionation factor (εp) of marine photosynthetic plankton needs to be determined. Therefore, the carbon isotopic compositions of dissolved oceanic CO2 and of primary photosynthate have to be assessed. Assuming ambient sea surface temperatures for the Early Aptian Vocontian Basin between 20°C and 13 30°C, the isotopic composition of oceanic dissolved CO2 can be calculated from δ Ccarb Based on the temperature-dependent fractionation factor of Romanek et al. (1992), the δ13C of dissolved CO2 lies in the range of -6.6 ‰ (20°C) to -5.4 ‰ (30°C) before and -5.1 ‰ (20°C) to -3.9 ‰ (30°C) during the positive isotope excursion. C28 steroids and short-chain n-alkanes are assumed to show an average depletion of ~ 4 ‰ compared to primary biomass resulting in values of -26.0 ‰ (pre-excursion, onset of segment C6) and -23.5 ‰ (excursion, end of segment C6) for the latter. Following the method of Freeman and Hayes (1992) the calculated

εp values have been converted into surface water [CO2 (aq)] based on the empirical relationship: log [CO2 (aq)] = 0.0551 * εp + 0.305. Finally, atmospheric CO2 concentrations where calculated by applying Henry’s Law.

Chapter 2 30 OAE 1a Interval 1a OAE p ε 15.0 19.0 23.0 -26.0 18 n-C 17 n-C n-alk (short-chain) C 13 δ -32.0 -30.0 -28.0 23 -26.0 n-C 24 n-C steroid + n-alk. C steroid

13 28 δ C -32.0 -30.0 -28.0 5 C7 C6 OM carb C5 C bulk 13 C C4 δ 13 δ 1234 -27.0 -25.0 -23.0 1.0 2.0 3.0 TOC (wt %) TOC

0.0

Lithology Meters 8 4 2 0 6 12 10

Fig. 3. Stratigraphy, lithology, total organic carbon (TOC) content, δ13C of bulk carbonate and OM,

13 δ C of individual biomarkers and calculated εp values across the OAE 1a interval, Serre Chaitieu section (Vocontian Basin, SE France). Carbon isotope values are reported in per mil versus VPDB. Biostratigraphy after Herrle and Mutterlose (2003). Dark grey bars refer to black shale horizons, pale grey area corresponds to the OAE 1a interval. Labels C4 to C7 indicate chemostratigraphic segments.

Chapter 2 31

The resulting estimates of palaeoatmospheric CO2 concentrations indicate that the Early

Aptian pCO2 level was about 3 to 4 times the pre-industrial level (~ 280 ppm). This result corresponds well to estimates of Freeman and Hayes (1992) for the mid-Cretaceous (~ 900 to 1200 ppm) based on the same method. Furthermore our data are in broad agreement with stomatal densitiy-derived pCO2 estimates (Beerling and Royer, 2002) as well as with results of geochemical modelling (Berner, 1994). Calculation of the pCO2 drop following the OAE 1a event results in a decrease of ~ 100-130 ppm or 10-15 % respectively. Based on sedimentary porphyrins from the Greenhorn Formation (USA) a comparable decrease in εp values of 1.5 ‰ has been calculated by Hayes et al. (1989) during the Cenomanian-Turonian black shale event (CTBE, OAE 2) and interpreted to reflect a ~ 20 % reduction in atmospheric pCO2 (Freeman and Hayes, 1992).

Estimated pCO2 variations for the Vocontian Basin are based on the assumption, that no physiological and environmental variables other than pCO2 affected εp values of marine phototrophs. This is in contrast to several authors (Hochuli et al., 1999; Kuypers et al., 2002; Pedersen and Calvert, 1990), who emphasize the important role of enhanced palaeoproductivity during formation of the mid-Cretaceous OAE black shales. This indicates that at least some portion of the estimated εp change in the Vocontian Basin might have been caused by increased algal growth rates and average cell sizes due to higher palaeofertility.

Consequently the estimated pCO2 decrease represents a maximum value.

6.5. Palynology

In order to trace changes in terrestrial climate patterns during and after formation of the OAE 1a interval, the relative and absolute abundance of climate-sensitive spore and pollen groups (incl. pteridophyte spores, bisaccate pollen and Classopollis spp.) have been analysed. Additionally, the composition of the entire palynofloral assemblage has been determined qualitatively. The current palynological findings are compared with the results of Hochuli et al. (1999) from the Cismon section. The cited percentages (%) refer to the total sporomorph counts.

6.5.1. Results of Palynological Analysis

In the Serre Chaitieu section the sporomorph assemblage accounts for only ~ 10 to 15 % of the particulate organic matter (excluding AOM). Pteridophyte spores represent an important

Chapter 2 32 constituent of the sporomorph assemblage and account for 20.6 to 41.5 % (mean 30.1). Classopollis spp. shows an increase from 25.9 to 43.8 % in the lower part followed by a subsequent decline in the upper part of the succession. Bisaccate pollen account for less than 12.9 % in most samples (max. 20.0 %) and display only minor variations throughout the section. The high abundance of pteridophyte spores – essentially represented by Deltoidospora spp. and Gleicheniidites spp. as well as the common occurrence of Classopollis-type pollen and the low percentage of bisaccates reflect a position in the southern part of the Southern Laurasian floral province. Some minor influence of the Northern Gondwana floral province is reflected by the rare, but consistent occurrences of Afropollis spp. and Ephedripites spp. Other typical elements of this province like Tucanopollis crisopolensis are scarce or absent.

6.5.2. Comparison of the palynological records from SE France and N Italy

Based on the chemostratigraphic correlation scheme, the sporomorph findings from the Serre Chaitieu section can be compared in detail with the palynological record of Hochuli et al. (1999) from the Cismon section (Fig. 4). The most distinct features are: (1) relatively high percentages of spores in the Serre Chaitieu section (mean of 30.1 % of total sporomorphs) compared to the Cismon section (mean of 7.1 % of total sporomorphs), (2) a prominent post- black shale increase in bisaccate pollen (from mean values of 18.6 % within to values of 72.2 % above the OAE 1a interval), accompanied by a decrease in Classopollis spp. in the Cismon section, (3) a rather uniform stratigraphic pattern of bisaccate pollen, Classopollis spp. and pteridophyte spores in the Serre Chaitieu section. Besides climatically-driven variations, hydrodynamic sorting processes can cause significant changes of the palynological assemblage in distal depositional settings. According to Tyson (Tyson, 1995), thick-walled spores are in general deposited near-shore in the vicinity of river mouths. The absence of large, thick-walled spores (e.g. Foveosporites spp, Impardecispora spp.) in the Serre Chaitieu assemblage suggests that some fractionation has already occurred. However, the percentage of pteridophyte spores is still relatively high considering the hemipelagic depositional environment of the Serre Chaitieu section (Tyson, 1995). This is interpreted to reflect the relatively high terrigenous flux to the basin (Bréhéret, 1994) and/or the enclosed nature of the Vocontian Basin. In contrast, the low amount of spores in the

Chapter 2 33

Cismon section (Hochuli et al. 1999)

% total sporomorphs

pteridophyte Classopollis bisaccate

meters Stage Foram.-zone Nanno.-zone Lithology spores pollen

Serre Chaitieu section (this study)

% total sporomorphs (NC7) Lithology Nanno.-zone Foram.-zone Stage meters

pteridophyte Classopollis 12 -258 bisaccate spores pollen

10 R. angustus (NC7A)

8 R. angustus Leupoldina cabri -263

Lower Aptian 6 Lower Aptian Leupoldina cabri

(NC6) 4 (NC6) Livello Selli Interval OAE 1a 050100050100 0 50 100 2 -268 C. litterarius C. litterarius Interval of selective sporomorph preservation 050100050100 050100 0 section not G. blowi accessible

Fig. 1 Correlation of frequency patterns (in percentage of total sporomorph counts) of climate- sensitive spores and pollen of the Serre Chaitieu and the Cismon record. Note the difference in bisaccate pollen abundance between the two records. Biostratigraphy of the Cismon section after Erba et al. (1999), lithology after Menegatti et al. (1998), pollen abundance after Hochuli et al. (1999) and Hochuli (unpubl. results). Biostratigraphy of the Serre Chaitieu section after Herrle and Mutterlose (2003). Solid lines indicate the chemostratigraphic correlation, stippled line reflects the boundary between NC6 and NC7. For lithological explanations see Fig. 3. pelagic sediments of the Belluno Trough (Cismon section) reflects the great distance of the depositional setting to continental areas. According to Vakhrameyev (1978; 1982) the abundance of Classopollis-producing cheirolepidacean plants display a climate-controlled increase towards low latitudes. High contents of Classopollis spp. (> 50 %) have been interpreted to indicate warm and arid climates. In contrast, bisaccates of pinaceous affinity represent a typical floral element of the boreal realm and therefore point to comparatively cool and humid climates (Batten, 1984; Brenner, 1976; Hochuli et al., 1999). Besides climatic effects, bisaccate pollen are strongly affected by selective sorting processes during transportation and deposition (Traverse, 1988; Tyson, 1995). The low amount and relatively

Chapter 2 34 stable distribution pattern of bisaccate pollen in the Serre Chaitieu section is in strong contrast to the observations from the Cismon record where a rapid and significant increase in bisaccate pollen (up to 86.3 % above the OAE 1a interval) is accompanied by a decrease in Classopollis-type pollen (1999). This strong increase in boreal floral elements has been attributed to a major cooling episode and/or a major reorganisation in the oceanographic circulation system of the W Tethys in the aftermath of the OAE 1a. Compared to the Cismon section, the composition of the observed spore-pollen association in the Serre Chaitieu section remains essentially unchanged across OAE 1a and the dominant forms persist throughout the studied record. This indicates that the continental hinterland of the Vocontian Basin was characterized by relatively stable vegetational patterns and associated palaeoenvironments. No signs of major climatic or oceanographic disturbances can be observed within the corresponding time-interval in the Vocontian Basin.

7. Integration of the chemostratigraphic, geochemical and palynological results

Based on geochemical evidence (Brass et al., 1982) and ocean general circulation model experiments (Barron and Peterson, 1990; Bice et al., 1997), deep water circulation in the mid- Cretaceous ocean was predominantly controlled by the formation of warm and saline waters in low latitude shelf areas. These unusual palaeoceanographic conditions favoured the formation of thinly laminated, OC-rich black shales in hemipelagic and pelagic environments, which reflect deposition under dysoxic-anoxic bottom water conditions. Short-lived periods of euxinia reaching the photic zone during the OAE 1a interval have been reported by Van Breugel et al. (2002). The occurrence of episodic oxygen deficiency in oceanic bottom waters has been interpreted to reflect periods of pronounced water column stratification (Erbacher et al., 2001) and/or a decrease in the rate of deep-water formation (Bralower and Thierstein, 1984). In addition, oxygen depletion and resulting anoxia due to enhanced productivity in ocean surface waters has been invoked as a possible mechanism for the formation of mid- Cretaceous black shales (Kuypers et al., 2002; Pedersen and Calvert, 1990). The detailed chemostratigraphic correlation scheme presented in this study demonstrates clearly that the occurrence of laminated horizons in the OAE 1a interval (PS-1 to PS-4) are equivalent to individual black shale layers in the Cismon section (-263.5 to -262.5 m). The coeval deposition of discrete black shale horizons at both locations indicates that comparatively short episodes of oxygen deficiency in bottom waters prevailed over large areas in the W Tethys Ocean independent of the depositional setting. In contrast to this, the

Chapter 2 35

Cismon record holds no equivalent to the upper black shale horizons (PS-5 to PS-6) occurring in the Serre Chaitieu section. This leads to the conclusion, that dysoxic-anoxic bottom water conditions episodically reoccurred in the Vocontian Basin in the aftermath of OAE 1a, whereas the deposition of carbonate-rich sediments in the Cismon section above OAE 1a points to a rapid reestablishment towards normal-marine conditions in the Belluno Trough. The palynological analysis across the Serre Chaitieu OAE 1a interval shows that neither the distribution of climate-sensitive pollen forms nor the qualitative composition of the palynofloral assemblage exhibit significant changes. These findings are supported by palynological results from time-equivalent, hemipelagic sediments from the Roter Sattel section (Prealps, Switzerland) where no prominent variations have been identified in the bisaccate pollen spectrum during or after formation of the OAE 1a black shale (Hochuli, unpubl. results). Furthermore, this in accordance with palynological data from shallow-water deposits from the Luz section (Algarve Basin, S Portugal) where bisaccate pollen account for less than 15 % during the late Early Aptian interval (Heimhofer et al., in prep.). These results indicate that the strong boreal pollen signal observed in the Cismon record is restricted to the pleagic deposits in the SW Tethys whereas no prominent changes are visible in the sections along the N Tethys and E Atlantic margins. Compared to the palaeolatitudinal position of the Cismon site, the Vocontian Basin was located ~ 8° to 10° more to the north. Hence, a southward shift of the Laurasian floral provinces due to climate cooling is expected to result in a significant increase in boreal pollen types along the N Tethys margin. However, neither in the Serre Chaitieu nor in the Roter Sattel assemblage a distinct trend towards a dominance of boreal sporomorphs can be observed. The stable spore-pollen distribution pattern across the

Serre Chaitieu OAE 1a interval is supported by the isotopically derived pCO2 estimates which point to a moderate decrease in atmospheric CO2 concentration of < 10 % – 15 %. This decrease is regarded to be insufficient to cause a severe global cooling, resulting in a major southward shift of the Laurasian floral provinces. In order to explain the discrepancy in the pollen spectra at the different locations we propose an alternative scenario. In offshore marine settings, continental runoff and marine currents are the main controlling factors for the spatial distribution of sporomorphs (Tyson, 1995). Bisaccate pollen have the capability to float for a relatively long time, which explains their relative increase in abundance on the shelf with increasing distance from the shoreline (Heusser and Balsam, 1977). According to Melia (1984) bisaccate conifer pollen can be transported over long distances and represent a rare but persistent constituent of deep-sea

Chapter 2 36 sediments. Consequently, the observed differences in the spore-pollen patterns might reflect changes in the oceanic current patterns and/or sea level rather than vast latitudinal shifts of the major floral belts. As mentioned above, the occurrence of dysoxic-anoxic bottom water conditions during formation of OAE 1a is probably linked to pronounced thermohaline stratification and decelerated renewal of relatively warm and saline deep waters. The abrupt termination of black shale deposition in the SW Tethys basin indicates a major reorganisation of the oceanic circulation patterns in this region. These paleoceanographic perturbations are accompanied by a major global sea-level rise during the Early Aptian (Hardenbol et al., 1998; Strasser et al., 2001). According to several authors, the formation of the OAE 1a itself is directly linked to this major transgression (Bréhéret, 1994; Erbacher and Thurow, 1997). The flooding of broad continental areas during sea-level rise resulted in the opening of gateways and deepening of existing connections between the Tethys and the adjacent ocean basins. The existence of N-S trending seaways which connected the W Tethys and the boreal oceans e.g. via the Polish Trough and the Moscow Platform during the mid-Cretaceous has been documented by palaeontological and palaeogeographic means (e.g. Marcinowski and Wiedmann, 1988). Hence, the high percentages of bisaccates in the distal facies of the Cismon record might reflect changes in the paleoceanographic current system (e.g. inflow of boreal water masses) and/or the effect of selective sorting due to a concomitant sea level rise.

8. Conclusions

The combination of organic geochemistry, carbon isotope analysis and palynology provides a valuable tool to study past changes of terrestrial environments during times of major oceanic perturbations. Our results indicate that during the Early Aptian OAE 1a the oceanic realm and its ecosystems were much more affected by severe disturbances than continental environments. Variations in palaeoatmospheric CO2 concentrations seem to be of minor importance. The most important findings of our study include the following conclusions.

(1) Carbon isotope records measured on different substrates (Ccarb, Corg, Cn-alk) show a similar positive excursion starting with the end of OAE 1a in the Vocontian Basin. The δ13C curve and its individual segments can be accurately correlated on a high-resolution with the existing Cismon record (SW Tethys). The chemostratigraphic correlation indicates the occurrence of short-term episodes of bottom water anoxia throughout the W Tethys

Chapter 2 37 independent of depositional setting. The abrupt termination of OAE 1a in the Cismon record contrasts with a more gradual reestablishment of normal-marine conditions in the Vocontian Basin.

(2) Estimations of palaeoatmospheric CO2 concentrations during the Early Aptian are in accordance with results from other palaeobarometeric proxies and yield pCO2 levels of about

3 to 4 times the preindustrial level. Calculated changes in pCO2 across the OAE 1a are only moderate and can not account for a major cooling accompanied by a southward shift of boreal floras in the western Tethys region. (3) The abundance patterns of climate-sensitive spores and pollen in combination with the observed palynofloral association reveal relatively stable patterns in vegetation and associated palaeoenvironments during times of black shale accumulation in the adjacent basin. Evidence for major climatic disturbances accompanied by prominent shifts of the floral belts is missing in the Vocontian Basin. (4) The contrasting pollen records of the Cismon and Serre Chaitieu sections are interpreted to reflect the reorganisation of oceanic circulation patterns in the aftermath of the OAE 1a and/or the effect of selective sorting of the palynological assemblages in the pelagic Cismon section due to a late Early Aptian sea level rise.

Acknowledgements

We thank Luc Zwank from the EAWAG for support with the irmGC-MS measurements and Christian Ostertag-Henning from the University of Münster for help with the sterane identification. This manuscript was significantly improved thanks to suggestions and reviews by R. V. Tyson and M. Pagani. Financial support from ETH-project TH-34./99-4 is greatfully acknowledged.

Chapter 2 38

References

Andersen, N., Müller, P.J., Kirst, G. and Schneider, R.R., 1999. Alkenone δ13C as a proxy for past PCO2 in surface waters: results from the Late Quaternary Angola Current. In: G. Fischer and G. Wefer (Editors), Use of proxies in paleoceanography: examples from the South Atlantic. Springer, Berlin/Heidelberg, pp. 469-488. Arnaud, H. and Lemoine, M., 1993. Structure and Mesozoic-Cenozoic evolution of the South-East France Basin (SFB). Géologie Alpine, Série spéciale "Colloques et Excursions", 3, 3-58. Arthur, M.A., Dean, W.E. and Pratt, L.M., 1988. Geochemical and climatic effects of increased marine organic carbon burial at the Cenomanian/Turonian boundary. Nature, 335, 714-717. Arthur, M.A., Dean, W.E. and Schlanger, S.O., 1985. Variations in the global carbon cycle during the Cretaceous related to climate, volcanism and changes in atmospheric CO2. In: E.T. Sundquist and W.S. Broecker (Editors), The carbon cycle and atmospheric CO2: natural variations Archean to present. AGU Geophys. Monogr., pp. 504-529. Barron, E.J. and Peterson, W.H., 1990. Mid-Cretaceous ocean circulation: results from model sensitivity studies. Paleoceanography, 5, 319-337. Batten, D.J., 1984. Palynology, climate and the development of Late Cretaceous floral provinces in the Northern Hemisphere: a review. In: P. Brenchley (Editor), Fossils and climate. John Wiley and Sons, New York, pp. 127-164. Baudin, F., Fiet, N., Coccioni, R. and Galeotti, S., 1998. Organic matter characterisation of the Selli Level (Umbria-Marche Basin, central Italy). Cretaceous Research, 19, 701-714. Beerling, D.J., Lomas, M.R. and Gröcke, D.R., 2002. On the nature of methane gas-hydrate dissociation during the Toarcian and Aptian oceanic anoxic events. American Journal of Science, 302, 28-49. Beerling, D.J. and Royer, D.L., 2002. Fossil plants as indicators of the Phanerozoic global carbon cycle. Annual Review of Earth and Planetary Sciences, 30, 527-556. Berner, R.A., 1994. GEOCARB II: a revised model of atmospheric CO2 over Phanerozoic time. American Journal of Science, 294, 56-91. Bice, K.L., Barron, E.J. and Peterson, W.H., 1997. Continental runoff and early Cenozoic bottom- water sources. Geology, 25, 951-954. Bidigare, R.R., Fluegge, A., Freeman, K.H., Hanson, K.L., Hayes, J.M., Hollander, D., Jasper, J.P., King, L.L., Laws, E.A., Milder, J., Millero, F.J., Pancost, R., Popp, B.N., Steinberg, P.A. and Wakeham, S.G., 1997. Consistent fractionation of 13C in nature and in the laboratory: growth- rate effects in some haptophyte algae. Global Biogeochemical Cycles, 11, 279-292. Bralower, T.J., CoBabe, E., Clement, B., Sliter, W.V., Osburn, C.L. and Longoria, J., 1999. The record of global change in mid-Cretaceous (Barremian-Albian) sections from the Sierra Madre, northeastern Mexico. Journal of Foraminiferal Research, 29, 418-437. Bralower, T.J. and Thierstein, H.R., 1984. Low productivity and slow deep-water circulation in mid- Cretaceous oceans. Geology, 12, 614-618. Brass, G.W., Southam, J.R. and Peterson, W.H., 1982. Warm saline bottom water in the ancient ocean. Nature, 296, 620-623. Bréhéret, J.G., 1988. Episodes de sédimentation riche en matière organique dans le marnes bleues d'age aptien et albien de la partie pélagique du bassin Vocontien. Bulletin de la Societe Geologique de France, 4, 349-356. Bréhéret, J.G., 1994. The mid-Cretaceous organic-rich sediments from the Vocontian Zone of the French southeast basin. In: A. Mascle (Editor), Hydrocarbon and petroleum geology of France. Springer, Berlin/Heidelberg, pp. 295-220. Bréhéret, J.G., 1997. L'Aptian et l'Albian de la fosse vocontienne (bordures au bassin): évolution de la sédimentation et enseignements sur les événements anoxiques. Publications de la Société Géologique du Nord, 25, 1-614. Brenner, G., 1976. Middle Cretaceous floral provinces and early migration of Angiosperms. In: C.B. Beck (Editor), Origin and early evolution of angiosperms. Columbia University Press, New York, pp. 23-44.

Chapter 2 39

Didyk, B.M., Simoneit, B.R.T., Brassell, S.C. and Eglinton, G., 1978. Organic geochemical indicators of palaeoenvironmental conditions of sedimentation. Nature, 272, 216-222. Eglinton, G. and Hamilton, C.P., 1967. Leaf epicuticular waxes. Science, 156, 1322-1335. Erba, E., 1994. Nannofossils and superplumes: the Early Aptian "nannoconid crisis". Paleoceanography, 9, 483-501. Erba, E., Channell, J.E.T., Claps, M., Jones, C., Larson, R., Opdyke, B., Premoli Silva, I., Riva, A., Salvini, G. and Torricelli, S., 1999. Integrated stratigraphy of the Cismon Apticore (southern Alps, Italy): a "reference section" for the Barremian-Aptian interval at low latitudes. Journal of Foraminiferal Research, 29, 371-391. Erbacher, J., Huber, B.T., Norris, R.D. and Markey, M., 2001. Increased thermohaline stratification as a possible cause for an ocean anoxic event in the Cretaceous period. Nature, 409, 325-327. Erbacher, J. and Thurow, J., 1997. Influence of oceanic anoxic events on the evolution of mid- Cretaceous radiolaria in the North Atlantic and western Tethys. Marine Micropaleontology, 30, 139-158. Farrimond, P., Eglinton, G., Brassell, S.C. and Jenkyns, H.C., 1990. The Cenomanian-Turonian anoxic event in Europe: an organic geochemical study. Marine and Petroleum Geology, 7, 75-89. Francis, J.E. and Frakes, L.A., 1993. Cretaceous climates. In: V.P. Wright (Editor). Sed. Rev. Blackwell, Oxford, pp. 17-30. Freeman, K.H. and Hayes, J.M., 1992. Fractionation of carbon isotopes by phytoplankton and estimates of ancient CO2 levels. Global Biogeochemical Cycles, 6, 185-198. Gelpi, E., Schneider, H., Mann, J. and Oro, J., 1970. Hydrocarbons of geochemical significance in microscopic algae. Phytochemistry, 9, 603-612. Gorin, G.E. and Feist-Burkhardt, S., 1990. Organic facies of early to middle Jurassic sediments in the Jura Mountains, Switzerland. Review of Palaeobotany and Palynology, 65, 349-355. Gröcke, D.R., Hesselbo, S.P. and Jenkyns, H.C., 1999. Carbon-isotope composition of Lower Cretaceous fossil wood: ocean-atmosphere chemistry and relation to sea-level change. Geology, 27, 155-158. Hallam, A., 1985. A review of Mesozoic climate. Journal of the Geological Society of London, 142, 433-445. Hardenbol, J., Thierry, J., Farley, M.B., Jacquin, T., de Graciansky, P.-C. and Vail, P.R., 1998. Cretaceous sequence chronostratigraphy. In: P.-C. de Graciansky, J. Hardenbol, T. Jacquin and P.R. Vail (Editors), Mesozoic and Cenozoic sequence stratigraphy of European basins. SEPM Spec. Publ. Hay, W.W., DeConto, R.M., Wold, C.N., Wilson, K.M., Voigt, S., Schulz, M., Rossby Wold, A., Dullo, W.C., Ronov, A.B., Balukhovsky, A.N. and Söding, E., 1999. Alternative global Cretaceous paleogeography. In: E. Barrera and C.C. Johnson (Editors), Evolution of the Cretaceous ocean-climate system. Geol. Soc. Am. Spec. Paper, pp. 1-47. Hayes, J.M., Popp, B.N., Takigiku, R. and Johnson, M.W., 1989. An isotopic study of biogeochemical relationships between carbonates and organic carbon in the Greenhorn Formation. Geochimica et Cosmochimica Acta, 53, 2961-2972. Heimhofer, U., Hochuli, P.A., Burla, S., Andersen, N. and Weissert, H., 2003. Terrestrial carbon- isotope records from coastal deposits (Algarve, Portugal): a tool for chemostratigraphic correlation on an intrabasinal and global scale. Terra Nova, 15, 8-13. Herbert, T.D., 1992. Paleomagnetic calibration of Milankovitch cyclicity in Lower Cretaceous sediments. Earth and Planetary Science Letters, 112, 15-28. Herrle, J.O., Kößler, P., Friedrich, O., Erlenkeuser, H. and Hemleben, C., 2004. High-resolution carbon isotope records of the Aptian to Lower Albian from SE France and the Mazagan Plateau (DSPD site 545): a stratigraphic tool for paleocanographic and paleobiologic reconstruction. Earth and Planetary Science Letters, 218, 149-161. Herrle, J.O. and Mutterlose, J., 2003. Calcareous nannofossils from the Aptian-Lower Albian of southeast France: palaeoecological and biostratigraphic implications. Cretaceous Research, 24, 1-22.

Chapter 2 40

Herrle, J.O., Pross, J., Friedrich, O. and Hemleben, C., 2003. Short-term environmental changes in the Cretaceous Tethyan Ocean: micropalaeontological evidence from the Early Albian Oceanic Anoxic Event 1b. Terra Nova, 15, 14-19. Heusser, L. and Balsam, W.L., 1977. Pollen distribution in the Northeast Pacific Ocean. Quaternary Research, 7, 45-62. Hochuli, P.A., Menegatti, A.P., Weissert, H., Riva, A., Erba, E. and Premoli Silva, I., 1999. Episodes of high productivity and cooling in the Early Aptian Alpine Tethys. Geology, 27, 657-660. Huber, B.T., Hodell, D.A. and Hamilton, C.P., 1995. Middle-Late Cretaceous climate of the southern high latitudes: stable isotopic evidence for minimal equator-to-pole thermal gradients. Geological Society of America Bulletin, 107, 1164-1191. Joachimski, M.M., Pancost, R.D., Freeman, K.H., Ostertag-Henning, C. and Buggisch, W., 2002. Carbon isotope geochemistry of the Frasnian-Famennian transition. Palaeogeography, Palaeoclimatology, Palaeoecology, 181, 91-109. Kohnen, M.E.L., Schouten, S., Sinninghe Damsté, J.S., de Leeuw, J.W., Merritt, D.A. and Hayes, J.M., 1992. Recognition of paleobiochemicals by a combined molecular sulfur and isotope geochemical approach. Science, 256, 358-362. Kössler, P., Herrle, J.O., Appel, E., Erbacher, J. and Hemleben, C., 2001. Magnetic records of climatic cycles from mid-Cretaceous hemipelagic sediments of the Vocontian Basin, SE France. Cretaceous Research, 22, 321-331. Kuypers, M.M.M., Pancost, R.D., Nijenhuis, I.A. and Sinninghe Damsté, J.S., 2002. Enhanced productivity led to increased organic carbon burial in the euxinic North Atlantic basin during the late Cenomanian oceanic anoxic event. Paleoceanography, 17, 1-13. Kuypers, M.M.M., Pancost, R.D. and Sinninghe Damsté, J.S., 1999. A large and abrupt fall in atmospheric CO2 concentration during Cretaceous times. Nature, 399, 342-345. Larson, R.L. and Erba, E., 1999. Onset of the mid-Cretaceous greenhouse in the Barremian-Aptian igneous events and the biological, sedimentary and geochemical responses. Paleoceanography, 14, 663-678. Leckie, R.M., Bralower, T.J. and Cashman, R., 2002. Oceanic anoxic events and plankton evolution: biotic response to tectonic forcing during the mid-Cretaceous. Paleoceanography, 17, 1-29. Levert, J. and Ferry, S., 1988. Diagenèse argileuse complexe dans le Mésozoïque subalpin révélée par cartographie des proportions relatives d'argiles selon des niveaux isochrones. Bulletin de la Societe Geologique de France, 8, 1029-1038. Marcinowski, R. and Wiedmann, J., 1988. Paleogeographic implications of the Albian ammonite faunas of Poland. In: J. Wiedmann and J. Kullmann (Editors), Cephalopods-present and past. Schweizerbart'sche Verlagsbuchhandlung, Stuttgart, pp. 491-504. Melia, M.B., 1984. The distribution and relationship between palynomorphs in aerosols and deep-sea sediments off the coast of Northwest Africa. Marine Geology, 58, 345-371. Menegatti, A.P., Weissert, H., Brown, R.S., Tyson, R.V., Farrimond, P., Strasser, A. and Caron, M., 1998. High-resolution δ13C stratigraphy through the Early Aptian "Livello Selli" of the Alpine Tethys. Paleoceanography, 13, 530-545. Pagani, M., Arthur, M.A. and Freeman, K.H., 1999. Miocene evolution of atmospheric carbon dioxide. Paleoceanography, 14, 273-292. Pedersen, T.F. and Calvert, S.E., 1990. Anoxia vs. productivity: what controls the formation of organic-carbon-rich sediments and sedimentary rocks? American Association of Petroleum Geologists Bulletin, 74, 454-466. Peters, K.E. and Moldowan, J.M., 1993. The Biomarker Guide. Prentice-Hall, Englewood Cliffs, NJ, 363 pp. Popp, B.N., Laws, E.A., Bidigare, R.R., Dore, J.E., Hanson, K.L. and Wakeham, S.G., 1998. Effect of phytoplankton cell geometry on carbon isotopic fractionation. Geochimica et Cosmochimica Acta, 62, 69-77. 13 Rau, G.H. and Takahashi, T., 1989. Latitudinal variations in plankton δ C: implications for CO2 and productivity in past oceans. Nature, 341, 516-518. Rohmer, M., Bisseret, P. and Neunlist, S., 1992. The hopanoids, prokaryotic triterpenoids and precursors of ubiquitous molecular fossils. In: J.M. Moldowan, P. Bisseret and S. Neunlist

Chapter 2 41

(Editors), Biological markers in sediments and petroleum. Prentice-Hall, Englewood Cliffs, NJ, pp. 1-17. Romanek, C.S., Grossman, E.L. and Morse, J.W., 1992. Carbon isotopic fractionation in synthetic aragonite and calcite: effects of temperature and precipitation rate. Geochimica et Cosmochimica Acta, 56, 419-430. Schlanger, S.O. and Jenkyns, H.C., 1976. Cretaceous oceanic anoxic events: causes and consequences. Geologie en Mijnbouw, 55, 179-184. Scholle, P.A. and Arthur, M.A., 1980. Carbon isotope fluctuations in Cretaceous pelagic limestones: potential stratigraphic and petroleum exploration tool. American Association of Petroleum Geologists Bulletin, 64, 67-87. Sinninghe Damsté, J.S., Kok, M.D., Köster, J. and Schouten, S., 1998. Sulfurized carbohydrates: an important sedimentary sink for organic carbon? Earth and Planetary Science Letters, 164, 7- 13. Strasser, A., Caron, M. and Gjermeni, M., 2001. The Aptian, Albian and Cenomanian of Roter Sattel, Romandes Prealps, Switzerland: a high-resolution record of oceanographic changes. Cretaceous Research, 22, 173-199. Traverse, A., 1988. Paleopalynology. Unwin Hyman, Boston, 600 pp. Tyson, R.V., 1995. Sedimentary organic matter. Chapman & Hall, London, 615 pp. Vakhrameyev, V.A., 1978. The climates of the Northern Hemisphere in the Cretaceous in the light of paleobotanical data. Paleontological Journal, 12, 143-154. Vakhrameyev, V.A., 1982. Classopollis pollen as an indicator of Jurassic and Cretaceous climate. International Geology Review, 24, 1190-1196. Van Breugel, Y., Schouten, S., Erba, E., Price, G.D. and Sinninghe Damsté, J.S., 2002. The global Aptian Selli event: molecular examination of the negative carbon isotope spike, Organic- carbon burial climate change and ocean chemistry (Mesozoic-Paleogene). Geol. Soc. Conf. Abstracts, pp. 59. Volkman, J.K., 1986. A review of sterol markers for marine and terrigenous organic matter. Organic Geochemistry, 9, 83-99. Weissert, H., Lini, A., Föllmi, K.B. and Kuhn, O., 1998. Correlation of Early Cretaceous carbon isotope stratigraphy and platform drowning events: a possible link? Palaeogeography, Palaeoclimatology, Palaeoecology, 137, 189-203. Wilson, P.A. and Norris, R.D., 2001. Warm tropical ocean surface and global anoxia during the mid- Cretaceous period. Nature, 412, 425-428.

Chapter 2 42

Chapter 3 43

Chapter 3

Palynological and calcareous nannofossil records across the late Early Aptian OAE 1a:

Implications for palaeoclimate, palaeofertility and detrital input

Abstract

High resolution records of terrestrial and marine-derived palynomorphs, particulate organic matter (OM) and calcareous nannoplankton provide new insights into the palaeoclimatic and palaeoceanographic conditions during deposition of the late Early Aptian oceanic anoxic event (OAE) 1a in the Vocontian Basin. The analysed spore-pollen assemblages indicate a rich and diverse flora, dominated by various ferns (e.g. Gleicheniaceae, Schizaeaceae, Osmundaceae), different types of cycads, bennettitales as well as by several conifer families (incl. Araucariaceae, Cheirolepidaceae, Podocarpaceae). The observed vegetation patterns remain essentially stable and the dominant pollen and spore types persist throughout the studied interval. The dinoflagellate cyst assemblage and diversity patterns as well as the calcareous nannofossil-based nutrient index provide no evidence for significant changes in the palaeofertility conditions across the OAE 1a. Based on congruent fluctuations in absolute abundances of terrestrial sporomorphs and marine organic-walled plankton, sedimentation rates (SR) and organic carbon mass accumulation rates (OC MAR) have been estimated tentatively. SR show significant fluctuations ranging from ~2.5 cm ka-1 in bioturbated marls to ~0.5 cm ka-1 in laminated, OC-rich horizons. Estimated OC MAR fluctuate between 0.02 to 0.06 gC cm-2 ka-1 and exhibit no evidence for increased OC accumulation during deposition of the OAE 1a black shales. Our results provide no evidence for enhanced surface water productivity due to accelerated climate-controlled nutrient fluxes during times of black shale deposition as previously suggested. In contrast, the concomitant occurrence of reduced detrital input and oxygen-deficient bottom waters rather suggests that fluctuations in sea-level and/or changes in runoff played a key role for the formation of OC-rich deposits during the late Early Aptian.

Keywords: OAE 1a; Aptian; palynology; dinoflagellate cysts; calcareous nannofossils; palaeoproductivity; Vocontian Basin

Chapter 3 44

1. Introduction

1.1. Palaeoclimatic conditions during the mid-Cretaceous

The Aptian to Turonian interval (~120-90 Ma, Gradstein et al., 1995) has been described as a time of warm climates (Hallam, 1985; Wilson and Norris, 2001), low equator-to-pole thermal gradients (Barron, 1983; Huber et al., 1995) and considerably high levels of atmospheric carbon dioxide (Beerling and Royer, 2002; Berner, 1994; Freeman and Hayes, 1992). Exceptional climatic conditions are also reflected in the composition and distribution of the fossil floras. The occurrence of ferns, conifers and cycadophytes throughout low to high latitudes in combination with the absence of expanded polar ice sheets points to more equable and warmer climates during the mid-Cretaceous in comparison to the present-day situation (Francis and Frakes, 1993; Hallam, 1985; Spicer and Parrish, 1986). However, Cretaceous climates were far from stable. Geochemical (Stoll and Schrag, 1996; Weissert and Lini, 1991; Wilson and Norris, 2001), micropalaeontological (Erba, 1994; Herrle et al., 2003b) as well as sedimentological evidence (Frakes and Francis, 1988; Kemper, 1987) indicates pronounced changes in the thermal state of the Cretaceous oceans and the climates of continental interiors. The episodic occurrence of organic carbon-rich intervals during the Barremian to Turonian has been interpreted to reflect major perturbations of the ocean-atmosphere system, accompanied by severe changes of the existing climatic patterns (Arthur et al., 1988; Herrle et al., 2003b; Kuypers et al., 1999; Weissert et al., 1998). These relatively short-lived intervals (~ 50 to 500 ka) of organic carbon (OC) accumulation were confined to marine pelagic and hemipelagic environments and have been termed Oceanic Anoxic Events (OAEs) by Schlanger and Jenkyns (1976). The late Early Aptian OAE 1a represents the first globally distributed black shale event of the Cretaceous and is accompanied by dramatic turnovers in nannoplankton (nannoconid-crisis of Erba, 1994) as well as in calcareous (Leckie et al., 2002; Premoli Silva et al., 1999) and siliceous plankton (Erbacher et al., 1996). In addition, a phase of carbonate platform demise has been documented from the northern Tethyan margin and circum-Atlantic regions (Weissert et al., 1998; Wissler et al., 2003) which shortly predates the OAE 1a. Prominent changes in the global carbon cycle during and after times of OAE 1a formation are reflected in the 13C/12C ratio of organic and carbonate carbon. The resulting δ13C pattern is characteristic for the Early Aptian and has been documented worldwide from various

Chapter 3 45 depositional settings (Bralower et al., 1999; Herrle et al., 2004; Menegatti et al., 1998; Price, 2003). The processes leading to the formation of the mid-Cretaceous OC-rich black shales are still a matter of debate. A variety of different palaeoceanographic models have been proposed during the last decades, most of which can be assigned to one of the two contrasting hypotheses. (1) The productivity model is based on the observation, that enhanced fertility in ocean surface waters results in an increased flux of OM to the sea floor. This in turn causes increasing oxygen deficiency within the water column and hence, increased OM preservation under dys- to anoxic bottom waters. The importance of enhanced oceanic productivity for the formation of mid-Cretaceous black shales has been emphasised by Arthur et al. (1987), Petersen and Calvert (1990), Premoli Silva et al. (1999) and Kuypers et al. (2002) among others. (2) In contrast to this, the stagnant ocean model argues with a reduction of deep-water renewal and/or the formation of thermohaline stratification. The decline in oxygen-rich deep water production prevents the aerobic degradation of organic matter within the water column and at the sediment-water interface, resulting in the accumulation of OC at the sea floor (e.g. Arthur et al., 1990; Bralower and Thierstein, 1984; Tyson, 1995).

1.2. Main objectives of the study

In this study, we present a detailed palynological record, encompassing the late Early Aptian OAE 1a interval in the Vocontian Basin (SE France). The spore-pollen record is combined with data on calcareous nannofossils, organic-walled plankton, particulate organic matter and with geochemical results. The main objectives of our study are: (i) to trace climate-induced variations in the terrestrial palynofloral assemblage and concomitant changes in the marine plankton associations during times of black shale deposition in the OAE 1a interval and (ii) to provide new insights from independent terrestrial and marine proxies on the controlling mechanisms for the formation of OC-rich deposits on regional and global scales. The chosen Serre Chaitieu section from the Vocontian Basin is particularly suitable to study changes in terrestrial vegetation patterns during times of widespread black shale formation. (i) Deposition in the Vocontian Basin was characterized by relatively high sedimentation rates due to prominent detrital fluxes from the adjacent continents. As a consequence, the occurring OM is well preserved and comprises relatively abundant terrestrial spores and pollen. (ii) Due to its favourable palaeobiogeographic position near the southern boundary of the Southern Laurasian floral province, the Vocontian Basin was sensitive to climate-induced shifts in the

Chapter 3 46 major floral belts. (iii) The succession has been studied in detail from bio- and chemostratigraphical as well as from sedimentological and palaeontological perspectives (Bréhéret, 1988; Bréhéret, 1997; Herrle and Mutterlose, 2003; Weissert and Bréhéret, 1991).

2. Methods

2.1. Palynology

16 samples from the Serre Chaitieu section were prepared for palynological analysis. Cleaned and weighed (10 to 12 g) samples were treated with hydrochloric and hydrofluoric acid following standard palynological preparation techniques (Traverse, 1988). The residue was sieved with an 11-µm mesh-sieve and a first set of strew mounts was prepared for palynofacies analysis. A short oxidation with HNO3 was performed on all residues before the preparation of a second set of strew mounts for palynological purposes. Lycopodium marker spores were added prior to preparation to receive absolute counts per gram sediment. For palynofacies analysis the following major categories of particulate OM were distinguished: Amorphous organic matter (AOM), opaque and translucent phytoclasts, cuticles, dinoflagellate cysts, other algae, foraminifera test linings and sporomorphs. Quantitative analysis involved three steps: (i) for palynofacies analysis, a minimum of 350 particles were counted per sample (excl. AOM), (ii) a minimum of 200 palynomorphs were counted for the determination of the absolute abundances of terrestrial and marine palynomorphs, (iii) a minimum of 200 sporomorphs were determined and counted for the pollen and spores assemblage and at least one slide per sample was screened for additional sporomorph taxa.

2.2. Calcareous nannoplankton

Quantitative analyses of calcareous nannofossils were performed on 16 samples using the random settling technique of Geisen et al. (1999). At least 300 individuals were counted per sample in random traverses at x1250 magnification. In addition to total abundance of calcareous nannoplankton, Discorhabdus rotatorius, Zeugrhabdotus erectus, Watznaueria barnesae, Assipetra infracretacea, Rucinolithus terebrodentarius and Nannoconus spp. have been counted separately because of their special palaeoecological and palaeoceanographic significance. In order to assess surface water productivity the nutrient index (NI) of calcareous nannofossils was calculated following Herrle et al. (2003b), where the high- productivity assemblage comprises Z. erectus, D. rotatorius, and the low-fertility assemblage

Chapter 3 47 consists of W. barnesae. To assess nannofossil preservation, light microscope identification of etching and overgrowth effects was used (Bown and Young, 1999).

2.3. Total organic carbon and carbonate carbon content

Inorganic carbon (IC) contents of 36 samples were determined using an UIC CM 5012 Coulomat; total carbon (TC) contents were measured on a CNS Elemental Analyser (Carlo Erba Instruments). Total organic carbon (TOC) contents were calculated from the difference between TC and IC contents.

3. Studied sections, lithology and stratigraphy

The studied OAE 1a interval has been sampled at the Serre Chaitieu section, which is located 20 km southeast of Die, about 1 km south of the village Lesches-en-Diois, Département Drôme, SE France (Fig. 1, 2). The studied interval encompasses 12 m and is mainly composed of dark-grey marls, which are highly bioturbated with Chondrites and Planolites as the most common trace fossils (Bréhéret, 1997). Six finely laminated, dark-grey to black horizons, ranging in thickness from 20 cm to 35 cm, are intercalated within the homogenous marly succession. The individual horizons exhibit submillimetre-scale lamination. They are essentially devoid of bioturbation and are referred to as paper-shales. According to Bréhéret (1994) the lowermost 4 paper shale horizons represent the expression of the OAE 1a in the Vocontian Basin, termed Niveau Goguel interval. Based on a hydrogen index (HI) of up to

500 mg HC/g TOC and a oxygen index (OI) of 0 to 50 mg CO2/g TOC, the sedimentary OM of the Niveau Goguel interval has been classified as type II kerogen, indicating a marine phytoplanktonic and/or bacterial origin (Bréhéret, 1994). Based on biostratigraphic results (Bréhéret, 1994; Bréhéret, 1997; Herrle and Mutterlose, 2003; Moullade, 1966) the studied interval comprises the Deshayesites deshayesi/Tropaeum bowerbanki ammonite Zones and the middle part of the Leupoldina cabri planktic foraminiferal Zone (Fig. 2). The first occurrence of Eprolithus floralis can be recognized between the paper shales PS-3 and PS-4 which marks the onset of the NC7A (Rhagodiscus angustus) and the end of the NC6 (Chiastozygus litterarius) calcareous nannofossil Zones.

Chapter 3 48

Drowned platform facies B Shallow open marine environments C 40° Deep open marine environments pre-Triassic basement Gravity reworked siliciclastics Massif Central Grenoble Internal zones Southern Laurasian province (subtropical to warm-temperate) Valence of the Alps Die 30° Serre Chaitieu section Vocontian Basin

Digne

Castellane 20° Provence Platform Nice Intermediate floral belt A N

an Sea Northern Gondwana province France (arid to semi-arid) Mediterrane 0 50 km N 10° B Spain 0 500 km

Fig. 1: (A) Location of the Vocontian Basin in SE France. (B) Spatial distribution of different depositional settings within the Vocontian Basin during the mid-Cretaceous. The location of the Serre Chaitieu section is marked with an asterisk. Map modified after Arnaud and Lemoine (1993). (C) Plate tectonic reconstruction of the W’ Tethys realm during the mid-Cretaceous (~115 Ma). The position of the Vocontian Basin is marked with an asterisk. Floral provinces and inferred climates after Brenner (1976) and Hochuli (1981). Map modified after Geomar map generator (www.odsn.de).

The studied interval displays a prominent δ13C excursion, which allows correlation with time- equivalent sections on a global scale (Herrle et al., 2004; Weissert and Bréhéret, 1991). Detailed chemostratigraphic correlation with the Cismon section of northern Italy shows, that the entire lower part (0 to 6.5 m) of the Serre Chaitieu section corresponds to the OAE 1a interval. Furthermore, the correlation reveals, that the lowermost part of the OAE 1a (incl. the global negative carbon isotope excursion) is not exposed at the Serre Chaitieu section (Heimhofer et al. submitted).

Chapter 3 49 Substage Ammonite Zone Foraminiferal Zone Nannofossil Zone Meters Lithology

P. n. T. b. Faisceau Nolan

Niveau Fallot

Niveau Noir Calcaire

Niveau Noir

Niveau Clairs Niveau Blanc

Niveau studied Goguel interval

black shale dark-grey marlstone marly limestone

Fig. 2: Lithological column and key beds of the Serre Chaitieu section (Vocontian Basin, SE France) plotted against biostratigraphy. The studied OAE 1a interval is located at the base of the Serre Chaitieu section and encompasses the laminated paper-shales of the Niveau Goguel. Planktic foraminiferal and ammonite biostratigraphy after Bréhéret (1997 and references therein), calcareous nannofossil zonation after Herrle and Mutterlose (2003). D. des., Deshayesites deshayesi; T. bow., Tropaeum bowerbanki; P. n., Parahoplithes nutfieldiensis; L. cabri, Leupoldina cabri; G. ferreolensis, Globigerinelloides ferreolensis; G. alger., Globigerinelloides algerianus; H. trocoidea, Hedbergella trocoidea; T. b., Ticinella bejaouaensis; C. litt., Chiastozygus litterarius.

4. Palaeogeographic and palaeophytogeographic framework

During the mid-Cretaceous, the Vocontian Basin was situated at a palaeolatitude of 25° to 30°N (Hay et al., 1999), forming part of the northern continental margin of the Alpine Tethys Ocean (Fig. 1c). The Marnes Bleues formation, a thick monotonous succession of grey to

Chapter 3 50

dark-grey marls intercalated with calcareous marls, limestones and numerous Corg-rich black shale horizons was deposited in the basin between Early Aptian and Early Cenomanian times (Bréhéret, 1997). Accumulation of fine-grained sediments in the Vocontian Basin was largely confined to pelagic and hemipelagic environments. The basin was surrounded by slope and platform settings, resulting in the intercalation of hemipelagic facies with shallow-water sediments in marginal settings (Arnaud and Lemoine, 1993). To the east, the basin was open towards the Tethys Ocean facilitating exchange with Tethyan water masses. The studied section was situated in the northern, deep marine part of the Vocontian Basin, several tens of km south of the northern palaeo-margin (Arnaud and Lemoine, 1993). According to Brenner (1976), four major floral provinces can be distinguished during the Barremian to Cenomanian, which includes the Northern and the Southern Laurasian provinces as well as the Northern and the Southern Gondwana provinces (Fig. 1c). These floral belts were broadly latitudinally arranged and exhibited progressive compositional changes with increasing latitude, despite low equator-to-pole thermal gradients (Batten, 1984). The studied area was located within the southern part of the Southern Laurasian province of Brenner (1976). The palynofloral assemblage of this province is characterised by numerous and varied pteridophyte spores - namely by spores of gleicheniaceous and schizaeaceous affinity - and by various bisaccate pollen of the Podocarpaceae and Pinaceae. Other gymnosperm pollen such as Classopollis (Cheirolepidiaceae) and Araucariacites (Araucariaceae) represent common elements of the floral assemblages of this province (Batten, 1984; Brenner, 1976; Vakhrameyev, 1991). According to Brenner (1976), Vakhrameyev (1978) and Chumakov et al. (1995) the Southern Laurasian province was characterized by a warm-temperate to subtropical humid climate. The Northern Gondwana province further south is dominated by gymnosperm pollen like Callialasporites, Araucariacites and large numbers of Classopollis. In addition, Ephedripites and Cycadopites represent highly diverse genera, whereas pteriodophyte spores show low diversity. Bisaccate pollen are virtually absent in these assemblages. The climate of the Northern Gondwana province is regarded as arid to semi-arid (Chumakov et al., 1995; Vakhrameyev, 1991). Hochuli (1981) identified an intermediate floral belt in between the Southern Laurasian and the Northern Gondwana provinces, which was characterised by palynofloral elements from both provinces.

Chapter 3 51

5. Results

5.1. Preservation of the particulate OM

Although the sporopollenin composition of pollen and spore walls makes them resistant to degradation, chemical and biological processes during transport and deposition as well as post-depositional alteration can corrode or even destroy palynomorphs (Traverse, 1988). In general, palynomorphs are more affected by degradation processes than refractory organic material (e.g. phytoclasts), which can result in an enrichment of the latter (Tyson, 1995). Furthermore, thin-walled pollen grains are less resistant to biological/chemical alteration and more easily decomposed than thick-walled spores and pollen, leading to preferential preservation of particular thick-walled sporomorph groups. In order to exclude a strong preservational bias of the studied fossil palynofloral assemblages, the preservation of the particulate OM has been carefully examined. Visually, the preservation of the palynomorphs is good to excellent. The consistent occurrence of well-preserved, fine-sculptured and thin-walled angiosperm pollen (e.g. Retimonocolpites spp; Clavatipollenites spp.) throughout the studied interval indicates the absence of a strong preservational bias towards more robust, thick-walled sporomorphs. In addition, the chemically less stable palynomorphs and the resistant phytoclasts fraction show similar variations in absolute abundances (particles/g sediment), which is expressed in the good correlation of the two particle groups (R2 = 0.66; Fig. 3a). This indicates that selective preservation of palynomorphs in OC-rich horizons does not control the observed distribution patterns of the spore-pollen assemblage.

Thermally unaltered conditions for the sedimentary OM in the section is inferred from Tmax values of 420 to 435°C (Bréhéret, 1994), unchanged colouring of the palynomorphs (thermal alteration index < 2) and moderate to strong UV fluorescence of the amorphous fraction and the palynomorphs. In addition, several biomarker maturity indices including the 22S/(22R +

22S)-hopane (C31) index, the Mor/(Mor + Hop) index and the Ts/(Ts + Tm) confirm the immature stage of the organic matter with respect to hydrocarbon generation (Peters and Moldowan, 1993).

Chapter 3 52

(A) (B)

40 R2 = 0.66 2 n = 16 R = 0.86 10.0 n = 16 30 8.0

6.0 20

4.0

Pollen (grains/mg sed.) 10

2.0 Sporomorphs (grains/mg sed.)

0.0 0 0 20 40 60 80 100 0 1020304050607080 Phytoclasts (grains/mg sed.) Dinoflagellate cysts (grains/mg sed.)

Fig. 3: (A) Cross-plot of phytoclast and pollen absolute abundances of the Serre Chaitieu section (Vocontian Basin, SE France). Squares correspond to samples from paper shales, dots represent samples from bioturbated, marly lithology. Note the notable correlation between the chemically less stable pollen grains and the more refractory phytoclast fraction. This indicates that degradation processes are of minor importance and that the observed variations in the palynomorphs assemblages are not controlled by selective preservation. (B) Cross-plot of dinoflagellate and sporomorph absolute abundances from the Serre Chaitieu section. The two different particle groups show a strong correlation, which emphasises the congruent pattern of the two records.

5.2. Composition and distribution of the particulate OM

The studied section is characterised by low to moderate CaCO3 (9.0 to 36.0 %) and TOC (0.4 to 2.3 %) contents, respectively (Fig. 4a). The bioturbated marls in the lower part of the section (0 to 4.3 m) show slightly enriched TOC values (mean 1.3 %). In contrast, the upper part (6.5 to 12.0 m) displays comparatively low TOC (mean 0.7 %), but increased CaCO3 contents. Higher TOC contents (up to 2.3 %) are restricted to the occurrence of finely laminated paper-shales. Only horizon PS-4 exhibits an exceptional low TOC content of only 0.5 %. The major constituent of the particulate OM is formed by amorphous organic matter (AOM), which accounts for ~95 % within the OAE 1a interval and for 70 to 90 % in the bioturbated marls above. Two different types of AOM can be distinguished. Type A is composed of glossy, inclusion-rich floccules of orange-brown colouring and exhibits moderate to strong fluorescence. In contrast, type B is characterised by matt, shard-like, grey-brown particles with weak to moderate fluorescence. Type B represents the major constituent in the upper part

Chapter 3 53 of the section. The dominance of type A-AOM in the lower part (OAE 1a interval) and within the paper-shales is interpreted to reflect dys- to anoxic bottom water conditions (e.g. Tribovillard and Gorin, 1991; Tyson, 1995). Again, horizon PS-4 represents an exception and comprises predominantly AOM of type B. The phytoclast fraction is dominated by equidimensionally shaped particles, predominantly < 20 µm in size. Together, opaque and translucent phytoclasts account for 32.2 to 52.0 % (mean 37.3 %) of the particulate OM (excl. AOM). Their frequency pattern displays no distinct trend or variations throughout the studied interval. The observed phytoclast assemblage is typical for deep-water sediments, which are generally characterized by the dominance of small, equidimensional, oxidized woody debris and some windblown charcoal (Habib, 1982; Tyson, 1995). The palynomorphs fraction (Fig. 4b) is clearly dominated by dinoflagellate cysts which range from 51.2 to 81.3 % (mean 67.4 %) in relative abundance. The high amount of dinoflagellate cysts emphasises the open marine conditions of the depositional setting. Sporomorphs (incl. spores and pollen) account for 10.7 to 38.9 % (mean 23.6 %) and are slightly enriched in the OAE 1a interval as well as in the laminated horizons (mean of 27.4 %) compared to the upper part (mean of 22.5 %). Foraminifera test linings display strong fluctuations, ranging from 18.4 % to complete absence (mean 9.0 %) in particular paper- shales. Absolute abundances of continent-derived sporomorphs and marine dinoflagellate cysts are displayed in Fig. 4c. Both palynomorph groups show a strong increase in absolute abundances within paper-shale horizons (PS-1, 2 and 5) compared to background values. Peak values of sporomorphs are as high as 4 × 104 sporomorphs/g sediment whereas dinoflagellate cysts account for up to 7 × 104 cysts/g sediment. In addition, continent-derived sporomorphs and marine-derived dinoflagellate cysts display essentially similar variations throughout the section, which is expressed in a strong correlation of the two records (R2 = 0.86, Fig. 3b). The geochemical and palynofacies results are summarized in Table 1.

Chapter 3 54 high 40 (D) 20 60 nutrient index (NI) low 0 calcareous nannofossil 4 10 x 8 4 10 x 6 4 10 x marine dinocysts pollen and spores (C) 4 4 abundance (grains/g sed.) 10 x 2 palynomorphs absolute 0 (B) 40 pollen and spores marine dinocysts foraminifera linings abundance 20 60 80 (% of total palynomorphs) palynomorphs relative 0100 40 23 (wt %)

(wt %) 3 20 30 (A) TOC CaCO

01 010 OAE 1a OAE

PS-6 PS-5 PS-4 PS-3 PS-2 PS-1

Lithology Meters 8 4 2 0 6

12 10

Fig. 4: Selected geochemical and palaeontological parameters across the OAE 1a interval, Serre

Chaitieu section (Vocontian Basin, SE France) plotted against lithology. (A) TOC and CaCO3 content, (B) relative abundances of dinoflagellate cysts, sporomorphs and foraminifera test linings expressed as percentages of the total palynomorphs fraction, (C) absolute abundances of sporomorphs and dinoflagellate cysts expressed as grains per g sediment, (D) calcareous nannofossil nutrient index. Dotted lines mark the position of laminated paper shales. For lithological explanations see Fig. 2.

Chapter 3 55

5.3. Composition and distribution of the palynoflora

We distinguished 18 groups of spores and 19 groups of pollen grains in the microflora of the Serre Chaitieu section (Fig. 5). The quantitatively most important group is represented by Classopollis spp. which accounts for 16.7 to 42.3 % (mean 29.2 %) of the entire assemblage. Classopollis spp. shows a gradual increase across the OAE 1a interval (from 24.6 up to 42.3 %) and a subsequent decline in the upper part of the section. Other common gymnosperm pollen include Araucariacites spp. (5.9 to 12.9 %; mean 7.2 %), Inaperturopollenites spp. (2.5 to 11.3 %; mean 5.7 %) and Sciadopityspollenites spp. (1.0 to 5.5 %; mean 2.8 %). Exesipollenites spp. (2.8 to 12.4 %; mean 6.0 %) displays low abundance across the OAE 1a interval (mean 3.7 %) but is relatively common in the upper part (mean 9.7 %). Various bisaccate pollen (incl. Podocarpidites spp., Alisporites spp.) account for less than 13.7 % in most samples. Slightly increased abundance of bisaccates (up to 20.8 %) is essentially restricted to the occurrence of paper shales (PS-1, 2 and 5). Common representatives of the angiosperm pollen group include Striatopollis spp., Clavatipollenites spp. and Retimonocolpites spp. and form a rare, but consistent element of the observed floral assemblage (< 2.0 %; mean 0.5 %). Pteridophyte spores represent another important constituent and exhibit a slight, but consistent increase within the laminated paper shales. Deltoidospora spp. (11.4 to 20.6 %; mean 15.1 %) and Gleicheniidites spp. (2.5 to 10.9 %; mean 6.1 %) dominate the spore spectrum, whereas other spores like Cicatricosisporites spp., Leptolepidites spp. and Retitriletes spp. are quantitatively of minor importance.

5.4. Composition of dinoflagellate cyst and calcareous nannofossil assemblages

The dinoflagellate assemblage of the Serre Chaitieu section has been studied qualitatively (Fig. 6). A total of 61 different dinoflagellate taxa have been identified on genera or species level. The relatively homogenous assemblage displays an Early Aptian composition and comprises many long ranging forms. In the studied interval, the most important dinoflagellate marker species for the Early Aptian include Pseudoceratium securigerum, Heslertonia heslertonensis, Oligosphaeridium asterigerum, Druggidium apicopaucicum and Rhynchodiniopsis aptian. The Achomosphaera spp. and Spiniferites spp. groups have not been differentiated on species level. The diversity distribution displays a relatively stable pattern throughout the succession (mean of 20 taxa per sample) with a slight increase towards the top (Fig. 6).

Chapter 3 56 spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. group Meters Stage Lithology Nannofossil Zone Foraminiferal Zone

12 PS-6 Bisaccate Pollen Araucariacites Classopollis Vitreisporites pallidus Other Spores Ephedripites Exesipollenites Inaperturopollenites Sciadopityspollenites Other Gymnosperms Afropollis Angiosperms Cicatricosisporites Deltoidospora Gleicheniidites Retitriletes

10 (NC7A)

PS-5

8 R. angustus

PS-4 6 PS-3

PS-2 Lower Aptian Leupoldina cabri PS-1 4 (NC6) OAE 1a

2 C. litterarius

0 0 1020304050%

Fig. 5: Quantitative distribution pattern of selected spore and pollen types across the OAE 1a interval, Serre Chaitieu section (Vocontian Basin, SE France). Note that the entire lower part of the section (0- 6.5 m) corresponds to the OAE 1a. Relative abundances of the spores and pollen are expressed as percentages of the total sporomorph assemblage. Biostratigraphy of the Serre Chaitieu section after Herrle and Mutterlose (2003).

An exceptional high diversity of 28 taxa can only be observed in the uppermost paper shale horizon of the OAE 1a interval (PS-4). The calcareous nannofossil assemblage is dominated by (in descending order) Watznaueria barnesae, Zeugrhabdotus erectus, Discorhabdus rotatorius, Assipetra infracretacea, Rucinolithus terebrodentarius and Nannoconus spp. representing 28.5 to 74.4 % (mean 46.6 %) of the total assemblage. The portion of the most dissolution-resistant species W. barnesae ranges from 17.5 to 39.2 % of the total assemblage. Following Thierstein (1980) and Roth & Bowdler (1981) portions of W. barnesae > 40 % often indicate dissolution to the extent that the original assemblages no longer yield a primary signal. Both the low percentages of W. barnesae and the etching and overgrowth ranking of E1 to E1-2 and O1 (slightly etched and

Chapter 3 57 overgrown of coccoliths elements) of the studied samples indicate a good preservation of the calcareous nannofossil assemblage. The calculated nutrient index (NI) varies between 24.5 and 48.1 % (mean 38.1 %). Low percentages (< 38 %) can be recognized in the lower part of the succession (Fig. 4d). Just below the onset of the paper-shales the percentages of the NI increase, characterized by minor fluctuations around 40 %. Highest percentages of A. infracretacea/R. terebrodentarius (up to 7.6 %) occur below the paper shale interval and between the paper shales PS-3 and PS-4. Nannoconus spp. is characterized by increasing percentages (up to 2.8 %) in the uppermost part of the studied succession.

6. Discussion

6.1. Palaeo-environmental and -climatic significance of the micropalaeontologcial results

The palynofloral assemblage of the studied succession reflects a rich and diverse flora. Besides various fern families (e.g. Gleicheniaceae, Schizaeaceae, Osmundaceae, Dicksoniaceae), different types of ginkgophytes, cycads, bennettitales and several conifer families (Araucariaceae, Cheirolepidaceae, Taxodiaceae and Podocarpaceae) can be identified (Balme, 1995). The rare but consistent occurrences of angiosperm pollen in the Early Aptian deposits mark the incipient radiation of this plant group. Based on the observed palynofloral association, a tentative interpretation of the corresponding habitats can be given. In Mesozoic assemblages, ferns are considered to be common elements of lush and moist vegetation along riversides and/or coastal lowlands (Mohr, 1989; Van Konijnenburg - Van Cittert and Van der Burgh, 1989). Therefore, the common occurrence of pteridophytes in the Serre Chaitieu section indicates humid and warm habitats in the corresponding hinterland. Evidence for predominantly lowland and/or coastal vegetation can be inferred from the abundant occurrence of various pollen of bennettitalean and araucariacean affinity (Abbink, 1998; Vakhrameyev, 1991). In contrast, the large quantities of Classopollis spp. are produced by the xerophythic (drought-resistant) and thermophythic Cheirolepidaceae, which are considered to reflect well-drained slope and upland environments (Vakhrameyev, 1982; Vakhrameyev, 1991) or mangrove-type, coastal vegetation (Watson, 1988). Abundance patterns of Classopollis pollen are a valuable indicator of the prevailing climate. High numbers of Classopollis spp are considered to reflect warm and arid conditions whereas low abundances correspond to cooler and more humid climates (Vakhrameyev, 1982). Bisaccate pollen-producing Podocarpaceae and Pinaceae are indicative of relatively dry upland

Chapter 3 58 10 20 30 (# taxa) diversity

0 Dinoflagellate

spp. Prolixosphaeridium Prolixosphaeridium

Microdinium opacum Microdinium

Kleithriasphaeridium loffrense Kleithriasphaeridium

spp. Chytroeisphaeridia Chytroeisphaeridia

Prolixosphaeridium parvispinum Prolixosphaeridium

Kleithriasphaeridium fasciatum Kleithriasphaeridium

Sepispinula huguoniotii Sepispinula

spp. Pseudoceratium Pseudoceratium

spp. Kalyptea Kalyptea

imparilis imparilis cf. cf.

Dingodinium cerviculum Dingodinium

spp. Coronifera Coronifera

spp. Dapsilidinium Dapsilidinium

Wallodinium lunum Wallodinium

Odontochitina Odontochitina

spp. Callaiosphaeridium Callaiosphaeridium

Rhynchodiniopsis aptiana Rhynchodiniopsis

spp. Hystrichosphaeropsis Hystrichosphaeropsis

Exochosphaeridium phragmites Exochosphaeridium

spp. Kleithriasphaeridium Kleithriasphaeridium

Hystrichodinium pulchrum Hystrichodinium

spp. Exochosphaeridium Exochosphaeridium

spp. Cleistosphaeridium Cleistosphaeridium

spp. Chlamydophorella Chlamydophorella

Palaeoperidinium cretaceum Palaeoperidinium

Kleithriasphaeridium simplicispinum Kleithriasphaeridium

Gonyaulacysta cretacea Gonyaulacysta

spp. Gardodinium

Florentina deanei Florentina

spp. Batiacasphaera Batiacasphaera

Pterodinium spp. Pterodinium

spp. Oligosphaeridium Oligosphaeridium

Kiokansium polypes Kiokansium

Hystrichosphaerina schindewolfii Hystrichosphaerina

Gardodinium trabeculosum Gardodinium

spp. Druggidium Druggidium

Druggidium apicopaucicum Druggidium

Cribroperidinium orthoceras Cribroperidinium

Coronifera oceanica Coronifera

Aptea polymorpha Aptea

spp. Trichodinium Trichodinium

Pinocchiodinium erbae Pinocchiodinium

Oligosphaeridium asterigerum Oligosphaeridium

Heslertonia heslertonensis Heslertonia

spp. Florentinia Florentinia

spp. Circulodinium Circulodinium

Tanyosphaeridium variecalamum Tanyosphaeridium

spp. Systematophora Systematophora

spp. Subtilisphaera Subtilisphaera

spp. Spiniferites Spiniferites

Pterodinium cingulatum Pterodinium

Pseudoceratium securigerum Pseudoceratium

Oligosphaeridium complex Oligosphaeridium

Odontochitina operculata Odontochitina

Gonyaulacysta helicoidea Gonyaulacysta

spp. Dingodinium Dingodinium

spp. Cribroperidinium Cribroperidinium

spp. Cometodinium Cometodinium

Cerbia tabulata Cerbia

Cassiculosphaeridia reticulata Cassiculosphaeridia

Callaiosphaeridium asymmetricum Callaiosphaeridium

spp. Achomosphaera Achomosphaera

PS-6 PS-5 PS-4 PS-3 PS-2 PS-1

Lithology

C. litterarius litterarius C. R. angustus angustus R. Nannofossil Zone Nannofossil (NC6) (NC7A)

Leupoldina cabri Leupoldina Foraminiferal Zone Foraminiferal

Stage Lower Aptian Lower Meters 8 4 2 0 6

12 10

Fig. 6: Stratigraphical distribution of dinoflagellate taxa in the Serre Chaitieu section (Vocontian Basin, SE France) ordered to first occurrences. Selected Early Aptian dinoflagellate marker species are printed in bold type. Dinoflagellate cyst diversity represents the number of taxa per sample. Biostratigraphy of the Serre Chaitieu section after Herrle and Mutterlose (2003). Dotted lines mark the position of laminated paper shales. For lithological explanations see Fig. 2.

Chapter 3 59 vegetation and generally dominate in boreal associations (Abbink, 1998; Vakhrameyev, 1991). In general, the vegetation patterns remain essentially stable and the dominant sporomorph forms persist throughout the studied record. Distinct changes can only be observed in the abundance of the thermophilic Cheirolepidaceae (Classopollis spp.) as well as in plants of questionable bennettitalean or taxodiacean affinity (Exesipollenites spp.). The increase in cheirolepidaceans during the OAE 1a interval suggests a shift towards more arid conditions. Due to the ambiguous botanical affinity and habitat preferences of the Exesipollenites- producing plants, a climatic interpretation can not be given. Besides these fluctuations, we observe no indication for major climatic disturbances, accompanied by prominent shifts in the major floral belts. The observed palynoflora is typical for the late Early Cretaceous Southern Laurasian floral province (Fig. 1c). Some minor influence of the Northern Gondwana province is reflected in the rare, but consistent occurrence of Afropollis spp. and Ephedripites spp. On the other hand, the pollen record provides no evidence for a southward dispersion of boreal vegetation (e.g. bisaccate pollen of Pineacean affinity) during or in the aftermath of the OAE 1a interval. Even though bisaccate abundance is in general considered as a palaeoclimatic indicator (e.g. Vakhrameyev, 1991) the observed variations in bisaccate pollen in the Serre Chaitieu section might rather reflect transportation bias than a real vegetation signal. Due to their specific morphology, bisaccate pollen can be dispersed easily by atmospheric or aquatic pathways (Traverse, 1988). According to Heusser and Balsam (1977) the capability of bisaccates to float for a relatively long time period explains their relative increase in shelf sediments with increasing distance from the shoreline. Hence, the observed increase in bisaccates might be related to sea-level fluctuations and/or changes in runoff patterns during times of black shale formation. The pollen record of the Vocontian Basin contrasts with the results of Hochuli et al. (1999) who reported a significant increase in bisaccate pollen abundance from ~20 % within to ~80 % above the OAE 1a interval at the the Cismon site (Belluno Trough, N Italy). Based on the findings from the Vocontian Basin, the Cismon pollen record is considered to reflect a major change in the paleoceanographic current pattern rather than a major floral shift due to global cooling (Heimhofer et al. submitted).

The common occurrence of the dinoflagellate groups Cribroperidinium spp. and Circulodinium spp. in all studied samples is interpreted to indicate inner neritic conditions whereas some open marine influence is reflected in the consistent occurrences of the

Chapter 3 60

Oligosphaeridium spp. and Spiniferites spp. groups (Wilpshaar and Leereveld, 1994). Similar assemblages have been documented from the Southern Alps (Cismon section) by Torricelli et al. (2000) as well as from SE France (Gare de Cassis section) by Masure et al. (1998). The observed dinoflagellate cyst assemblage and diversity patterns display no distinct variations across the OAE 1a interval (Fig. 6). We observe neither a significant impoverishment nor a strong diversity increase of the organic-walled plankton within or above the OAE 1a interval. The slight increase in diversity towards the top of the studied interval is considered to reflect the response of the dinoflagellate cyst assemblage to a rising sea level (Tyson, 1995). This is in accordance with the results of Wilpshaar and Leereveld (1994) who report a significant shift of the dinoflagellate cyst assemblages towards a more oceanic association due to a late Early Cretaceous sea-level rise in the Vocontian Basin. The quantitative analysis of Toricelli (2000) from the Cismon site (Belluno Trough, N Italy) displays a decrease in dinoflagellate cyst diversity in the lowermost part of the OAE 1a interval (not accessible in the Serre Chaitieu section) and, similar to the Vocontian Basin record, a gradual diversity increase throughout the black shale interval and the overlying strata.

The calcareous nannofossil nutrient index displays no evidence for a major change in surface water productivity during the formation of the OAE 1a interval (Fig. 4d). We observe no consistent pattern in nutrient index corresponding to the occurrence of individual paper shale horizons. In comparison to earlier studies on the OAE 1b from the Vocontian Basin by Herrle et al (2003b), the observed variations of surface water productivity across the OAE 1a interval are rather moderate. The calculated calcareous nannofossil nutrient index indicates low to moderate surface water productivity conditions during formation of the OAE 1a and a subsequent increase in the aftermath of the black shale episode. These findings are in accordance to earlier studies on calcareous micro- and nannofossils. According to Luciani et al. (2001), W’ Tethys surface waters were characterised by moderate palaeofertility conditions during the OAE 1a interval. Premoli Silva et al. (1999) pointed out, that eutrophication during the Early Aptian OAE 1a was less intense compared to the OAE 2 interval (Cenomanian-Turonian boundary event).

In summary, the palynological results imply, that the continental hinterland of the Vocontian Basin was characterised by diverse but relatively stable palaeoenvironments during the studied interval. A change towards more arid climatic patterns during the OAE 1a interval is

Chapter 3 61 documented in the rising abundance of Classopollis spp. In contrast to Hochuli et al. (1999), we observe no prominent increase in boreal pollen forms in the aftermath of the OAE 1a. Minor variations in the dinoflatellate cyst assemblages as well as the calcareous nannofossil nutrient index provide no evidence for enhanced surface water productivity during the late Early Aptian black shale episode in the Vocontian Basin.

6.2. Changes in sedimentation rates and OC accumulation across OAE 1a

In general, the flux of pollen and spores to the depositional environment is closely tied to detrital input, both predominantly controlled by continental runoff (Traverse, 1994). However, along arid coasts as well as in hemipelagic to pelagic environments, atmospheric transportation of pollen can be of significant importance (Dupont and Wyputta, 2003; Melia, 1984). This results in the decoupling of fluvial siliciclastic input and airborne pollen flux. In the Serre Chaitieu section, prominent changes are displayed in the absolute abundances of terrestrial sporomorphs and marine dinoflagellate cysts (Fig. 4c). Even though the two palynomorphs groups are affected by completely different processes during transportation and deposition, they display congruent variations in absolute particle abundance. We assume that a large part of the terrestrial sporomorph fraction has been transported via atmospheric pathways to the depositional setting and that the input fluxes of both, marine dinoflagellate cysts and sporomorphs were roughly constant. In consequence, the observed fluctuations in both particle groups are essentially controlled by changes in sedimentation rates. This in turn gives way to a tentative estimation of sedimentation rates (SR) and OC accumulation across the OAE 1a interval. Based on time series analysis and biostratigraphic data, an average SR of 3.0 to 3.5 cm ka-1 has been calculated for the Late Aptian part of the Serre Chaitieu section by Kössler et al. (2001) and Herrle et al. (2003a). A similar SR is assumed for the average background sedimentation represented by dark-grey, bioturbated marls in the upper part of the section (6.5 to 12 m). In combination with absolute palynomorph abundances (tentatively regarded as a constant flux), this results in significant SR changes. Increased palynomorph abundances within paper-shales correspond to very low sedimentation rates (SR as low as 0.5 cm ka-1) and therefore reduced dilution by siliciclastic detrital material. In contrast, decreased abundances within bioturbated marls indicate periods of higher sediment flux and increased siliciclastic input (SR between 2.0 and 3.0 cm ka-1). This is also displayed in fluctuations of the carbonate carbon content record. Peak values in CaCO3 content within OC-rich paper- shales reflect lowered siliciclastic dilution of the pelagic carbonate sedimentation.

Chapter 3 62

TOC content and SR exhibit a notable inverse correlation (R2 = 0.61; Fig. 7). Low SR corresponds to increased TOC contents and vice versa. The occurrence of well-preserved OM during periods of reduced SR seems to be somehow contradictory. To prevent the OM from degradation during its relatively long exposure at the sea floor, strongly oxygen-depleted conditions are required (Tyson, 1995). Evidence for anoxic bottom water conditions is provided by the occurrence of finely laminated, non-bioturbated facies accompanied by relatively low abundances of foraminiferal test linings, both indicating decreased benthic activity. Based on organic facies analysis, similar low-oxygen bottom water conditions have been inferred for several horizons of the time-equivalent Livello Selli interval in Italy (Baudin et al., 1998; Hochuli et al., 1999; Menegatti et al., 1998) and for the Early Albian Niveau Paquier (OAE 1b) in the Vocontian Basin (Tribovillard and Gorin, 1991).

3.0 R2 = 0.61 n = 16 2.5

2.0

1.5 TOC (wt. %) TOC 1.0

0.5

0.0 0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 Sedimentation rate (cm ka-1)

Fig. 7: Cross-plot of inferred sedimentation rate and TOC content of the Serre Chaitieu section (Vocontian Basin, SE France). Squares correspond to samples from paper shales, dots represent samples from bioturbated, marly lithology. The inverse correlation between the two parameters clearly indicates that OC accumulation in the studied section was not controlled by the effect of increasing sedimentation rate.

Chapter 3 63

To evaluate the sedimentary OC contents separately from the input of other components, organic carbon mass accumulation rates (OC MAR) have been calculated. OC MAR fluctuate between 0.02 and 0.06 gC cm-2 ka-1 throughout the section. Comparable OC MAR have been determined for Lower Aptian sediments located in the Eastern Atlantic (Stein et al., 1986) and Pacific oceans (Bralower and Thierstein, 1987). Our estimates for the Serre Chaitieu section are in the range of OC MAR for the present day Panama and Canary Basins (Bralower and Thierstein, 1987; Stein et al., 1986). Minor variations in OC MAR can be observed throughout the studied interval. Whereas marly, bioturbated lithologies exhibit values of ~0.03 to 0.06 gC cm-2 ka-1, the paper shale horizons (except PS-3) show similar or even lower OC MAR between ~0.02 and 0.04 gC cm-2 ka-1. These results indicate that the accumulation of OC during the OAE 1a interval was not enhanced compared to post-OAE times in the Vocontian Basin.

6.3. Palaeoceanographic implications

According to several authors (e.g. Bellanca et al., 2002; Jenkyns, 1999; Weissert et al., 1998) the formation of the Early Aptian OAE 1a black shale reflects the complex interplay of enhanced hydrological cycling and accelerated continental weathering during a period of exceptional warmth. The intensified transport of continent-derived detrital material towards the basins e.g. during episodes of increased runoff is interpreted to result in enhanced nutrient levels of oceanic surface waters. High nutrient availability in turn has been interpreted to cause enhanced phytoplankton productivity in surface waters, leading to the deposition of OC-rich sediments. In the Vocontian Basin, several lines of evidence contradict a causal link between accelerated climate-controlled nutrient fluxes, high oceanic palaeoproductivity and the deposition of the OAE 1a black shales. Neither the palynofloral record nor the dinoflagellate cyst and calcareous nannofossil assemblages indicate strongly increased hydrological cycling accompanied by a significant increase in surface water primary productivity in the Vocontian Basin. Furthermore, the estimated changes in OC accumulation during and after formation of the OAE 1a provide no evidence for increased palaeoproductivity. Even though, the above mentioned scenario could explain an increase in continent-derived sporomorphs (enhanced runoff) paralleled by an increase in organic-walled plankton (enhanced productivity), the almost straight proportional dependency of the two different proxies (Fig. 3b) suggests similar

Chapter 3 64 input fluxes rather than a complex biologically feedback mechanism to account for the observed pattern. The reduced SR, which characterise the lower part of the OAE 1a and particularly the paper- shale horizons, are interpreted to reflect episodes of pronounced condensation due to decreased siliciclastic input. Furthermore, the episodic occurrence of well-developed bottom water anoxia is accompanied by low SR. Based on the current findings two alternative scenarios are proposed which could account for the OC accumulation during the OAE 1a interval in the Vocontian Basin. Our interpretation suggests fluctuations in (i) sea-level and/or (ii) runoff to account for the above mentioned observations.

(i) Sea-level fluctuations have been addressed by various authors to play a key role for the formation of OC-rich deposits in hemipelagic to pelagic settings during OAE 1a (e.g. Bréhéret, 1994; Erbacher et al., 1996; Strasser et al., 2001). Based on the analysis of stacking patterns, Bréhéret (1994) considered amalgamation and condensation processes to cause the formation of the OAE 1a paper-shales in the Vocontian Basin. According to his model, the deposition of the paper-shales is related to small-scale sea level rises which are superimposed on a major transgressive pulse, or maximum flooding (2nd order sequence). Small- and large- scale sea level rises are supposed to cause a relative decrease in detrital input due to an increase in accommodation space, resulting in condensation in the basinal environments of the Vocontian Basin. Similarly, Strasser et al. (2001) identified several higher-frequency sea- level changes superimposed on a major transgression, which had a marked influence on the formation of the OAE 1a interval along the northern margin of the Alpine Tethys Ocean. The concomitant occurrence of sea-level rise and bottom water anoxia observed in hemipelagic settings has been related to various mechanisms. This includes vertical and lateral shifts of the oxygen minimum zone onto the shelf during transgressive phases (Schlanger and Jenkyns, 1976), reduced mixing of shelf waters due to increasing water depth (Arthur et al., 1987; Tyson, 1995) or increased nutrient flux from coastal lowlands, resulting in productivity-driven anoxia (Erbacher et al., 1996; Jenkyns, 1980). Even though, the observed fluctuations in SR can be well explained with the occurrence of high-frequency sea-level variations, the superimposed low-frequency sea-level rise is not well expressed in a reduction of the estimated SR in the Serre Chaitieu section.

Chapter 3 65

(ii) An alternative explanation for reduced SR involves distinct changes in runoff patterns. Evidence for less precipitation and drier climatic conditions is reflected in the frequency patterns of Classopollis-type pollen (Vakhrameyev, 1982; Vakhrameyev, 1991). The increase in Classopollis spp. from 25 % to > 40 % towards the top of the OAE 1a interval can be interpreted to reflect a shift towards more arid conditions whereas the decline above the OC- rich interval might indicate a return to more humid climate patterns. Such a climatic change is supposed to result in reduced runoff and therefore in a decline of siliciclastic input to the basin. A general increase in aridity during formation of the OAE 1a could probably account for the observed dys- to anoxic conditions documented from various ocean basins. The formation of black shales due to enhanced thermohaline stratification and concomitant oxygen-deficiency in bottom waters during periods of increased aridity has been invoked in previous studies (e.g. Barron and Peterson, 1990; Brass et al., 1982).

7. Conclusions

In the Vocontian Basin, several lines of evidence contradict the previous held view, that the OAE 1a black shales reflect the complex interplay of accelerated hydrological cycling, increased climate-controlled nutrient fluxes and high oceanic primary productivity. Results from the analysis of dinoflagellate cyst and calcareous nannoplankton assemblages as well as tentative estimates of OC accumulation indicate a rather reduced or unchanged palaeoproductivity during times of OAE 1a formation. Similarly, the pollen-based reconstruction of the vegetation patterns in the corresponding hinterland provide no evidence for enhanced humidity and intensified precipitation. In contrast, the observed increase in Classopollis-type pollen across the OAE 1a interval points rather to a shift towards a more arid climate during deposition of the black shales in the adjacent basin. Tentatively estimated sedimentation rates display significant fluctuations across the studied interval and are particularly reduced within laminated, non-bioturbated, OC-rich horizons. The concomitant occurrence of reduced detrital input and oxygen-deficient bottom waters indicates that low- frequency sea-level fluctuations and/or changes in riverine runoff play a key role in the formation of the OAE 1a.

Chapter 3 66

height (m) TOC (%) CaCO3 (%) phytoclasts sporomorphs sporomorphs dino-cysts dino-cysts foraminifera foraminifera sum lycopods dino-cysts sporomorphs NI

% kerogen % palynos total counts % palynos total counts % palynos total counts grains/mg sed. grains/mg sed.

NS-36 11.75 0.8 27.3 NS-35 11.35 0.6 25.7 NS-34 10.95 0.7 30.7 51.1 18.3 43 70.6 166 11.1 26 235 48 20.8 5.4 40.3 NS-33 10.55 0.6 35.7 NS-32 10.15 0.6 22.0 52.0 20.6 43 67.0 140 12.4 26 209 64 13.1 4.0 48.1 NS-31 9.75 0.8 20.1 NS-30 9.35 1.2 15.5 NS-29 9.15 1.8 12.6 37.6 38.9 95 55.3 135 5.7 14 244 27 32.3 22.7 43.9 NS-28 9.05 0.9 18.8 NS-27 8.6 1.1 16.7 NS-26 8.2 0.5 21.1 40.2 10.7 26 72.3 175 16.9 41 242 109 10.4 1.5 37.7 NS-25 7.8 0.4 26.4 NS-24 7.4 0.6 21.0 43.0 18.6 44 70.8 167 10.6 25 236 93 10.8 2.8 38.3 NS-23 7 0.9 23.1 37.5 15.1 34 71.1 160 13.8 31 225 67 14.3 3.0 41.4 NS-22 6.55 0.9 15.4 NS-21 6.45 0.5 16.8 35.6 13.3 32 81.3 195 5.4 13 240 53 22.1 3.6 39.7 NS-20 6.25 0.9 21.5 NS-19 6 1.1 21.7 36.5 25.0 56 61.2 137 13.8 31 224 29 28.3 11.6 41.0 NS-18 5.6 1.6 13.5 NS-17 5.5 2.2 25.8 35.2 27.0 58 73.0 157 0.0 0 215 32 31.7 11.7 44.0 NS-16 5.4 1.3 21.5 NS-15 5.3 1.4 17.4 36.0 22.5 58 67.8 175 9.7 25 258 32 32.8 10.88 33.3 NS-14 5.2 1.7 10.2 NS-13 5.1 2.1 31.5 NS-12 5 1.9 28.4 32.2 31.0 81 51.7 135 17.2 45 261 14 63.1 37.87 46.5 NS-11 4.9 1.3 15.0 NS-10 4.75 1.3 12.5 NS-9b 4.5 2.2 30.9 33.2 30.5 73 69.0 165 0.4 1 239 15 72.0 31.85 30.4 NS-9a 4.4 0.8 8.7 NS-8 3.9 1.6 18.2 38.3 30.3 74 51.2 125 18.4 45 244 33 22.7 13.5 42.8 NS-7 3.35 1.7 15.4 NS-6 2.8 1.7 14.8 36.7 23.6 54 73.4 168 3.1 7 229 39 25.8 8.3 24.5 NS-5 2.15 0.7 13.3 NS-4 1.65 1.0 12.0 42.1 23.5 52 71.5 158 5.0 11 221 46 20.6 6.8 30.3 NS-3 1.1 1.8 14.2 NS-2 0.55 1.1 13.2 37.3 27.9 64 71.2 163 0.9 2 229 38 25.7 10.1 27.6

Table 1: Geochemical and palynofacies data from the Serre Chaitieu section (Vocontian Basin, SE France)

Chapter 3 67

Acknowledgements

Financial support from ETH-project TH-34./99-4 is greatfully acknowledged.

References

Abbink, O.A., 1998. Palynological investigations in the Jurassic of the North Sea region, Universiteit Utrecht, Utrecht, 192 pp. Arnaud, H. and Lemoine, M., 1993. Structure and Mesozoic-Cenozoic evolution of the South-East France Basin (SFB). Géologie Alpine, Série spéciale "Colloques et Excursions", 3, 3-58. Arthur, M.A., Dean, W.E. and Pratt, L.M., 1988. Geochemical and climatic effects of increased marine organic carbon burial at the Cenomanian/Turonian boundary. Nature, 335, 714-717. Arthur, M.A., Jenkyns, H.C., Brumsack, H.-J. and Schlanger, S.O., 1990. Stratigraphy, geochemistry and paleoceanography of organic-carbon rich Cretaceous sequences. In: R.N. Ginsburg and B. Beaudoin (Editors), Cretaceous Resources, Events and Rhythms. NATO ASI Series C. Kluwer Academic Publishers, MasDordrecht, pp. 75-119. Arthur, M.A., Schlanger, S.O. and Jenkyns, H.C., 1987. The Cenomanian-Turonian Oceanic Anoxic Event. Palaeoceanographic controls on organic-matter production and preservation. In: J. Brooks and A.J. Fleet (Editors), Marine Petroleum Source Rocks. Geological Society Special Publications. Geological Society of London, London, UK, pp. 401-420. Balme, B.E., 1995. Fossil in situ spores and pollen grains: an annotated catalogue. Review of Palaeobotany and Palynology, 87, 81-323. Barron, E.J., 1983. A warm, equable Cretaceous: The nature of the problem. Earth-Science Reviews, 19, 305-338. Barron, E.J. and Peterson, W.H., 1990. Mid-Cretaceous ocean circulation: Results from model sensitivity studies. Paleoceanography, 5, 319-337. Batten, D.J., 1984. Palynology, climate and the development of Late Cretaceous floral provinces in the Northern Hemisphere: A review. In: P. Brenchley (Editor), Fossils and climate. John Wiley and Sons, New York, pp. 127-164. Baudin, F., Fiet, N., Coccioni, R. and Galeotti, S., 1998. Organic matter characterisation of the Selli Level (Umbria-Marche Basin, central Italy). Cretaceous Research, 19, 701-714. Beerling, D.J. and Royer, D.L., 2002. Fossil plants as indicators of the phanerozoic global carbon cycle. Annual Review of Earth and Planetary Sciences, 30, 527-556. Bellanca, A., Erba, E., Neri, R., Premoli, S.I., Sprovieri, M., Tremolada, F. and Verga, D., 2002. Palaeoceanographic significance of the Tethyan "Livello Selli" (Early Aptian) from the Hybla Formation, northwestern Sicily: Biostratigraphy and high-resolution chemostratigraphic records. Palaeogeography, Palaeoclimatology, Palaeoecology, 185, 175-196. Berner, R.A., 1994. GEOCARB II: A revised model of atmospheric CO2 over Phanerozoic time. American Journal of Science, 294, 56-91. Bown, P.R. and Young, J.R., 1999. Techniques. In: P.R. Bown (Editor), Calcareous Nannofossil Biostratigraphy. British Micropalaeontological Society Publications. Kluwer Academic Publishers, Dordrecht, Netherlands, pp. 16-28. Bralower, T.J., CoBabe, E., Clement, B., Sliter, W.V., Osburn, C.L. and Longoria, J., 1999. The record of global change in Mid-Cretaceous (Barremian-Albian) sections from the Sierra Madre, northeastern Mexico. Journal of Foraminiferal Research, 29, 418-437. Bralower, T.J. and Thierstein, H.R., 1984. Low productivity and slow deep-water circulation in mid- Cretaceous oceans. Geology, 12, 614-618. Bralower, T.J. and Thierstein, H.R., 1987. Organic-carbon and metal accumulation in Holocene and Mid-Cretaceous marine sediments: palaeoceanographic significance. In: J. Brooks and A.J. Fleet (Editors), Marine petroleum source rocks. Geological Society Special Publications. Geological Society of London, London, UK, pp. 345-369.

Chapter 3 68

Brass, G.W., Southam, J.R. and Peterson, W.H., 1982. Warm saline bottom water in the ancient ocean. Nature, 296, 620-623. Bréhéret, J.G., 1988. Episodes de sédimentation riche en matière organique dans le marnes bleues d'age aptien et albien de la partie pélagique du bassin Vocontien. Bulletin de la Societe Geologique de France, 4, 349-356. Bréhéret, J.G., 1994. The Mid-Cretaceous organic-rich sediments from the Vocontian Zone of the French southeast basin. In: A. Mascle (Editor), Hydrocarbon and petroleum geology of France. The European Association of Petroleum Geoscientists Special Publication. Springer- Verlag, Berlin/Heidelberg, pp. 295-220. Bréhéret, J.G., 1997. L'Aptian et l'Albian de la fosse vocontienne (bordures au bassin): Évolution de la sédimentation et enseignements sur les événements anoxiques. Publications de la Société Géologique du Nord, 25, 614. Brenner, G., 1976. Middle Cretaceous floral provinces and early migration of Angiosperms. In: C.B. Beck (Editor), Origin and early evolution of angiosperms. Columbia University Press, New York, pp. 23-44. Chumakov, N.M., Zharkov, M.A., Herman, A.B., Doludenko, M.P., Kalandadze, N.M., Lebedev, E.L., Ponomarenko, A.G. and Rautian, A.S., 1995. Climatic belts of the mid-Cretaceous time. Stratigraphy and Geological Correlation, 3, 241-260. Dupont, L.M. and Wyputta, U., 2003. Reconstructing pathways of aeolian pollen transport to the marine sediments along the coastline of SW Africa. Quaternary Science Reviews, 22, 157- 174. Erba, E., 1994. Nannofossils and superplumes: The early Aptian "nannoconid crisis". Paleoceanography, 9, 483-501. Erbacher, J., Thurow, J. and Littke, R., 1996. Evolution patterns of radiolaria and organic matter variations: A new approach to identify sea-level changes in Mid-Cretaceous pelagic environments. Geology, 24, 499-502. Frakes, L.A. and Francis, J.E., 1988. A guide to Phanerozoic cold polar climates from high-latitude ice-rafting in the Cretaceous. Nature, 333, 547-549. Francis, J.E. and Frakes, L.A., 1993. Cretaceous climates. In: V.P. Wright (Editor), Sedimentology Review. Blackwell Scientific Publications, Oxford, pp. 17-30. Freeman, K.H. and Hayes, J.M., 1992. Fractionation of carbon isotopes by phytoplankton and estimates of ancient CO2 levels. Global Biogeochemical Cycles, 6, 185-198. Geisen, M., Bollmann, J., Herrle, J.O., Mutterlose, J. and Young, J.R., 1999. Calibration of the random settling technique for calculation of absolute abundances of calcareous nannoplankton. Micropaleontology, 45, 437-442. Gradstein, F.M., Agterberg, F.P., Ogg, J.G., Hardenbol, J., van Veen, P., Thierry, J. and Huang, Z., 1995. A Triassic, Jurassic and Cretaceous time scale. In: W.A. Berggren, D.V. Kent, M.-P. Aubry and J. Hardenbol (Editors), Geochronology, time scales and global stratigraphic correlation. SEPM (Society for Sedimentary Geology), Tulsa, OK, United States, pp. 95-126. Habib, D., 1982. Sedimentary supply origin of Cretaceous black shales. In: S.O. Schlanger and M.B. Cita (Editors), Nature and origin of Cretaceous carbon-rich facies. Academic Press, New York, NY, pp. 113-127. Hallam, A., 1985. A review of Mesozoic climate. Journal of the Geological Society of London, 142, 433-445. Hay, W.W., DeConto, R.M., Wold, C.N., Wilson, K.M., Voigt, S., Schulz, M., Wold, A.R., Dullo, W.C., Ronov, A.B., Balukhovsky, A.N. and Soeding, E., 1999. Alternative global Cretaceous paleogeography, Evolution of the Cretaceous ocean-climate system. Geological Society of America (GSA), Boulder, pp. 1-47. Herrle, J.O., Kößler, P., Friedrich, O., Erlenkeuser, H. and Hemleben, C., 2004. High-resolution carbon isotope records of the Aptian to Lower Albian from SE France and the Mazagan Plateau (DSPD site 545): a stratigraphic tool for paleocanographic and paleobiologic reconstruction. Earth and Planetary Science Letters, 218, 149-161.

Chapter 3 69

Herrle, J.O. and Mutterlose, J., 2003. Calcareous nannofossils from the Aptian-Lower Albian of southeast France: Palaeoecological and biostratigraphic implications. Cretaceous Research, 24, 1-22. Herrle, J.O., Pross, J., Friedrich, O. and Hemleben, C., 2003a. Short-term environmental changes in the Cretaceous Tethyan Ocean: Micropalaeontological evidence from the Early Albian Oceanic Anoxic Event 1b. Terra Nova, 15, 14-19. Herrle, J.O., Pross, J., Friedrich, O., Kössler, P. and Hemleben, C., 2003b. Forcing mechanisms for Mid-Cretaceous black shale formation; evidence from the upper Aptian and lower Albian of the Vocontian Basin (SE France). Palaeogeography, Palaeoclimatology, Palaeoecology, 190, 399-426. Heusser, L. and Balsam, W.L., 1977. Pollen distribution in the Northeast Pacific Ocean. Quaternary Research, 7, 45-62. Hochuli, P.A., 1981. North Gondwanan floral elements in Lower to Middle Cretaceous sediments of the Southern Alps (Southern Switzerland, Northern Italy). Review of Palaeobotany and Palynology, 35, 337-358. Hochuli, P.A., Menegatti, A.P., Weissert, H., Riva, A., Erba, E. and Premoli Silva, I., 1999. Episodes of high productivity and cooling in the early Aptian Alpine Tethys. Geology, 27, 657-660. Huber, B.T., Hodell, D.A. and Hamilton, C.P., 1995. Middle-Late Cretaceous climate of the southern high latitudes: Stable isotopic evidence for minimal equator-to-pole thermal gradients. Geological Society of America Bulletin, 107, 1164-1191. Jenkyns, H.C., 1980. Cretaceous anoxic events: from continents to oceans. Journal of the Geological Society of London, 137, 171-188. Jenkyns, H.C., 1999. Mesozoic anoxic events and palaeoclimate. Zentralblatt für Geologie und Paläontologie, 1997, 943-949. Kemper, E., 1987. Das Klima der Kreide-Zeit. Geologisches Jahrbuch, Reihe A, 96, 5-185. Kössler, P., Herrle, J.O., Appel, E., Erbacher, J. and Hemleben, C., 2001. Magnetic records of climatic cycles from mid-cretaceous hemipelagic sediments of the Vocontian Basin, SE France. Cretaceous Research, 22, 321-331. Kuypers, M.M.M., Pancost, R.D., Nijenhuis, I.A. and Sinninghe Damste, J.S., 2002. Enhanced productivity led to increased organic carbon burial in the euxinic North Atlantic basin during the late Cenomanian oceanic anoxic event. Paleoceanography, 17, 1-13. Kuypers, M.M.M., Pancost, R.D. and Sinninghe, D.J.S., 1999. A large and abrupt fall in atmospheric CO2 concentration during Cretaceous times. Nature, 399, 342-345. Leckie, R.M., Bralower, T.J. and Cashman, R., 2002. Oceanic anoxic events and plankton evolution: Biotic response to tectonic forcing during the mid-Cretaceous. Paleoceanography, 17, 13-1 - 13-29. Luciani, V., Cobianchi, M. and Jenkyns, H.C., 2001. Biotic and geochemical response to anoxic events: The Aptian pelagic succession of the Gargano promontory (Southern Italy). Geological Magazine, 138, 277-298. Masure, E., Raynaud, J.-F., Pons, D. and De Reneville, P., 1998. Palynologie du stratotype historique de l'Alptien inférieur dans la région de Cassis - La Bédoule (SE France). Géologie Méditerranéenne, 25, 263-287. Melia, M.B., 1984. The distribution and relationship between palynomorphs in aerosols and deep-sea sediments off the coast of Northwest Africa. Marine Geology, 58, 345-371. Menegatti, A.P., Weissert, H., Brown, R.S., Tyson, R.V., Farrimond, P., Strasser, A. and Caron, M., 1998. High-resolution δ13C stratigraphy through the early Aptian "Livello Selli" of the Alpine Tethys. Paleoceanography, 13, 530-545. Mohr, B.A.R., 1989. New palynological information on the age and environment of Late Jurassic and Early Cretaceous vertebrate localities of the Iberian Peninsula (eastern Spain and Portugal). Berliner Geowissenschaftliche Abhandlungen (A), 106, 291-301. Moullade, M., 1966. Etude stratigraphique et micropaléontologique du Crétacé inférieur de la "Fosse Vocontienne". Documents des Laboratoires de Géologie de la Faculté des Sciences de Lyon, 15, 1-369.

Chapter 3 70

Pedersen, T.F. and Calvert, S.E., 1990. Anoxia vs. productivity: What controls the formation of organic- carbon-rich sediments and sedimentary rocks? American Association of Petroleum Geologists Bulletin, 74, 454-466. Peters, K.E. and Moldowan, J.M., 1993. The Biomarker Guide. Prentice-Hall, Englewood Cliffs, NJ, 363 pp pp. Premoli Silva, I., Erba, E., Salvini, G., Locatelli, C. and Verga, D., 1999. Biotic changes in Cretaceous oceanic anoxic events of the Tethys. Journal of Foraminiferal Research, 29, 352-370. Price, G.D., 2003. New constraints upon isotope variation during the early Cretaceous (Barremian- Cenomanian) from the Pacific Ocean. Geological Magazine, 140, 513-522. Roth, P.H. and Bowdler, J.L., 1981. Middle Cretaceous calcareous nannoplankton biogeography and oceanography of the Atlantic ocean. Society of Economic Palaeontologists and Mineralogists, Special Publications, 32, 517-546. Schlanger, S.O. and Jenkyns, H.C., 1976. Cretaceous oceanic anoxic events: Causes and consequences. Geologie en Mijnbouw, 55, 179-184. Spicer, R.A. and Parrish, J.T., 1986. Paleobotanical evidence for cool North Polar climates in Middle Cretaceous (Albian-Cenomanian) time. Geology, 14, 703-706. Stein, R., Rullkötter, J. and Welte, D.H., 1986. Accumulation of organic-carbon-rich sediments in the Late Jurassic and Cretaceous Atlantic Ocean: a synthesis. Chemical Geology, 56, 1-32. Stoll, H.M. and Schrag, D.P., 1996. Evidence for glacial control of rapid sea level changes in the Early Cretaceous. Science, 272, 1771-1774. Strasser, A., Caron, M. and Gjermeni, M., 2001. The Aptian, Albian and Cenomanian of Roter Sattel, Romandes Prealps, Switzerland: A high-resolution record of oceanographic changes. Cretaceous Research, 22, 173-199. Thierstein, H.R., 1980. Selective dissolution of Late Cretaceous and Earliest Tertiary calcareous nannofossils: experimental evidence. Cretaceous Research, 2, 165-176. Torricelli, S., 2000. Lower Cretaceous dinoflagellate cyst and acritarch stratigraphy of the Cismon APTICORE (Southern Alps, Italy). Review of Palaeobotany and Palynology, 108, 213-266. Traverse, A., 1988. Paleopalynology. Unwin Hyman, Boston, 600 pp. Traverse, A., 1994. Sedimentation of Organic Particles. Cambridge Universtiy Press, Cambridge, 1- 544 pp. Tribovillard, N.P. and Gorin, G.E., 1991. Organic facies of the early Albian Niveau Paquier, a key black shales horizon of the Marnes Bleues Formation in the Vocontian Trough (Subalpine Ranges, SE France). Palaeogeography, Palaeoclimatology, Palaeoecology, 85, 227-237. Tyson, R.V., 1995. Sedimentary organic matter. Chapman & Hall, London, 615 pp. Vakhrameyev, V.A., 1978. The climates of the Northern Hemisphere in the Cretaceous in the light of paleobotanical data. Paleontological Journal, 12, 143-154. Vakhrameyev, V.A., 1982. Classopollis pollen as an indicator of Jurassic and Cretaceous climate. International Geology Review, 24, 1190-1196. Vakhrameyev, V.A., 1991. Jurassic and Cretaceous floras and climates of the Earth. Cambridge University Press, Cambridge, 318 pp. Van Konijnenburg - Van Cittert, J.H.A. and Van der Burgh, J., 1989. The flora from the Kimmeridgian (Upper Jurassic) of Culgower, Sutherland, Scotland. Review of Palaeobotany and Palynology, 61, 1-51. Watson, J., 1988. The Cheirolepidiaceae. In: C.B. Beck (Editor), Origin and evolution of the Gymnosperms. Columbia University Press, New York, pp. 382-447. Weissert, H. and Bréhéret, J.G., 1991. A carbonate-isotope record from Aptian-Albian sediments of the Vocontian Trough (SE France). Bulletin de la Societe Geologique de France, 162, 1133- 1140. Weissert, H. and Lini, A., 1991. Ice age interludes during the time of Cretaceous greenhouse climate? In: D.W. Müller, J.A. McKenzie and H. Weissert (Editors), Controversies in modern geology: Evolution of geological theories in sedimentology, earth history and tectonics. Academic Press, London, UK, pp. 173-191.

Chapter 3 71

Weissert, H., Lini, A., Foellmi, K.B. and Kuhn, O., 1998. Correlation of Early Cretaceous carbon isotope stratigraphy and platform drowning events: A possible link? Palaeogeography, Palaeoclimatology, Palaeoecology, 137, 189-203. Wilpshaar, M. and Leereveld, H., 1994. Palaeoenvironmental change in the Early Cretaceous Vocontian Basin (SE France) reflected by dinoflagellate cysts. Review of Palaeobotany and Palynology, 84, 121-128. Wilson, P.A. and Norris, R.D., 2001. Warm tropical ocean surface and global anoxia during the mid- Cretaceous period. Nature, 412, 425-428. Wissler, L., Funk, H. and Weissert, H., 2003. Response of Early Cretaceous carbonate platforms to changes in atmospheric carbon dioxide levels. Palaeogeography, Palaeoclimatology, Palaeoecology, 200, 187-205.

Chapter 3 72

Chapter 4 73

Chapter 4

Terrestrial carbon-isotope records from coastal deposits (Algarve, Portugal):

A tool for chemostratigraphic correlation on an intrabasinal and global scale*

Abstract

The carbon-isotope signature of terrestrial organic matter (OM) offers a valuable tool to develop stratigraphic correlations for near-shore deposits. A mid-Cretaceous coastal succession of the western Algarve Basin, Portugal, displays a marked negative δ13C excursion ranging from -21.2‰ to -27.8‰ in the Early Aptian followed by two shifts towards higher values (up to -19.3‰) during the Early and Late Aptian, respectively. The dominance of cuticle and leaf debris in the bulk OM fraction is confirmed by optical studies, Rock-Eval pyrolysis and by comparison with the δ13C signature of four different types of fossilized land- 13 plant particles. Correlation of two terrestrial δ Cbulk OM records from different study sites leads to a significant enhancement of the intrabasinal stratigraphic correlation within the Algarve Basin. Three prominent excursions in the Portuguese records can be correlated with existing δ13C curves from pelagic and terrestrial environments. The general carbon-isotope pattern is superimposed by small-scale fluctuations which can be explained by compositional variations within the OM.

Keywords: carbon-13, phytoclasts, chemostratigraphy, Aptian, terrestrial environment

* published as: Heimhofer, U., Hochuli, P. A., Burla, S., Andersen, N. and Weissert, H. (2003). Terrestrial carbon-isotope records from coastal deposits (Algarve, Portugal): A tool for chemostratigraphic correlation on an intrabasinal and global scale. Terra Nova, 15, 8-13

Chapter 4 74

1. Introduction

Several major carbon-isotope excursions, initially reported from marine carbonates (Ccarb) and accompanying marine organic carbon (Corg) have been recognized recently in plant material of terrestrial origin. In these studies, different types of vascular land-plant debris served as an isotopic substrate, including fossil wood (Gröcke et al., 1999), jet, coal and charcoal (Hesselbo et al., 2000), cuticle and vitrinite (Jahren et al., 2001) or bulk terrestrial OM (Ando et al., 2002; Hasegawa, 1997). Even though the δ13C composition of land-plants is affected by different ecophysiological, taphonomic and diagenetic effects, the fossilized tissues serve to trace changes in the carbon-isotope composition of the ocean–atmosphere system through earth history. One of the best-studied carbon-isotope records covers the Aptian stage (121-112 Ma), a time of major perturbations of the global carbon cycle as documented in the widespread deposition of organic-carbon rich shales in the world oceans (Arthur et al., 1990; Bralower et al., 1994) occurrence of marine biocalcification crises (Erba, 1994; Weissert et al., 1998) and accompanying biological turnover (Erbacher et al., 1996; Hochuli et al., 1999). The corresponding carbon isotope records are marked by several pronounced negative and positive excursions with an overall amplitude of ~4.0‰ in marine carbonates and up to

~7.5‰ in marine Corg. The diagnostic isotope pattern has been recognized and described in detail from pelagic and hemipelagic successions (Bralower et al., 1999; Menegatti et al., 1998; Weissert and Breheret, 1991) as well as from time-equivalent shallow water environments (Ferreri et al., 1997; Jenkyns, 1995). More recently, similar δ13C variations measured in terrestrial plant OM have been correlated with marine isotope records (Ando et al., 2002; Gröcke et al., 1999) and emphasize the close linkage between the oceanic and atmospheric carbon reservoirs.

In contrast to marine Ccarb and Corg, the carbon-isotope geochemistry of terrestrial OM has not yet been widely applied as a tool for high-resolution stratigraphy and correlation between different depositional environments (Hesselbo et al., 2000; Hesselbo et al., 2002). This application is of special interest for near-shore deposits, which often lack an adequate stratigraphic resolution due to the rare occurrence or absence of reliable biostratigraphic markers. In this study, we investigate the organic carbon-isotope geochemistry of an Early Cretaceous coastal succession. We use the isotopic signature of a variety of vascular land-plant materials

Chapter 4 75 and bulk terrestrial OM to demonstrate that near-shore successions can be accurately dated with organic carbon-isotope records of terrestrial origin.

2. Study sites

Two sections from the Algarve region (southern Portugal) have been chosen as terrestrial archives spanning the Aptian time window. Both study sites (Luz and Burgau) are located within the western part of the Algarve Basin (Fig. 1) and have been described in detail by Rey (1983; 1986) from a sedimentological and biostratigraphic perspective. The sedimentary succession consists mainly of varicolored clays and marls with some intercalated siltstone and limestone beds (Luz Marls Formation). These sediments were deposited in a shallow lagoonal to brackish marsh environment with only minor open-marine episodes. The Luz Marls gradually evolve into a carbonate-dominated tidal flat setting, documented in the deposition of thick-bedded shallow-water limestone and calcareous marls (Porto de Mos Formation). Both sections represent pronounced near-shore depositional settings. Evidence for sedimentary gaps is restricted to the occurrence of several hardgrounds in the upper carbonate-dominated unit and to a depositional discontinuity at the base of a graded limestone bed within the Luz Marls. The uniform sedimentary setting and the occurrence of characteristic depositional patterns allow an accurate lithostratigraphic correlation of the two sections over a distance of ~ 6.5 km.

010 km r v e N a Iberia g l Portimão A Lagos Algarve Burgau basin Luz section Sagres 37°N Burgau section

9°W 8°30'W

Fig. 1: Location map of the western Algarve basin on the Iberian Peninsula. Studied sections are marked with an arrow.

Due to the lack of common index fossils, biostratigraphy has been based on dinoflagellate cysts (Berthou and Leereveld, 1990), benthic foraminifera and calcareous algae (Rey, 1983;

Chapter 4 76

Rey, 1986). In combination with our new palynological data (will be published elsewhere) these results suggest an Early Aptian age for the lower, and a Late Aptian age for the upper part of the Luz Marls Formation.

3. Methods

Closely spaced samples (~ 1 m to 2 m) from Luz and Burgau were measured for the carbon- 13 isotope composition of bulk OM. To avoid possible diagenetic alteration effects of the δ Cbulk signature, reddish and purple colored horizons were excluded from this analysis. For bulk OM determinations, 400 mg of each sample was treated twice with 1 N HCl for 24 h to remove the carbonate carbon. 1-20 mg of the residue was analyzed via combustion for δ13C in a CNS Elemental Analyzer (Carlo Erba Instruments) connected to an isotope ratio mass spectrometer (Optima/Micromass). Carbon-isotope ratios were expressed in the standard δ notation in per mil (‰) relative to the international VPDB isotope standard. The δ13C values were calibrated against a laboratory internal standard (Atropina; δ13C = -28.48‰) and an international standard (NBS 22; δ13C = -29.74‰); analytical reproducibility was ±0.2‰. Inorganic and total organic carbon content (IC/TOC) was measured on a UIC CM 5012 Coulomat. To assess the origin of the OM as well as the compositional variations within, bulk parameter measurements including visual kerogen analysis and Rock-Eval pyrolysis were combined with the carbon-isotope analysis of various types of vascular land-plant particles. Following the method of Jahren et al. (2001) we compared the land-plant δ13C signature with that of the bulk OM signal to determine its main components. If no macroscopic fossil wood fragments were available, the sample was acid macerated (24 h with 3 N HCl), rinsed and sieved (>62 µm). Following this treatment, the isolated phytoclasts were picked by hand under the microscope. Four types of different land-plant particles have been distinguished including charcoal, lignite, translucent cuticle and opaque leaf fragments. All phytoclasts were measured for their carbon-isotope composition via combustion using the same procedure as for bulk OM. If possible, repeated measurements were carried out and the standard error of the means was calculated.

4. Characterization of the sedimentary organic matter

Both studied sections represent siliciclastic-dominated coastal environments with the sedimentary OM strongly dominated by terrestrial material. Total organic carbon (TOC)

Chapter 4 77 content (dry wt. %) of the sediment varies between 0.1% and 0.9% throughout the entire succession with a mean value of 0.2%. Palynofacies analysis displays a high abundance of opaque phytoclasts, cuticle fragments, spores and pollen grains. These results are supported by low HI values (< 150 mg HC/g TOC) indicating a strong terrestrial contribution to the sedimentary OM. The δ13C values of the phytoclasts were compared to the bulk OM signal obtained from the same horizons (Fig. 2). The isotopic composition of leaves (mean of -23.3‰) and translucent cuticle (mean of –23.1‰) is very similar to the average bulk OM signature (mean of -23.4‰), although the variability in the phytoclasts is larger (1.6‰). In contrast to this, charcoal (mean of –20.8‰) and lignite (mean of –21.4‰) show a mean offset of 1.6‰ in the Luz Marls Formation and of 2.8‰ in the Porto de Mos Formation. In comparison to bulk OM both particle types display similar shifts throughout the section. Variations in the isotopic offset between bulk OM and individual phytoclast types can occur due to changes in the proportion of the different phytoclasts or result from additional OM to the bulk fraction from a different source, most likely marine. The congruence of the bulk OM isotope signature and the cuticle/leaf particles clearly indicates that the measured OM is predominantly composed of foliage debris of continental origin. Therefore its δ13C signature can be interpreted as to represent a terrestrial signal. Furthermore, the consistency of the isotope shifts in bulk OM and land-plant particles demonstrates that fluctuations in the bulk terrestrial OM record are not solely controlled by variations in the mixing ratio of terrestrial and marine OM. Fossilized plant cuticle has been proposed as an ideal substrate for carbon-isotopic studies due to its high resistance to decay and degradation processes (Arens et al., 2000; Upchurch et al., 1997). Evidence for the primary nature of the measured δ13C phytoclast signature is given by the consistent isotopic difference of ~2.0‰ between translucent cuticle and lignite. A similar depletion in 13C of about 2.5‰ to 3.5‰ between cuticle and leaves relative to whole wood plant carbon has been reported from extant as well as from fossil plants (Leavitt and Long, 1982; Upchurch et al., 1997) suggesting an insignificant diagenetic alteration of the isotopic signal. Thermally unaltered conditions for the sedimentary OM are indicated by unchanged coloring of the palynomorphs (TAI < 2), strong UV fluorescence of the amorphous OM fraction and 13 low Tmax values (mean of 424.5°C). The absence of any significant correlation between δ C values and CaCO3- or TOC-content of the samples indicates the independence of the carbon- isotope signature from lithological variations.

Chapter 4 78

height (m) charcoal lignite cuticules leafs

21 0

19 0

17 0 Porto de Mos Fm. Lower Albian Lower 15 0

13 0

11 0 Upper Aptian

90 Luz Marls

70

50 bulk

Lower Aptian phytoclasts 30 -27 -25 -23 -21 -19 -27 -25 -23 -21 -19 -27 -25 -23 -21 -19 -27 -25 -23 -21 -19 δ13 δ13 δ13 δ13 Corg (%0 VPDB) Corg (%0 VPDB) Corg (%0 VPDB) Corg (%0 VPDB)

13 Fig. 2: δ Corg measurements of different types of land-plant particles (closed symbols) and the 13 corresponding bulk OM signature (open symbols) from the same horizon. Error bars of δ Corg values represent standard errors of means of repeated measurements.

5. Intrabasinal chemostratigraphic correlation

13 In order to test the terrestrial δ Cbulk data for its consistency as well as for its potential as a chemostratigraphic correlation tool, the carbon-isotope records of the Luz and Burgau 13 sections have been compared in detail (Fig. 3). Even though the δ Corg record of the Luz section is rather noisy and records an overall variation of ~8.5‰, pronounced shifts in the 13 magnitude of 5.0‰ to 7.0‰ can be observed. The most significant features of the δ Cbulk OM curve are an isotopic minimum with values down to –27.8‰ (from 17 m to 37 m) in the Early Aptian, followed by two prominent and abrupt shifts towards higher values (–19.4‰ at 37 m; -19.3‰ at 120 m) in the Early and Late Aptian, respectively. Furthermore, intervals with strong δ13C variability (70 m – 78 m, variation of ~4.5‰; 120 m – 157 m, variation of ~5.1‰) and intervals displaying more constant values (78 m - 120 m, variation of ~2.4‰) can be recognized. Comparison with the Burgau carbon-isotope record (overall variation of

Chapter 4 79

~6.0‰) reveals that both curves do not only exhibit a similar shape with its distinct isotopic shifts, but also show correspondence in the small-scale fluctuations. The obvious congruence 13 of the two records is furthermore supported by the similarity of the averaged δ Corg values of -22.8‰ (Luz) and -23.2‰ (Burgau). Based on the lithostratigraphic framework, the correlation of large- and small-scale isotope excursions enables the establishment of 10 chemostratigraphic segments (I to X) resulting in a significant enhancement of the intra- basinal stratigraphic correlation. The well preserved OM and the occurrence of a similar isotopic pattern at two separate study sites rule out a diagenetic control on the δ13C curve. Our preliminary palynological results provide no evidence for major floral turnovers or input of a group of plants with exceptional carbon-isotope compositions, which could explain abrupt shifts in the Aptian terrestrial δ13C record. Even though the bulk terrestrial OM in the Luz section is predominatly composed of cuticle and leaf debris, occasional input of isotopically less negative lignite and charcoal particles results in δ13C shifts, which contribute to the small-scale fluctuations occurring throughout the record. The strong δ13C variability in segment VII, IX and X of the Luz record is interpreted to reflect compositional changes of the bulk OM due to fluctuations in the ratio of marine to terrestrial OM. Horizons of purely terrestrial material alternate with intervals containing a significant amount of isotopically light amorphous OM of presumably marine origin. These alternations result in abrupt and brief carbon isotope shifts in the bulk OM record. Despite a variety of factors contributing to small-scale fluctuations of the terrestrial δ13C record, the overall trend of the curve with its prominent excursions can not be explained solely by compositional variations of the OM, changes in floral assemblage or diagenetic alteration.

6. Global significance

Based on the palynostratigraphic framework, the Portuguese carbon-isotope profiles are 13 compared to an existing terrestrial δ Cwood curve (Gröcke et al., 1999) and to a marine 13 δ Ccarb reference record, including data of Erba et al. (1999) and Bralower et al.(1999). The different curves exhibit essentially the same characteristic Aptian carbon-isotope pattern with its distinctive anomalies (Fig. 4). A first marked negative δ13C excursion (1) in the marine

Chapter 4 80

E Algarve basin W ~ 6.5 km Luz -28 -26 -24 -22 -20 -18 formation stage Spores & Pollen height (m) Dinocysts ment g

150 Albian Burgau 140

-28 -26 -24 -22 -20 hemostrat. se c 100 X 130

90 IX 120

80 110 VIII 70 100

60 90 VII 505 80 VI

40 V 70 IV 30 60 III 20 50 II 10 40 I

Lower AptianLower Upper Aptian 0 m 30 -28 -26 -24 -22 -20

. δ 13C bulk OM (‰ VPDB) sp

20 . dividus cf . jardinus aff

10 Marls Luz Lower Marls Luz Upper limestone marl silt Tricolpits vulgaris Tricolpits Dichastopollenites Brenneripollis reticulatus Brenneripollis Afropollis jardinus Afropollis peroreticulatus Brenneripollis Brenneripollis 0 Hystrichosphaeridium arborispinum corollum Protoellipsodinium Dinopterygium cladoides pariata Muderongia Subtilisphaera cheit Cyclonephelium paucimarginatum staurota Muderongia Odontochitina ancala tenuiceras Tehamadinium claystone (<25% CaCO3) -28 -26 -24 -22 -20 -18 correlation of marker beds δ 13C bulk OM (‰ VPDB)

13 Fig. 3: Palynostratigraphy, lithological logs and terrestrial δ Cbulk OM data of two sections covering the Aptian Luz Marls Formation of the western Algarve Basin, Portugal. Dotted lines correspond to the lithostratigraphic framework. Shaded bars indicate the chemostratigraphic correlation. Only palynomorphs with stratigraphic significance are displayed. The occurrence of Ctenidodinium elegantulum, Rhynchodiniopsis aptiana and Pseudoceratium securigerum 30 m below the base of the displayed section indicate an Early Bedoulian age for the lowermost part of the Luz Marls Formation. reference record occurs in the Lowermost Aptian nannofossil Zone NC6, base of the Leupoldina cabri planktic foraminiferal Zone. The negative excursion is followed by two prominent shifts towards more positive values (2) at the transition from nannofossil Zone NC6 to NC7, upper Leupoldina cabri planktic foraminiferal Zone and (3) in the uppermost

Chapter 4 81

Composite Isle of Wight Algarve marine record (UK) (Portugal) δ13 Ccarb (‰ VPDB) 13 δ Cwood (‰ VPDB) δ13Cbulk OM (‰ VPDB) Zonation

2.0 3.0 4.0 5.0 height (m) Pk Foram. Zone stage stage stage Nannofos. Zone height (m) Ammonite -30 -26 -22 -18 -28 -26 -24 -22 -20 160 Alb.

Albian 300 Alb.

140 jacobi 250 120 nutfield- H. plan. - T. bej. H. plan. - T. iensis Upper 200 100

Aptian NC7 NC8 G. alg. martinioides 80 150 bowerbanki G. ferr. deshayesi 60 100 forbesi 40 Lower L. cabri ?

Aptian 50 Lower Aptianfissicostatus Upper Aptian NC6 Lower Aptian20 Upper Aptian

Barr. G. blowi 0 0 Barr. NC5

13 Fig. 4: Tentative correlation of mid-Cretaceous terrestrial δ Corg records from the Isle of Wight, United Kingdom (Gröcke et al., 1999) and from the Portuguese Algarve Basin (this study). In addition, both terrestrial records are correlated with a composite marine reference curve, based on 13 δ Ccarb measurements from Mexican and Tethyan sites (Bralower et al., 1999; Erba et al., 1999). A 3 point moving average was applied to the Algarve record to compensate for the noisiness of the curve due to compositional variations in the measured bulk terrestrial OM. Shaded areas illustrate the correlation between the records. Terrestrial curves are plotted against thickness (m) in different scales. nannofossil Zone NC7, Ticinella bejaouaensis planktic foraminiferal Zone. These marked isotope excursions can be correlated with an isotopic minimum (segment II) and with two shifts towards more positive values (segment III and IX) in our terrestrial record. The occurrence of a distinct Aptian isotope pattern in the Portuguese record facilitates a well- defined chemostratigraphic correlation with existing marine and terrestrial δ13C curves and results in a significant increase in the stratigraphic resolution of these near-shore deposits.

7. Conclusions

Carbon-isotope studies on terrestrial OM obtained form near-shore depositional settings hold a strong potential to serve as continental high-resolution records during earth history. Our results demonstrate that despite a multitude of environmental and diagenetic factors affecting

Chapter 4 82 the carbon-isotope signature of bulk terrestrial OM in coastal depositional systems, the overall trend of the δ13C record can serve as a reliable chemostratigraphic correlation tool. This is confirmed by an intrabasinal correlation of coastal deposits using the δ13C signature of continental-derived bulk OM. Comparison with existing mid-Cretaceous carbon-isotope curves results in a significant increase of the stratigraphic resolution of the Portuguese near- shore succession and points to the global significance of the terrestrial δ13C record. The higher resolution will offer the opportunity to study the response of a near-shore sedimentary system to major perturbations of the ocean-atmosphere-biosphere system.

Acknowledgements

We thank P. Steinmann from the University Neuchâtel for Rock-Eval pyrolysis determinations; J. Dinis from Coimbra University and R. Gonzales from Algarve University for field assistance. This manuscript was significantly improved thanks to suggestions and reviews by D.R. Gröcke and an anonymous reader. Financial support from ETH-Project TH- 34./99-4 is greatfully acknowledged.

References

Ando, A., Kakegawa, T., Takashima, R. and Saito, T., 2002. New perspective on Aptian carbon isotope stratigraphy: Data from δ13C records of terrestrial organic matter. Geology, 30, 227- 230. Arens, N.C., Jahren, A.H. and Amundson, R., 2000. Can C3 plants faithfully record the carbon isotopic composition of atmospheric carbon dioxide? Paleobiology, 26, 137-164. Arthur, M.A., Jenkyns, H.C., Brumsack, H.-J. and Schlanger, S.O., 1990. Stratigraphy, geochemistry and paleoceanography of organic-carbon rich Cretaceous sequences. In: R.N. Ginsburg and B. Beaudoin (Editors), Cretaceous Resources, Events and Rhythms. NATO ASI Series C. Kluwer Academic Publishers, MasDordrecht, pp. 75-119. Berthou, P.Y. and Leereveld, H., 1990. Stratigraphic implications of palynological studies on Berriasian to Albian deposits from western and southern Portugal. Review of Palaeobotany and Palynology, 66, 313-344. Bralower, T.J. et al., 1994. Timing and paleoceanography of oceanic dysoxia/ anoxia in the late Barremian to early Aptian (Early Cretaceous). Palaios, 9, 335-369. Bralower, T.J. et al., 1999. The record of global change in Mid-Cretaceous (Barremian-Albian) sections from the Sierra Madre, northeastern Mexico. Journal of Foraminiferal Research, 29, 418-437. Erba, E., 1994. Nannofossils and superplumes: The early Aptian "nannoconid crisis". Paleoceanography, 9, 483-501. Erba, E. et al., 1999. Integrated stratigraphy of the Cismon Apticore (southern Alps, Italy): A "reference section" for the Barremian-Aptian interval at low latitudes. Journal of Foraminiferal Research, 29, 371-391.

Chapter 4 83

Erbacher, J., Thurow, J. and Littke, R., 1996. Evolution patterns of radiolaria and organic matter variations: A new approach to identify sea-level changes in Mid-Cretaceous pelagic environments. Geology, 24, 499-502. Ferreri, V., Weissert, H., D'Argenio, B. and Buonocunto, F.P., 1997. Carbon isotope stratigraphy; a tool for basin to carbonate platform correlation. Terra Nova, 9, 57-61. Gröcke, D.R., Hesselbo, S.P. and Jenkyns, H.C., 1999. Carbon-isotope composition of Lower Cretaceous fossil wood: Ocean-atmosphere chemistry and relation to sea-level change. Geology, 27, 155-158. Hasegawa, T., 1997. Cenomanian-Turonian carbon isotope events recorded in terrestrial organic matter from northern Japan. Palaeogeography, Palaeoclimatology, Palaeoecology, 130, 251- 273. Hesselbo, S.P. et al., 2000. Massive dissociation of gas hydrate during a Jurassic oceanic anoxic event. Nature, 406, 392-395. Hesselbo, S.P., Robinson, S.A., Surlyk, F. and Piasecki, S., 2002. Terrestrial and marine extinction at the Triassic-Jurassic boundary synchronized with major carbon-cycle perturbations: a link to initiation of massive volcanism? Geology, 30, 251-254. Hochuli, P.A., Menegatti, A.P., Riva, A., Weissert, H. and Erba, E., 1999. High-productivity and cooling episodes in the Early Aptian Alpine Tethys European Union of Geosciences conference abstracts; EUG 10, European Union of Geosciences conference; EUG 10. Strasbourg, France. March 28-April 1, 1999. Journal of Conference Abstracts. Cambridge Publications. Cambridge, United Kingdom. 1999., pp. 219. Jahren, A.H., Arens, N.C., Sarmiento, G., Guerrero, J. and Amundson, R., 2001. Terrestrial record of methane hydrate dissociation in the Early Cretaceous. Geology, 29, 159-162. Jenkyns, H.C., 1995. Carbon-isotope stratigraphy and paleoceanographic significance of the Lower Cretaceous shallow-water carbonates of Resolution Guyot, Mid-Pacific Mountains. In: E.L. Winterer, W.W. Sager, J.V. Firth and J.M. Sinton (Editors), Proceedings of the Ocean Drilling Program, Scientific Results. Proceedings of the Ocean Drilling Program, Scientific Results. Texas A & M University, Ocean Drilling Program, College Station, TX, United States, pp. 99- 104. Leavitt, S.W. and Long, A., 1982. Evidence for 13C/12C fractionation between tree leaves and wood. Nature, 298, 742-744. Menegatti, A.P. et al., 1998. High-resolution δ13C stratigraphy through the early Aptian "Livello Selli" of the Alpine Tethys. Paleoceanography, 13, 530-545. Rey, J., 1983. Le Crétacé de l'Algarve: Essai de Synthèse. Comunicações dos Serviços Geológicos de Portugal, 69, 87-101. Rey, J., 1986. Micropaleontological assemblages, paleoenvironments and sedimentary evolution of Cretaceous deposits in the Algarve (southern Portugal). Palaeogeography, Palaeoclimatology, Palaeoecology, 55, 233-246. Upchurch, G.R., Marino, B.D., Mone, W.E. and McElroy, M.B., 1997. Carbon isotope ratios in extant and fossil plant cuticule. American Journal of Botany, 84, 143-144. Weissert, H. and Breheret, J.G., 1991. A carbonate-isotope record from Aptian-Albian sediments of the Vocontian Trough (SE France). Bulletin de la Societe Geologique de France, 162, 1133- 1140. Weissert, H., Lini, A., Foellmi, K.B. and Kuhn, O., 1998. Correlation of Early Cretaceous carbon isotope stratigraphy and platform drowning events: A possible link? Palaeogeography, Palaeoclimatology, Palaeoecology, 137, 189-203.

Chapter 4 84

Chapter 5 85

Chapter 5

A well-dated and continuous early angiosperm pollen record from mid-Cretaceous

coastal deposits (Lusitanian and Algarve Basins, Portugal):

Implications for the timing of the early angiosperm radiation

Abstract

Detailed and continuous palynological records from two well-dated successions in the Portuguese Algarve and Lusitanian Basins are presented, which document the diversification of early angiosperm pollen during the Barremian to Albian time interval. Based on dinoflagellate cysts biostratigraphy, an accurate stratigraphic framework has been established for the studied near-shore deposits resulting in distinct changes of the stratigraphic position of individual units. The qualitative and quantitative analysis of the palynofloras of the two sections revealed a total of 60 different types of angiosperm pollen. Most of them (51 taxa) are monoaperturate grains of magnoliid or monocot affinity. In both records eudicots, represented by various tricolpate taxa (9 taxa), are restricted to the post-Aptian part of the sections. Angiosperm pollen display a distinct increase in both, diversity (up to 18 taxa per sample) and relative abundance (up to 12 %) between the Late Barremian and Middle Albian. Comparison with published studies shows strong similarities with regard to floral composition and the timing of first appearances of particular angiosperm pollen forms. Our results ask for a new age interpretation of the well-known angiosperm mesofossil floras from the Portuguese Estremadura region which have been interpreted as Barremian or possibly Aptian in age. Several lines of evidence, including sequence- and biostratigraphy as well as palynology, indicate a post-Aptian age for these assemblages (incl. the Famalicão, Buarcos and Vale de Agua mesofloras), hence demonstrating a major radiation phase during the Early Albian.

Key words: early angiosperms; radiation; mid-Cretaceous; palynology; biostratigraphy; Portugal Chapter 5 86

1. Introduction

The mid-Cretaceous diversification of angiosperms marks the profound change from Mesozoic floras dominated by ferns, conifers and cycads to the modern, angiosperm- dominated ecosystems of the Cenozoic era (e.g. Crane et al., 1995; Lidgard and Crane, 1988; Willis and McElwain, 2002). Evidence for early angiosperms in the fossil record has been essentially obtained from the analysis of fossil palynofloras from continental to shallow-water deposits. Pollen grains of unambiguous angiosperm origin have been reported from Barremian strata from various localities including equatorial and northern Africa (e.g. Doyle et al., 1977; Gübeli et al., 1984; Penny, 1986; Schrank and Mahmoud, 2002) as well as northwestern Europe (Hughes et al., 1979; Hughes and McDougall, 1990). These early assemblages consist of monoaperturate pollen types with reticulate-semitectate or columellate-tectate wall structure and display strong similarity to pollen of extant magnoliids or monocotyledons. The occurrence of presumed eudicots is documented by the appearance of triaperturate pollen grains from younger, post-Barremian deposits (e.g. Brenner, 1963; Brenner, 1996; Doyle and Robbins, 1977; Penny, 1986). Quantitative analyses of genera and species richness of numerous Cretaceous macrofossil floras display a step-wise increase of angiosperms diversity during the mid-Cretaceous interval. Whereas flowering plants were of only subordinate importance in Barremian to Aptian terrestrial ecosystems (on average less than 10 %), they experienced a rapid and extensive diversification during the Albian to Cenomanian. By the end of the Cenomanian angiosperms dominated in typical low-latitude floras, accounting for about 70 % of the encountered species (Crane and Lidgard, 1989; Lidgard and Crane, 1988). Fossil floras from the Portuguese Estremadura region play a key role for investigating the late Early Cretaceous angiosperm evolution and diversification. The continental deposits of the Lusitanian Basin have been intensely studied with regard to macrofossil leaf floras (Teixeira, 1948) as well as with regard to the pollen and spores content (Groot and Groot, 1962). In more recent times, several rich and well-preserved mesofossil floras including various in situ pollen have been described in detail by Friis et al. (1997; 1994; 1999; 2000a; 2001) from continental sediments of an inferred Barremian or possibly Aptian age. These floras display a relatively high diversity of in situ pollen, accounting for up to 30 % of the total floral diversity. Triaperturate pollen types represent about ~15 % of the angiosperm pollen diversity. According to these authors, the observed fossil angiosperm reproductive structures Chapter 5 87

(incl. flowers, stamens, anthers, fruits) represent the oldest unequivocal evidence for the occurrence of flowering plants in the fossil record. The proposed Barremian-Aptian age of this diverse angiosperm record contrasts with the previously held view that the major increase in angiosperm diversity occurred during the Albian. Furthermore, the consistent occurrence of triaperturate pollen in the in situ assemblages is in contrast to the absence of this type of pollen in most contemporaneous dispersed palynofloras. However, many of the early angiosperm records lack independent stratigraphic control due to the absence of adequate markers in the fossil-bearing strata. This hampers detailed comparison between dispersed palynofloras and the plant macro- or mesofossil records. Furthermore, a more precise dating of the early angiosperm diversification pattern would allow for a correlation with major climatic or tectonic events during the mid-Cretaceous, which might have had significant influence on the evolution and rapid diversification of the flowering plants (Crane et al., 1995; Lupia et al., 2000). Here, we present independently dated palynological records which document the early angiosperm diversification in Portugal on a previously not attained temporal resolution. Changes in palynofloral composition during the Late Barremian to Early Albian interval are traced throughout two coastal marine successions from the Algarve and Lusitanian Basins. In a first step, the existing stratigraphic model of both successions is revised based on dinoflagellate cyst biostratigraphy. The new results significantly change the stratigraphic assignment of several lithological units. In a second step the palynological content of the two sections is analysed with regard to composition, diversity and relative abundance. Both successions provide well-preserved and diverse angiosperm palynofloras. Our angiosperm pollen records are compared with previously published records from widespread localities including palynofloras from the North American Potomac Group (Brenner, 1963; Doyle and Robbins, 1977) as well as with the in situ pollen assemblages from Portugal (Friis et al. 1997; 1994; 1999; 2000a). Based on the revised stratigraphy and palynological arguments, we provide evidence for a significantly younger, post-Aptian age of the Portuguese mesofossil floras from the northern Lusitanian Basin. Chapter 5 88

2. Studied sections

Two Portuguese localities have been chosen for the present biostratigraphic and palynological study, both representing mixed carbonate-siliciclastic coastal successions and covering Barremian to Albian strata. The first section (Cresmina) is located in the Lusitanian Basin, western Portugal whereas the second section (Luz) is exposed in the Algarve Basin, southern Portugal (Fig. 1).

A 40°30' B 8° R I A Angiosperm mesofossil site E B Palaeozoic Basement I B 25 km Buarcos Coimbratugal

Figueira da Foz Por N C

40°00'

é Fault Zone Famalicão Nazar Nazaré Vale de Agua

39°30'

Peniche

Santa Cruz São Julião Torres Vedras section Catefica 39°00' Ericeira

Lisboa Cascais Lower Tagus Basin Cresmina section

38°30'

-9°45' -9°15'-8°45' -8°15'

C Silves

Lagos Tavira Faro -37°00' Luz section 25 km -9°00' -8°30' -8°00' -7°30'

Fig. 1: (A) Location of the Lusitanian and Algarve Basins in western and southern Portugal. (B) Map of the Estremadura region with the locations of the studied Cresmina and São Julião sections (arrows) and sites of angiosperm mesofossil floras (asterisks). (C) Map of the Algarve region with the location of the Luz section. Chapter 5 89

2.1 Cresmina section

The Cresmina section is well-exposed along the coastal cliffs north of Cabo Raso, about 5 km northeast of the village Cascais. The studied succession spans from the cliffs below the Forte da Cresmina along the beach towards the cliffs of Ponta da Galé (Ramalho et al., 1981; Rey, 1972). Due to unfavourable outcrop conditions at the Forte da Cresmina, the Praia da Lagoa Member has been sampled near São Julião in the Ericeira area, about 25 km north of the Forte da Cresmina site. Here, the Praia da Lagoa Member is exposed along the coastal cliffs below the small village São Julião, ~0.5 km south of the Ribeira do Porto river mouth. Similarly, the Rodízio Formation has been sampled south of Ericeira along the Praia dos Banhos, ~4.0 km north of the Ribeira do Porto river mouth (Rey, 1972). The Cresmina section has been studied in detail by Rey (1972; 1992) from a sedimentological, palaeontologcial and stratigraphical perspective. The section comprises ~200 m of Barremian to Albian sediments and can be separated into five major lithological units (Fig. 2). According to Rey (1972; 1992), the lower part of the section corresponds to the Cresmina Formation, which itself is composed of three individual lithostratigraphic units, including the Cobre, the Ponta Alta and the Praia da Lagoa Member. The upper part of the Cresmina section includes the Rodízio Formation and the lower part of the Galé Formation (Agua Doce Member). The individual lithostratigraphic units are named according to Rey (1972). The Cobre Member is mainly composed of strongly bioturbated and impure limestones, alternating with oyster-rich marls, siltstones and few well-sorted conglomerate layers. These sediments are interpreted to reflect a mixed carbonate-siliciclastic near-shore depositional environment. The overlying Ponta Alta Member consists of massy, thick-bedded, rudist-rich limestones. Besides various rudist taxa, the limestones comprise a diverse macrofauna including stromatoporoids, scleractinian corals and nerinean gastropods, indicating an open- platform depositional setting (Rey, 1979). The top of the Ponta Alta Member is marked by a prominent hardground, which allows precise correlation with sections in the northern part of the Lusitanian Basin (Rey, 1992). This widespread discontinuity is covered by the sediments of the Praia da Lagoa Member, mainly calcareous, orbitolinid-rich marls and fossiliferous sandy limestones. The marine deposits of the Praia da Lagoa Member are overlain disconformly by the coarse-grained siliciclastics and lignite-rich mudstones of the Rodízio Formation. According to Dinis and Trincão (1995), the boundary between the marine deposits of the Praia da Lagoa Member and the continental conglomerates of the Rodízio Formation Chapter 5 90 represents a major unconformity of superregional significance. The coarse-grained siliciclastics evolve gradually into the coastal marine silts, marls and limestones of the lower Agua Doce Member. The upper part of the Agua Doce Member is composed of fossiliferous marly limestones with intercalated rudist-rich horizons, indicating deposition in an inner- to mid-shelf environment.

2.2. Luz section

The Luz section is well exposed along the coastal cliffs southwest of the village Lagos in the western Algarve region. The entire sedimentary succession is slightly tilted towards the east and most of the studied section is accessible along a ~2.5 km long strip between the Praia da Luz (east of the village Luz) and the Praia da Porto de Mós (2 km southwest of Lagos). Only the lowermost part of the section (incl. the Choffatella decipiens Marls and the Palorbitolina Beds) has been sampled along the cliffs at Ponta da Calheta, 0.5 km north of the Praia da Luz (Rocha et al., 1983). Earlier sedimentological and biostratigraphical studies of these deposits have been carried out by Rey and Ramalho (1974), Ramalho and Rey (1981) and Rey (1983; 1986). Following Rey (1983) the ~260 m thick sedimentary succession can be separated into 5 lithostratigraphic units, including the Choffatella decipiens Marls, the Palorbitolina Beds, the Lower and the Upper Luz Marls as well as the Porto de Mós Formation (Fig. 2). The Choffatella decipiens Marls are mainly composed of alternating beds of gypsiferous marls, bioclastic limestones and dolomicrites, which have been deposited in a shallow marine to lagoonal setting. The overlying Palorbitolina Beds are represented by massive, oblique- bedded coastal sandstones, containing abundant nerinean gastropod coquinas. Above a distinct hardground, the Luz Marls consist of a monotonous succession of variegated marls and claystones with few intercalated silt- and limestone beds. The boundary between the Lower and Upper Luz Marls is marked by a distinct interval of thick-bedded fossiliferous limestones with a conglomeratic horizon at the base. The abundant occurrence of charophytes, miliolinid foraminifera and ostracods throughout the Luz Marls indicates deposition in a restricted lagoonal to brackish marsh environment with few open-marine episodes. The Upper Luz Marls are overlain by the Porto de Mós Formation, which is composed of thick-bedded, bioturbated limestones alternating with calcareous marls. Typical sedimentary structures include laminations, bored hardgrounds, desiccation cracks as well as fenestrae, indicating a carbonate-dominated tidal flat depositional environment. Chapter 5 91

Algarve Basin siltstone marl claystone (< 25% CaCO3) sandstone/conglomerate

Meter Stage Formation Lithology limestone first occurrence (FO) 260 last occurrence (FO) Dinopterygium cladoides (consistent occurrence) 240 Lusitanian Basin

220 Member Meter Stage Formation Lithology Porto de Mos Fm. Lower Albian Lower 200 200

180 Dinopterygium cladoides 180 Galé Fm.

160 Hystrichosphaerina schindewolfii 160 Middle Albian

Muderongia staurota Xiphophoridium alatum 140 Chichaouadinium vestitum 140 Subtilisphaera perlucida

Tehamadinium tenuiceras Agua Doce Mb. 120

120 Upper Luz Marls

100 100 Lower Albian Dinopterygium cladoides (consistent occurrence) Pseudoceratium securigerum 80

80 Fm.

Rodisio Ctenidodinium elegantulum Callaiosphaeridium trycherium Heslertonia heslertonensis PdL Upper BedoulianAptian Upper Pseudoceratium securigerum 60 60 Pachytraga paradoxa

Lower Praecaprina varians rudists Bedoulien

Lower Luz Marls Caprina douvillei Ponta Alta Pseudoceratium pelliferum 40 40 Beds Cresmina Fm. 20 Palorbitolina Rhynchodiniopsis aptiana 20 Ctenidodinium elegantulum Cerbia tabulata Cobre Mb. Callaiosphaeridium trycherium Odontochitina operculata

Marls Palaeoperidinium cretaceum Upper Barremian Lower Bedoulian 0 C. decipiens 0

Fig. 2: Simplified lithological log with biostratigraphic events of the Luz section (Algarve Basin) and the Cresmina section (Lusitanian Basin). Age-dignostic rudists in the Ponta Alta Member are displayed in grey. PDL, Praia da Lagoa Member; C. decipiens Marls, Choffatella decipiens Marls. Chapter 5 92

3. Palaeophytogeographic and palaeoclimatic framework

During the mid-Cretaceous, the Algarve and Lusitanian Basins were situated at a palaeolatitude of about 20ºN to 25ºN, forming part of the eastern margin of the evolving North Atlantic (Fig. 3). According to Brenner (1976) and Batten (1984), both Portuguese study sites were part of the southernmost Southern Laurasian floral province, which was restricted to the mid-latitudes of northern hemisphere during Aptian to Albian times. The boundary between the Northern Gondwana province in the south and the Southern Laurasian province in the north is represented by a transitional zone, which incorporates floral elements from both provinces (Batten, 1984; Hochuli, 1981). Palynofloral assemblages from the Southern Laurasian province typically contain abundant bisaccates of Pinacean affinity, conifer pollen such as Classopollis spp. and Araucariacites spp. as well as numerous and varied pteriodophyte spores - especially representatives of the Schizaeaceae and Gleicheniaceae. In contrast, high abundances of various gymnosperm pollen of the Ephedripites, Cycadopites and Araucariacites group indicate a Northern Gondwana affinity. In addition, large numbers of Classopollis spp. as well as the common occurrence of Afropollis spp. characterise this southern floral province whereas pteridophyte spores exhibit generally low diversity and abundance. Bisaccate pollen of Pinacean affinity are virtually absent. According to the palaeoclimatic reconstructions of Chumakov et al. (1995), the Northern Gondwana floral province corresponds to a broad zone of arid to semi-arid conditions (equatorial hot arid belt) during the Aptian interval. This is consistent with the results of Ruffel and Batten (1990) who proposed, based on sedimentological and palynological observations, a Barremian to mid-Aptian phase of aridity for the western European realm. During the Albian, the development of an equatorial humid belt represents a significant change in the palaeoclimatic patterns of low-latitudes (Chumakov et al., 1995). In general, this pattern is supported by palaeobotanical results which indicate an arid to semi-arid climate for the Northern Gondwana province, whereas the Southern Laurasian province was characterised by subtropical to warm-temperate conditions during the Aptian to Albian interval (Brenner, 1976; Chumakov et al., 1995; Vakhrameyev, 1978). Chapter 5 93

-75ºE -60ºE-45ºE -30ºE -15ºE 0ºE

Southern Laurasian province (subtropical to warm-temperate) 5

4 30ºN

1 2 transitional zone 3 15ºN

Northern Gondwana province (arid to semi-arid) 6 8 0ºN

7

-15ºN 9

Fig. 3: Palaeogeographic reconstruction of the North Atlantic and Tethyan realm during the mid- Cretaceous at ~115 Ma (modified after Geomar map generator; www.odsn.de). Asterisks mark locations of early angiosperm palynofloras which are used for comparison. 1, Lusitanian Basin, Portugal (Friis et al. 1999 and this study); 2, Algarve Basin, Portugal (this study); 3, DSDP sites 417 and 418, North Atlantic Basin (Hochuli and Kelts 1980); 4, Potomac Group, United States (Doyle and Robbins, 1977); 5, Wealden Group, England (Hughes et al. 1979); 6, Qattara Depression, Egypt (Ibrahim 1996); 7, Dakhla Oasis, Egypt (Schrank and Mahmoud 2002); 8, northern Negev, Israel (Brenner 1996); 9, Cocobeach system, Gabon (Doyle 1977). Major floral provinces and corresponding climate after Brenner (1976) and Batten (1984).

4. Material and methods

A total of 57 rock samples from the Cresmina section and 61 rock samples from the Luz section were prepared for palynological analysis. Despite the selection of apparently well- suited samples, numerous samples were barren of palynomorphs (27 in the Cresmina section; 27 in the Luz section). Cleaned, crushed and weighed samples (20 to 80 g) were treated with HCl and HF following standard palynological preparation techniques (e.g. Traverse, 1988). The residue was sieved with a 11 µm mesh-sieve and a first set of strew mounts was prepared for kerogen analysis. Following this, a short oxidation with HNO3 was performed on all residues. A second set of strew mounts was prepared for palynological analysis. All Chapter 5 94 productive samples were studied for their palynological content (dinoflagellate cysts, spores and pollen). Special attention was paid to the occurrence of angiosperm pollen. In a second step, a minimum of 200 (average of 240) sporomorphs was determined and counted. Light photomicrographs were taken using an Olympus BX 51 light microscope (LM) equipped with an Olympus DP 12 digital camera. The preservation of the studied palynomorphs is fairly good to excellent. Individual grains exhibit no obvious signs of post-depositional degradation. Thermally unaltered conditions of the OM are indicated by the virtually unchanged colouring of the palynomorphs (TAI < 2) as well as by strong UV fluorescence of the amorphous OM.

5. Dinoflagellate cyst biostratigraphy

In order to establish a refined biostratigraphic framework for the mid-Cretaceous strata of the Lusitanian and Algarve Basins, all productive palynological samples were analysed for the distribution of dinoflagellate cysts. In the Cresmina section a total of 78 different dinoflagellate cyst taxa have been distinguished (Fig. 4), whereas in the Luz section 74 taxa were determined (Fig. 5). No evidence for reworking has been observed. First and last occurrences (FO and LO) of age-diagnostic dinoflagellate cyst taxa are diplayed in Fig. 2. In the present context aiming for an independent age framework for the pollen record, stratigraphic evidence from pollen is not considered in this study. As already reported by Berthou and Leereveld (1990) the observed dinoflagellate cyst assemblages reflect a Boreal rather than Tethyan character. Therefore, comparison and correlation refers mainly to associations from the Boreal realm and corresponding biostratigraphic zonation schemes. To some extent comparison with Tethyan associations has been included. The comprehensive biostratigraphic zonation scheme of Monteil and Foucher (1998) including Boreal and Tethyan dinoflagellate taxa serves as a biostratigraphic baseline. In addition, the zonation schemes of Costa and Davey (1992), Stover et al. (1996) and Leereveld (1995) are applied for comparison and correlation. For regional stratigraphic considerations, the encountered associations are compared with the results of earlier studies of Berthou and Leereveld (1990), Hasenboehler (1981) and Berthou et al. (1980). For the Barremian to Aptian interval, our results are compared with assemblages from N Italy (Torricelli, 2000), SE France (Davey and Verdier, 1974; Masure et al., 1998) and south- western Morocco (Below, 1981). For the nomenclature of the mentioned taxa, we refer to Chapter 5 95

Williams et al. (1998). In the Ponta Alta Member, the occurrences of age diagnostic rudist species corroborate the ages indicated by palynology.

5.1. Cresmina section

Cobre Member The stratigraphic position of this unit is confined by the FOs of the age-diagnostic dinoflagellate cyst taxa Cerbia tabulata, Odontochitina operculata (Pl. VII; 3) and Palaeoperidinium cretaceum in the basal part (at 5.7 m) as well as by the LO of Pseudoceratium pelliferum (Pl. VII; 9) in the uppermost part of the Cobre Member (at 45.0 m). These findings indicate a Late Barremian age for this part of the section. According to Leereveld (1995), the FO of C. tabulata occurs just below the Early to Late Barremian boundary in both, the Boreal and the Tethyan realm. This corresponds to the FO of C. tabulata applied in the biostratigraphic schemes of Monteil and Foucher (1998), Costa and Davey (1992) and Stover et al. (1996). The FO of P. cretaceum at the same level is in agreement with the results of Monteil and Foucher (1998) as well as with Costa and Davey (1992). The FO of O. operculata is known to occur above the Early-Late Barremian boundary (Leereveld, 1995; Torricelli, 2000). The LO of P. pelliferum represents a frequently used stratigraphic event indicating the Barremian-Aptian boundary in the Tethyan realm (Costa and Davey, 1992; Leereveld, 1995; Stover et al., 1996). In the Boreal realm, the LO of P. pelliferum is less consistent and has been reported from latest Barremian (Monteil and Foucher, 1998) as well as from Bedoulian strata (Lister and Batten, 1988; Stover et al., 1996).

Ponta Alta Member The stratigraphic position of the Ponta Alta Member is determined by the occurrence of several age-diagnostic rudist species including Pachytraga paradoxa, Praecaprina varians and Caprina douvillei. These rudist taxa have been reported from sediments of latest Barremian to Bedoulian age (Masse and Chartrousse, 1998; Skelton and Masse, 1998), suggesting a similar age for the Ponta Alta Member. Palynological samples from this interval were barren.

Praia da Lagoa Member The stratigraphic position of the Praia da Lagoa Member is defined by the LOs of Pseudoceratium securigerum (at 63.0 m; Pl. VII; 2), Heslertonia heslertonensis (at 64.9 m; Chapter 5 96

Pl. VII; 1), Callaiosphaeridium trycherium (at 66.2 m) and Ctenidodinium elegantulum (at 68.5 m). This association indicates an age not younger than Early Bedoulian for this part Member. The encountered taxa are considered as important marker species in most biostratigraphic zonation schemes. Masure et al. (1998) used the LO of C. elegantulum together with Rhynchodiniopsis aptiana as a key event in their dinoflagellate cyst zonation for the Bedoulian stratotype, where it defines the top of the securigerum dinoflagellate zone in the lower part of the Late Bedoulian (Deshayesi ammonite zone). Davey and Verdier (1974) and Masure et al. (1998) reported P. securigerum as characteristic species from Bedoulian strata. The LO of H. heslertonensis is applied by Masure et al. (1998) as a marker species for earliest Bedoulian. According to Stover et al. (1996) the LO of C. elegantulum corresponds to the Early Bedoulian. In the Boreal realm, the LOs of C. elegantulum and C. trycherium have been placed into the Bedoulian by Monteil and Foucher (1998). In the zonation schemes of Costa and Davey (1992) and Monteil and Foucher (1998) the LO of H. heslertonensis is placed below the D. deshayesi ammonite zone within the Early Bedoulian.

Rodízio Formation The stratigraphic position of the Rodízio Formation is determined by the FO of Dinopterygium cladoides (at 86.0 m; Pl. VII; 5) occurring just above the second conglomerate horizon. From this horizon onwards, D. cladoides occurs consistently throughout the upper part of the succession. Except for an isolated record, Below (1981) reported the occurrence of D. cladoides (reported as Oodnadattia tuberculata) from Moroccan deposits from the Albian onwards. In addition, D. cladoides has been documented consistently from Cenomanian to Santonian chalks in southern England (Clarke and Verdier, 1967). According to the range chart of Monteil and Foucher (1998) the FO of D. cladoides is placed into the late Early Albian in the Boreal realm. This event clearly indicates an Early Albian or younger age for the upper part of the Rodízio Formation.

Agua Doce Member (Lower Galé Formation) The stratigraphic position of the Agua Doce Member is determined by the LOs of Subtilisphaera perlucida (at 130.6 m; Pl. VII; 4) and Hystrichosphaerina schindewolfii (at 158.8 m; Pl. VII; 7) as well as by the FOs of Chichaouadinium vestitum (at 136.4 m) and Xiphophoridium alatum (at 140.4 m). The observed association indicates an Early to Middle Chapter 5 97

Albian age for the Agua Doce Member. The boundary between the Early and Middle Albian is placed at the FO of X. alatum. According to Leereveld (1995), the LO of H. schindewolfii marks the latest Early Albian in the Boreal realm (D. mammilatum ammonite zone). Other authors place the LO of H. schindewolfii earlier within the Early Albian (Costa and Davey, 1992; Stover et al., 1996). Similarly, the LO of S. perlucida has been considered typical for Early Albian (Costa and Davey, 1992). In the Boreal realm, Monteil and Foucher (1998) report the FO of X. alatum from the late Middle Albian (E. lautus ammonite zone). According to Stover et al. (1996) the FOs of X. alatum and C. vestitum occur as late as earliest Late Albian. " spp. spp. spp. spp. spp. spp. spp. spp.

(A) spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. Meter Stage Sample Member Formation Achomosphaera neptuni Callaiosphaeridium trycherium Cerbia tabulata Cribroperidinium edwardsii Ellipsoidictyum Kiokansium polypes Kleithriasphaeridium readei Kleithriasphaeridium Muderongia staurota Odontochitina operculata Oligosphaeridium asterigerum Palaeoperidinium cretaceum Pseudoceratium securigerum Subtilisphaera perlucida Circulodinium brevispinosum Cribroperidinium Gonyaulacysta cretacea Hystrichodinium furcatum Muderongia Oligosphaeridium Spiniferites Subtilisphaera Trichodinium Callaiosphaeridium Pseudoceratium pelliferum Achomosphaera Batiacasphaera Cribroperidinium orthoceras Exochosphaeridium Hystrichodinium pulchrum Hystrichosphaerina schindewolfii Tehamadinium Kleithriasphaeridium corrugatum Pterodinium Callaiosphaeridium asymmetricum Dingodinium Heslertonia heslertonensis Hystrichosphaerina Microdinium opacum Aptea polymorpha Cometodinium Muderongia pariata Pseudoceratium securigerum "nudum Ctenidodinium elegantulum Geiselodinium Dapsilidinium deflandrei Dinopterygium cladoides Oligosphaeridium pulcherrimum Oligosphaeridium totum Protoellipsodinium corollum Subtilisphaera cheit Dapsilidinium warrenii Exochosphaeridium phragmites Florentina deanei Pinocchiodinium erbae Trichodinium castanea Odontochitina imparilis Odontochitina Cyclonephelium vannophorum Ascodinium Chichaouadinium vestitum Circulodinium Vozzhennikovia Xiphophoridium alatum Cyclonephelium paucispinum Diconodinium Dissiliodinium Polysphaeridium Cleistosphaeridium Tanyosphaeridium variecalamum Chlamydophorella Cepadinium ventriosa Chichaouadinium Microdinium Odontochitina ancala Oligosphaeridium complex Florentinia 200 Kalyptea

L-97 L-91 L-88 180

L-69 + 160 L-66 L-60

Middle Albian L-55 + L-52 + 140 L-48 Galé Fm. L-43 L-40 L-37 120 Agua Doce Mb. L-31 L-25 L-19 L-16 + + + + + + + + L-13 + + + 100 L-5 L-1 Lower Albian K-3.1

80 K-2.1 Fm.

Rodisio E-1s E-6 E-3 + + + + + + + + +

PdL E-1f + + + + + + + + + + + + + + 60 E-1c Lower Bedoulien

Ponta Alta D-62 40

20 Cresmina Fm. D-13 Cobre Mb. D-7 0 Upper Barremian

Fig. 4: Stratigraphical distribution of dinoflagellate cysts in the Cresmina section. Horizontal bars in the sample column represent productive palynological samples, crosses correspond to palynologically barren samples. Chapter 5 98

5.2. Luz section

Choffatella decipiens Marls and Palorbitolina Beds The stratigraphic position of the Choffatella decipiens Marls is defined by the LO of Callaiosphaeridium trycherium (at 10.6 m) as well as by the LOs of Rhynchodiniopsis aptiana and Ctenidodinium elegantulum at (17.5 m). A similar age-diagnostic dinoflagellate cyst assemblage has been observed in the Praia da Lagoa Member (see above), indicating an Early Bedoulian or older age for the Choffatella decipiens Marls. The overlying Palorbitolina Beds contain no appropriate lithologies for palynological analysis.

Luz Marls The biostratigraphic interpretation of the Luz Marls is based on the FO of Tehamadinium tenuiceras (at 129.8 m; Pl. VII; 8) and on the LOs of Pseudoceratium securigerum (at 83.4 m) and Muderongia staurota (at 148.2 m). The FO of T. tenuiceras represents an important event in most biostratigraphic zonations. According to Leereveld (1995), it is slightly diachronous and appears during the Late Bedoulian in the Boreal realm (D. deshayesi ammonite zone). This is consistent with the reported FO of T. tenuiceras (reported as Occisucysta tenuiceras) at the top of the D. deshayesi ammonite zone in the scheme of Lister and Batten (1988). In addition, T. tenuiceras has been documented from the Gargasian of south-western Morocco by Below (1981). Masure et al. (1998) propose a tenuiceras dinoflagellate biozone for SE France, which is defined by the FO of T. tenuiceras marking the Bedoulian-Late Aptian boundary. In the biostratigraphic zonation schemes of Monteil and Foucher (1998) and of Costa and Davey (1992) the LO of M. staurota is placed within the Late Bedoulian at the boundary between the D. deshayesi and T. bowerbanki ammonite zones. Following Masure et al. (1998), we use the FO of T. tenuiceras as a marker for the Bedoulian-Late Aptian boundary. A Bedoulian age for large parts of the Luz Marls is supported by the LOs of M. staurota and P. securigerum.

Porto de Mós Formation The stratigraphic position of the Porto de Mós Formation is determined by the FO of Dinopterygium cladoides (at 183.9 m) and its consistent occurrence in the upper part of the succession (from 240 m onwards), indicating an Early Albian or younger age.

Chapter 5 99 spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. (B) spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. Meter Formation Stage Sample Achomosphaera Batiacasphaera Callaiosphaeridium asymmetricum Callaiosphaeridium Callaiosphaeridium trycherium Cepadinium ventriosa Circulodinium brevispinosum Ctenidodinium elegantulum Hystrichodinium pulchrum Kiokansium polypes Oligosphaeridium complex Pseudoceratium securigerum Rhynchodiniopsis aptiana Subtilisphaera perlucida Subtilisphaera Cerbia tabulata Dingodinium albertii Florentina deanei Hystrichosphaerina schindewolfii Palaeoperidinium cretaceum Spiniferites Trichodinium Oligosphaeridium Protoellipsodinium corollum Cometodinium Cleistosphaeridium Cyclonephelium vannophorum Hystrichosphaeridium arborispinum Pterodinium Subtilisphaera cheit Subtilisphaera pirnaensis Valvaeodinium Canningia Hystrichosphaeridium Tehamadinium Coronifera Dissiliodinium Muderongia pariata Aptea polymorpha Microdinium Circulodinium Apteodinium Chlamydophorella Cribroperidinium Cyclonephelium paucimarginatum Dingodinium Heterosphaeridium Kalyptea Muderongia staurota Odontochitina Oligosphaeridium asterigerum Subtilisphaera senegalensis Tehamadinium tenuiceras Cepadinium Kleithriasphaeridium Odontochitina operculata Oligosphaeridium pulcherrimum Oligosphaeridium totum Ovoidinium Coronifera oceanica Cyclonephelium paucispinum Exiguisphaera Systematophora Ellipsoidictyum Chichaouadinium Exochosphaeridium Geiselodinium Dinopterygium cladoides Odontochitina ancala Trichodinium castanea Muderongia Chytroeisphaeridia Cribroperidinium edwardsii

260 A-201 + + + + + + + + + + + +

A-196 + + + + + + + + + +

A-194 + + + + + + + + + + + + + + A-193 + + + + + + + + + + + 240 A-188 + + + + + + + + + + +

A-179 + + + + + + + + + +

A-176 + + + + + + + + 220

A-172 + + + + + + + + + + + + + + + + + + + + + Lower Albian Lower

Porto de Mos Fm. A-169 + + +

A-162 + + + +

200 A-154 + + + + + + + + + + + + +

A-148 + + + + + + + + + + + +

A-137 + + + + + + + + + + + + + + + + + + + A-134 + + + + + + + + + + + 180

A-125 + + + + A-121 + + + + + + + + + + + + + + + + + +

A-115 + + + + + + + + + + + + + + + 160 A-114 A-112 + + + + + + + + + + + + + + + +

A-110 + +

A-108 + + + + + + + + + + + + + + + + + + + + + + + + A-106 + + +

140 A-101 + + + + + + + + + + + + + + + + + + + + + + +

A-97 A-94 + + + + + + + + + + + + + + + + + + + + + + + + + + + +

120 Upper Luz Marls A-81 A-79 + + + + + + + + + +

100 A-59 + + + + +

A-46 + + + + + + + + + + + + + A-41 + + + + + + + + + + + + + A-37 80 A-33 + + Upper BedoulianAptian Upper 60 Lower Luz Marls

40 Beds Palorbitolina

20 B-13 + + + + + + + + + + + + +

B-8 + + + + + + + + + + + + + + + Marls Lower Bedoulian 0 C. decipiens

Fig. 5: Stratigraphical distribution of dinoflagellate cysts in the Luz section. For explanations see Fig. 4. Chapter 5 100

6. Discussion of the biostratigraphic results

6.1. Cresmina section

The mid-Cretaceous deposits of the Lusitanian Basin have been studied in detail from a bio- and sequence-stratigraphic perspective by Rey (1972; 1992), Rey et al. (1977), Berthou and Schroeder (1979) and Dinis et al. (2002). Palynostratigraphic studies have been carried out by Hasenboehler (1981), Berthou et al. (1980) and Berthou and Leereveld (1990). Our results, presented in this study provide not only a refinement of the biostratigraphic framework (Fig. 6). Distinct changes in the age of individual lithological members will lead to a better understanding of the temporal evolution of the depositional history.

(1) The well-constrained Late Barremian age of the Cobre Member corroborates the biostratigraphic results of Rey (1992), who attributed an latest Barremian to Bedoulian age to this interval. Berthou and Leereveld (1990) interpreted the occurring dinoflagellate cyst assemblage (including Pseudoceratium pelliferum) to represent a Bedoulian age and commented on the lack of evidence for Upper Barremian strata in the Western Portuguese Basin. However, in most recent biostratigraphic zonation schemes, the consistent occurrence of P. pelliferum is regarded as a marker for Barremian or older strata, thus supporting a Late Barremian age for the Cobre Member. (2) Based on the occurrence of Pachytraga paradoxa, Praecaprina varians and Caprina douvillei, the Bedoulian age reported by Rey (1992) for the rudist-bearing limestones of the Ponta Alta Member can be confined to Early Bedoulian. (3) Based on the occurrence of several age-diagnostic dinoflagellate cysts, an Early Bedoulian age for the Praia da Lagoa Member is well-constrained, refining a previously reported Bedoulian to early Late Aptian age (Berthou and Leereveld, 1990; Rey, 1992). (4) Up till now, the stratigraphic position of the Rodízio Formation was loosely defined and an age range between early Late Aptian and Middle Albian has been inferred from palynological evidence (Dinis and Trincão, 1995; Hasenboehler, 1981). Based on the occurrence of Dinopterygium cladoides, an Early Albian or younger age can now be assigned to the top of the Rodízio Formation. The assemblages from the lower part of this formation do not contain age-diagnostic dinoflagellate cysts. However, the spore-pollen assemblages from the uppermost and the lower part of this formation show very similar compositions, indicating an Early Albian age for the entire Rodízio Formation. Consequently , the presence of a major Chapter 5 101 hiatus can be located between the marine marls of the Praia da Lagoa Member and the coarse siliciclastics deposits of the Rodízio Formation. This sedimentary gap encompasses at least the Late Bedoulian and entire Late Aptian and corresponds to the so-called “Lower Aptian unconformity” of Dinis and Trincao (1995). The existence of a major hiatus between the two formations has already been suggested by Berthou and Leereveld (1990). (5) Some stratigraphic discrepancies exist with regard to the Agua Doce Member. Palynological evidence, including the consistent occurrence of Dinopterygium cladoides indicates an Early Albian or younger age for the lower part of this member. Furthermore, Berthou and Leereveld (1990) reported the occurrence of Ovoidinium diversum, O. rhakodes and O. tuberculata from the upper part, which corresponds to a Middle Albian age. These results contradict the orbitolinid-derived biostratigraphy of Rey et al. (1977) and Berthou and Schroeder (1979), who attributed a Late Albian age to the upper part of the Agua Doce Member. Generally, Late Albian dinoflagellate assemblages include several distinct marker species. The absence of these markers in both palynological studies strongly suggests a Middle Albian age.

Stage Rey (1992) Berthou and This study Leereveld (1990)

Upper Albian Agua Doce Agua Doce Mb. Agua Doce Mb.

Middle Albian

Fm.

Rodizio Fm. Rodizio izio izio

Lower Albian od R ? Upper Aptian Praia da Lagoa Mb. Praia da Lagoa Mb. Ponta Alta Mb. ? Ponta Alta Mb. Upper Lower Bedoulian Praia da Lagoa Mb. Cobre Mb.

Aptian Cobre Mb. Lower ? Ponta Alta Mb. Bedoulian Cobre Mb.

Upper Barremian

Fig. 6: Comparison of the different biostratigraphic assignments for the lithostratigraphic units of the Cresmina section (Lusitanian Basin). Grey bars represent stratigraphic rangees of individual units, cross hatch indicates hiatuses, sinuous line indicates major unconformities. Uncertain stratigraphic ranges at are marked with a question mark. Chapter 5 102

6.2. Luz section

In comparison to the Lusitanian Basin, the biostratigraphic assignment of the mid-Cretaceous deposits of the Algarve Basin is less precise. This is mainly a consequence of the proximal position and the resulting restricted conditions of the depositional environment. For this reason, the stratigraphy of the Luz Marls and the Porto de Mós Formations has been mainly based on orbitolinids, calcareous algae and ostracods (Damotte et al., 1988; Ramalho and Rey, 1981; Rey, 1983; Rey, 1986). Additional stratigraphic information is provided by the palynological study of Berthou and Leereveld (1990) and the chemostratigraphic results of Heimhofer et al. (2003). Fig. 7 provides an overview of the refined stratigraphic framework in comparison to earlier studies.

(1) The LOs of several age-diagnostic dinoflagellate cysts within the Choffatella decipiens Marls indicate an Early Bedoulian age. In contrast, a Barremian age has been reported by Rey (1983; 1986) for this unit based on the occurrence of calcareous algae and orbitolinid assemblages. The overlying Palorbitolina beds have been dated as Bedoulian by the same author. (2) The Luz Marls Formation is Late Bedoulian to Late Aptian in age. The FO of Tehamadinium tenuiceras, indicating the Late Bedoulian to Early Aptian boundary corresponds well with the carbon-isotope data (Heimhofer et al., 2003). These results are in general agreement with the Bedoulian to Gargasian biostratigraphic assignment of Rey (1983; 1986) and Damotte et al. (1988) as well as with the undifferentiated Aptian age reported by Berthou and Leereveld (1990). (3) Although the Porto de Mós Formation comprises a relatively diverse dinoflagellate assemblage, only one age-diagnostic marker has been identified. An Early Albian age is based on the common occurrence of Dinopterygium cladoides throughout the Porto de Mós Formation. Negative evidence for Middle Albian is provided by the lack of Middle Albian dinoflagellate cyst markers. The position of the Aptian-Albian transition is marked by a characteristic negative shift in the δ13C record (Herrle et al., 2004). These results are in contrast with earlier age interpretations of the Porto de Mós Formation. Based on microfossil assemblages including calcareous algae, benthic foraminifera and ostracods, Rey (1983; 1986) and Damotte et al. (1988) proposed a Gargasian to Clansayesian age. However, the combined evidence from the dinoflagellate cyst biostratigraphy presented Chapter 5 103 here and the independent carbon-isotope record (Heimhofer et al., 2003) support an Early Albian age.

Stage Rey (1986) Berthou and This study

Leereveld (1990)

Fm.

s o

Middle Albian

orto de M de orto P

Lower Albian Fm. Mos de to Por Luz Malrs Fm. & Porto de Mos Fm. Luz Marls Fm. Luz Malrs Fm.

Upper Aptian Beds Beds Marls Upper Marls

Lower Bedoulian Palorbiolina Palorbiolina Aptian ?? Lower C. depiens ? Bedoulian C. depiens

Upper Barremian

Fig. 7: Comparison of the different biostratigraphic assignments for the lithostratigraphic units within the Luz section (Algarve Basin). For explanations see Fig. 6.

7. Palynological results of the studied sections

54 samples of both successions are analysed quantitatively with regard to the occurring spores and pollen. Gymnosperm pollen and pteridophyte spores were determined on the genera level and several forms were treated as groups (e.g. Classopollis group, bisaccate group). Special attention was paid to the occurring angiosperm pollen assemblages, which were analysed with regard to composition, relative abundance and diversity. The recorded angiosperm taxa and their distinctive morphologic features are listed in Table. 1. The two studied successions comprise a total of 60 different angiosperm pollen types within the Upper Barremian to Middle Albian interval. The most important group (incl. 51 taxa) is represented by monoaperturate grains of probably magnoliid and monocotyledonous affinity. Presumed eudicotyledons are represented by 9 tri- and one stephanocolpate taxa. The angiosperm palynoflora is dominated by columellate-tectate and reticulate-semitectate forms with ornamented or smooth muri. The occurrence of a striate, verrucate or crotonoid pattern is restricted to few taxa. Due to their ambiguous systematic position, pollen of the Afropollis group are not included in the angiosperm assemblage. Chapter 5 104

7.1. Cresmina section

In the sediments of the Cresmina section, 16 different types of gymnosperm pollen, 25 types of spores and a total of 48 angiosperm pollen taxa have been differentiated. Based on quantitative distribution of the major pollen groups four different local pollen zones (LPZ) are distinguished (Fig. 8). The lowermost Upper Baremian to Lower Bedoulian LPZ I (0 m to 68.5 m) is characterised by high abundances of Classopollis spp. (up to 65 %) and bisaccate pollen grains (up to 45 %). The uppermost part this zone (63 m to 68.5 m, Praia da Lagoa Member,) displays a significant increase in Inaperturopollenites spp and Perinopollenites spp. (up to 20 %). Other gymnosperm pollen and trilete spores account for less than 10 %, respectively. Above the major unconformity (MU), LPZ II (78.5 m to 104.5 m) comprises Lower Albian strata. The palynoflora of LPZ II is characterised by an increase in Classopollis spp. (from < 5 % up to ~50 %), high abundance of Inaperturopollenites spp. (~30 %) and a decline in Perinopollenites spp. (from ~25 % to less than 10 %). Other gymnosperm pollen (incl. bisaccate pollen, Araucariacites spp., Exesipollenites spp.) occur in low number whereas trilete spores account for 10 % to 20 %. LPZ III includes the Lower to Middle Albian interval between 104.5 m and 145.5 m. A prominent peak in trilete spores up to ~55 % (incl. Cicatricosisporites spp., Leptolepidites spp., Concavisporites spp. and Echinatisporites spp.) represents the most remarkable feature within this interval. This increase is accompanied by a strong decline in Classopollis spp., Inaperturopollenites spp. and Perinopollenites spp. whereas bisaccate pollen remain essentially stable. In contrast, Araucariacites spp. and Exesipollenites spp. display a slight increase and account for 5 % to 15 %, respectively. The palynoflora of the Middle Albian LPZ IV (145.5 m to 191 m) is characterised by a rapid increase of Exesipollenites spp. up to 45 % and the subsequent decline to less than 10 % towards the top of the succession. This decline is accompanied by a significant increase in Inaperturopollenites spp. as well as by an increase in trilete spores, which account for 10 % to 15 % in this part of the section. Perinopollenites spp. is virtually absent in this zone. The distribution, relative abundance and diversity of angiosperm pollen taxa in the Cresmina section are shown in Fig. 9. The Upper Barremian sediments (3 samples) comprise only two types of angiosperm pollen grains, which are both attributed to the Clavatipollenites group. In the Lower Bedoulian (3 samples) the assemblage is characterised by the appearance of several additional forms of the Retimonocolpites, Asteropollis and Pennipollis groups. Almost all of the taxa recorded in the Barremian to Lower Aptian interval are common throughout the Chapter 5 105 upper part of the section. In the Lower Albian (11 samples), above the major unconformity (MU), the observed angiosperm palynoflora is significantly enriched and additional monoaperturate pollen genera can be distinguished, including Dichastopollenites, Stellatopollis and Racemonocolpites. In addition, various forms of the Retimonocolpites, Clavatipollenites and Asteropollis group are identified. The first tricolpate angiosperm pollen appear in the Lower Albian, including forms of the Tricolpites, Senectotetradites and Striatopollis groups. The Middle Albian (9 samples) interval exhibits further diversification of the angiosperm palynoflora. Numerous FOs of monocolpate pollen are observed within the Retimonocolpites and Dichastopollenites groups whereas the association of tricolpates remains essentially the same as in the Lower Albian.

Gymnosperm pollenAngiosperm Trilete spores pollen spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. Afropollis Meter Sample Lithology Stage Bisaccate Pollen Classopollis Araucariacites Exesipollentes Inaperturopoll. Perinopollenites % % Angiosperms % Trilete spores LPZ 200 Podocarpidites Vitreisporites pallidus Sciadopityspollenites Ephedripites Eucommiidites Callialasporites Cicatricosisporites Deltoidospora Gleicheniidites Leptolepidites Plicatella Rubinella Concavisporites Costatoperforosporites Impardecispora Trilete spores indet. Biretisporites Concavissimisporites Reticulatisporites Foveosporites Cingutriletes Echinatisporis Hamulatisporites Ischyosporites Klukisporites Converrucosisporites Verrucosisporites L-97 L-91 L-88 180

L-69 IV 160 L-66 L-60 Middle Albian L-55 L-52 L-48 140 L-43 Galé Fm. L-40 L-37 III 120 L-31 L-25 L-19 L-16 L-13

100 L-5 L-1 Lower Albian K-3.1 II

80 K-2.1

Rodisio E-1s MU E-6 E-3 E-1f 60 E-1c Lower Bedoulien D-62 40 barren interval I 20 Cresmina Fm. D-13

D-7 Upper Barremian 0 0 100 % Fig. 8: Biostratigraphic interpretation, lithology and quantitative distribution of spores and pollen (Cresmina section). Grey bars mark palynologically barren intervals. Relative abundance of the spores and pollen are expressed in percentages of the total sporomorph assemblage. Local Pollen Zones (LPZ) are marked with dotted lines. MU, major unconformity. Chapter 5 106

The Barremian to Middle Albian deposits show distinct changes in angiosperm pollen relative abundance and diversity. The prominent shift in both parameters corresponds to the major hiatus. The Barremian interval is characterised by the sporadic occurrence of very few angiosperm pollen grains (relative abundance < 2 %). In the Lower Bedoulian angiosperm pollen content is still low (< 2 %) whereas diversity is slightly increased (up to 4 taxa per sample).

Lusitanian monocots and magnoliids eudicots angiosperm relative Basin diversity abundance

exoticus ghazalatensis fredericksburgensis . excelsus hughesii minutus tenellis cf. cf. sp. 2 sp. 1 spp. sp. 4 cf. sp. 5 sp. 6 sp. 5 spp. sp. 6 sp. 4 cf. cf. sp. 11 spp. sp. 3 sp. 8 sp. 5 sp. 2 sp. 10 sp. 12 sp. 11 sp. 1 sp. 13 sp. 16 sp. 7 aff. vermimurus spp cf. cf. cf. sp. A cf. sp. 3 asteroides aff. spp. sp.1 cf. sp. 2 sp. 4 spp. spp. sp. 2 sp. 3 spp. LPZ Stage Meter Formation Lithology Sample Clavatipollenites Clavatipollenites Clavatipollenites Retimonocolpites Asteropollis Pennipollis Retimonocolpites Asteropollis Dichastopollenites Stellatopollis Asteropollis Dichastopollenites Retimonocolpites Retimonocolpites Retimonocolpites Asteropollis asteroides Asteropollis Clavatipollenites Dichastopollenites Racemonocolpites Retimonocolpites Retimonocolpites Stellatopollis Clavatipollenites Dichastopollenites Retimonocolpites Dichastopollenites dunveganensis Stellatopollis barghoornii Pennipollis Retimonocolpites Dichastopollenites Retimonocolpites Retimonocolpites Retimonocolpites Clavatipollenites Dichastopollenites Retimonocolpites Retimonocolpites Retimonocolpites Dichastopollenites Dichastopollenites Retimonocolpites 200 Tricolpites Senectotetradites Striatopollis trochuensis Stephanocolpites Striatopollis Retitricolpites

L-97 + + L-91 + + + L-88 + + + + + 180

L-69 + + + + + IV

160 L-66 + + + + + + + + + L-60 + + + + + + + + + Middle Albian L-55 + + + + + + + L-52 + + + + + + + + + + + + + 140 L-48 + L-43 Galé Fm.

L-40 + + + +

L-37 + + + + + + + + + + + + + III 120 L-31 L-25 L-19 + L-16 + + L-13

100 L-5

L-1 + + + + + + Lower Albian

K-3.1 + + + II

80 K-2.1 + + Rodisio E-1s E-6 + MU E-3 + + E-1f + + 60 E-1c Lower Bedoulien D-62 40 barren interval I Cresmina Fm. 20

D-13

D-7 Upper Barremian 0 0 30 0 25%

Fig. 9: Distribution, diversity and within-palynofloral abundance of angiosperm pollen types plotted against biostratigraphic interpretation and lithology (Cresmina section). Angiosperm diversity represents the number of taxa per sample; relative abundance reflects the percentage of angiosperm pollen within the total palynoflora. Note the abrupt increase in angiosperm pollen diversity and within-pollen abundance above the major unconformity (MU). For explanations see Fig. 8. Chapter 5 107

In the Lower Albian strata above the hiatus, angiosperm pollen represent a common element of the palynoflora and account for 5 % to 8 % of the entire pollen assemblage. At the same time, pollen diversity reaches up to 16 taxa per sample. This increasing trend continues into the Middle Albian part of the section, where peak diversity (up to 18 taxa per sample) and highest relative abundance (up to 12 %) are observed.

7.2. Luz section

20 different types of gymnosperm pollen, 39 different types of spores and a total of 55 different angiosperm pollen taxa have been distinguished in the Luz section (Fig. 10). The Classopollis group accounts for 40 % to 85 % (mean of ~70 %) of the palynoflora throughout the section. Despite some fluctuations, a general increase in the relative abundance of the Classopollis group from ~60 % in the Bedoulian to ~75 % in the Upper Aptian to Lower Albian part can be recognized. This general pattern is interrupted by the virtual absence of Classopollis spp in a single sample at 96.5 m. In contrast, bisaccate pollen (incl. Podocarpidites spp.) display a declining trend from up to ~20 % in the lower part to < 10 % in the upper part of the Luz section. The relative abundance of Araucariacites spp. (up to ~25 %) fluctuates in the opposite direction to the frequency pattern of the Classopollis group. Other gymnosperm pollen (e.g. Exesipollenites spp., Ephedripites spp., Inaperturopollenites spp.) occur rarely and account together for < 5 %. Similarly, pollen of the Afropollis spp. group are sporadically observed. Trilete spores represent a subordinate element of the palynoflora and account for < 10 % on average. Two peaks in relative spore abundance (at 96.5 m and 184 m) reflect increased abundances of Cicatricosisporites spp. and Concavisporites spp., respectively. Relative abundance and diversity of the angiosperm pollen are shown in Fig. 11. In the Lower Bedoulian (2 samples) the assemblage is characterised by several monocolpate types of the Clavatipollenites, Retimonocolpites and Asteropollis groups. In addition, taxa of Pennipollis and Stellatopollis appear in the lowermost samples. These pollen types are relatively common and occur throughout the entire record. The Upper Bedoulian (5 samples) comprises several additional forms of the above mentioned groups. In the Upper Aptian (9 samples), the angiosperm palynoflora shows further diversification which is displayed in the FOs of several forms of the Clavatipollenites, Retimonocolpites, Pennipollis, Stellatopollis and Racemonocolpites groups. Chapter 5 108

Gymnosperm pollenAngiosperm Trilete spores pollen spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. spp. Meter Stage Lithology Sample Bisaccate Pollen Classopollis Araucariacites % Angiosperms % Trilete spores Vitreisporites pallidus Podocarpidites Inaperturopollenites Perinopollenites Exesipollenites Cicatricosisporites Cingutriletes Converrucosisporites Deltoidospora Echinatisporis Foveosporites Plicatella Leptolepidites Trilete spores indet. Costatoperforosporites Klukisporites Retitriletes Gleicheniidites Verrucosisporites Densoisporites Impardecispora Staplinisporites caminus Nodosisporites Neoraistrickia Crybelosporites pannuceus Hamulatisporites Clavatisporites Concavisporites 260 A-201 A-196 A-194 A-193 240 A-188

A-179 A-176 220 A-172 A-169 A-162

200 Albian Lower A-154 A-148

A-137 A-134 180 A-125 A-121

A-115 A-114 160 A-112 A-110 A-108 A-106 140 A-101 A-97 A-94

120 A-81 A-79

100 A-59 A-46 A-41 A-37 80 A-33

60 Upper BedoulianAptian Upper barren interval 40

20 B-13 B-8 Lower Bed. 0 0 100 %

Fig. 10: Biostratigraphic interpretation, lithology and quantitative distribution of spores and pollen (Luz section). For explanations see Fig. 8.

The Lower Albian (12 samples) is characterised by several new forms of the Dichastopollenites group and other monocolpates. The appearance of tricolpate pollen in the Lower Albian strata marks an important event in the composition of the angiosperm palynoflora. Tricolpate forms are predominantly represented by the Tricolpites and the Senectotetradites groups, whereas Rousea spp. and Phimopollenites spp. occur only sporadically. Both, the relative abundance and the diversity of angiosperm pollen display a distinct increase throughout the Luz section. During the Aptian, angiosperm pollen represent a minor constituent of the palynoflora and account for less than 2 % in most Bedoulian samples and Chapter 5 109 for < 4 % in the Upper Aptian samples. Similarly, their diversity is low in the Bedoulian (< 5 taxa in most samples) and increases throughout the Upper Aptian (5 to 10 taxa per sample). From the Upper Aptian to Lower Albian transition onwards, angiosperm pollen represent a consistent and important element (between 5 to 10 %). Their diversity displays a similar trend, reaching 10 to 15 (max. 18) taxa per sample in the Lower Albian part of the succession.

Algarve monocots and magnoliids eudicots angiosperm relative diversity abundance Basin sp. 4

sp. 6 spp. exoticus ghazalatensis

hughesii hughesii tenellis . 3 minutus cf. sp. 1 spp. cf. aff. cf. sp. 1 sp. 2 spp. aff. sp. 3 sp. 9 sp. 10 sp. 6 sp. 3 sp. 7 sp. 1 sp. 15 sp. 16 sp. 13 spp. sp. 8 aff. aff. spp. cf. cf. sp. 1 cf. sp. A cf. sp. 2 sp sp. sp. 1 crisopolensis

asteroides spp. cf. cf. sp. 1 sp. 4 cf. sp. 1 spp. sp. 3 sp. 2 sp. 1 sp. 4 sp. 3 aff. sp. 3 spp. spp. Senectotetradites Sample Formation Lithology Asteropollis Clavatipollenites Clavatipollenites Pennipollis Retimonocolpites Retimonocolpites Asteropollis Clavatipollenites Pennipollis Retimonocolpites sp. 4 Retimonocolpites Stellatopollis Retimonocolpites Retimonocolpites Clavatipollenites Retimonocolpites Tucanopollis Asteropollis asteroides Clavatipollenites Stellatopollis barghoornii Asteropollis Clavatipollenites Pennipollis Pennipollis Clavatipollenites Dichastopollenites Angiosperme inc. sed. 3 Racemonocolpites Stellatopollis Retimonocolpites Asteropollis Retimonocolpites Dichastopollenites Retimonocolpites Stellatopollis Angiosperme inc. sed. 1 Angiosperme inc. sed. 2 Clavatipollenites Dichastopollenites Asteropollis Retimonocolpites Retimonocolpites Clavatipollenites Dichastopollenites Retimonocolpites Retimonocolpites Dichastopollenites Pennipollis Retimonocolpites Tricolpites Tricolpites vulgaris Rousea Senectotetradites Phimopollenites Aff. Stage Meter

260 A-201 + +

A-196 +

A-194 + + ++ + + A-193 +++++ 240 A-188 ++++++ + + +

A-179 +++

A-176 + + ++ + + + 220 A-172 ++ + + +

A-169 + + + Lower Albian Lower Porto de Mos Fm.

A-162 + ++++++

200 A-154 + + + + +

A-148 + + + + + +

A-137 ++ + +++ A-134 + + 180 A-125 A-121 ++ + + + +

A-115 ++ + ++ 160 A-114 A-112 ++ ++ + + A-110

A-108 + +++ A-106 +++++++

140 A-101 + +

A-97 A-94

120 Upper Luz Marls A-81 A-79 +++++++++

100 A-59 A-46 A-41 + A-37 80 A-33 Upper BedoulianAptian Upper 60 barren interval Lower Luz Marls

40

20 B-13

B-8 + + Palorbitolina Beds 0 Lower Bedoulian 030 0 25%

Fig. 11: Distribution, diversity and within-palynofloral abundance of angiosperm pollen types plotted against biostratigraphy and lithology (Luz section). For explanations see Fig. 8 and 9. Chapter 5 110

8. Discussion of the palynological results

8.1. A continuous pollen record of early angiosperm diversification from Portugal

Despite strong differences with respect to depositional setting, tectonic history and overall vegetation patterns, the Cresmina and Luz sections display two closely comparable records of dispersed angiosperm pollen. It is not only the composition of the assemblage, but also the relative abundance and diversity of angiosperm pollen which reflects similar patterns. The occurrence of similar palynofloras at two different localities provides strong evidence, that the observed increase in relative abundance and diversity primarily reflects the incipient dispersion of angiosperm plants in the hinterland of the studied coastal settings. The observed changes in the angiosperm community appear to be less affected by physical environmental factors, including changes in sea level and depositional environment. Combination of the palynological findings from the two studied successions results in a composite record, which covers the Late Barremian to Middle Albian time interval. The here proposed bio- and chemostratigraphic framework allows to trace the successive changes of the angiosperm pollen association through time. Our results are comparable with existing palynological studies of dispersed angiosperm pollen from geographically widespread locations ranging from Barremian to Cenomanian in age (Fig. 3). These sites from palaeolatitudes between ~10°S and ~60°N are mostly situated along the margins of the Tethys Ocean and the evolving North Atlantic and include palynofloras from N-America (Singh, 1971; Singh, 1983; Srivastava, 1977), the North Atlantic basin (Hochuli and Kelts, 1980), northern and western Africa (Doyle et al., 1977; Ibrahim, 1996; Penny, 1986; Schrank and Mahmoud, 2002), Middle East (Brenner, 1996) as well as south and north-western Europe (Friis et al., 1994; Friis et al., 1999; Friis et al., 2000a; Groot and Groot, 1962; Hughes et al., 1979; Hughes and McDougall, 1990; Laing, 1975). Our results from Portugal are compared in detail with published material from the Potomac Group, which represents the oldest exposed Cretaceous unit of the Atlantic coastal plain of the United States. Both palynofloras originate from the southern part of the Southern Laurasian floral province and have been deposited in coastal to continental settings along the margins of the North Atlantic Basin. A first comprehensive description of the Potomac palynology including a pollen-based zonation has been established by Brenner (1963). Subsequent studies of the palynological and macrofossil content of the Potomac Group resulted in additional angiosperm pollen records and further Chapter 5 111 refinement of the pollen-based biostratigraphy (Doyle, 1992; Doyle and Hickey, 1976; Doyle and Robbins, 1977). When possible, the observed pollen grains were assigned to published and described forms. However, many of the presented taxa from the Portuguese successions can not be assigned to any previously reported pollen grains, resulting in a large number of informal species. This seems to be caused by the improved resolution of present-day LM as well as by the fragmentary documentation and partly imprecise description of earlier records. Furthermore, detailed correlation with previously published pollen floras is hampered by the lack of independent stratigraphic control for many records. Due to the absence of adequate marker fossils the age of the pollen-bearing, predominantly continental deposits was based on the occurring palynofloral assemblage (incl. angiosperm pollen) and therefore has to be considered carefully.

- Barremian

The Late Barremian angiosperm pollen assemblage of the Portuguese record is restricted to the occurrence of two monocolpate forms including Clavatipollenites spp. and Clavatipollenites cf. hughesii (Pl. I; 1-2). Small, columellate-tectate pollen grains of the Clavatipollenites group represent a common constituent in early angiosperm assemblages and have been reported by various authors from pre-Aptian sediments around the world (e.g. Doyle et al., 1977; Hughes et al., 1979; Schrank and Mahmoud, 2002). According to Brenner (1996) Clavatipollenites-type pollen occur in sediments from Israel, dated as old as Late Hauterivian. In accordance to our results, Clavatipollenites hughesii and C. cf. hughesii constitute early elements in angiosperm pollen records from the Western North Atlantic (Hochuli and Kelts, 1980) as well as in deposits from the North American Potomac Group (Brenner, 1963; Doyle and Robbins, 1977). Based on comparison with extant forms and in situ palynological records, a strong affinity of the early Clavatipollenites pollen types to pollen of the extant Chloranthaceae family has been inferred (Pedersen et al., 1991).

- Aptian

The Early Bedoulian of the Portuguese successions is characterised by a significant increase in angiosperm pollen diversity, including the FOs of the monocolpate taxa Pennipollis sp. 2 (Pl. IV; 12-13), Clavatipollenites cf. minutus (Pl. I; 3-4) Asteropollis cf. asteroides (Pl. III; Chapter 5 112

13-14), Stellatopollis spp. as well as several forms of the Retimonocolpites group including Retimonocolpites sp. 4 (Pl. II; 3-4), R. aff. sp. 3 and R. sp. 6 (Pl. IV; 1). The genera Pennipollis of Friis Pedersen & Crane (2000a) corresponds to the former Peromonolites (Brenner, 1963) and comprises reticulate-acolumellate pollen with pronounced supratectal sculpture elements. This characteristic pollen type has been frequently reported from Early Aptian palynofloral assemblages (e.g. Brenner, 1996; Doyle and Robbins, 1977; Hochuli and Kelts, 1980; Hughes et al., 1979; Ibrahim, 1996). Doyle (1992) highlights the potential importance of the acolumellate Pennipollis group as a stratigraphic marker for post- Barremian strata. The observed taxa Pennipollis sp. 2 can be compared to Peromonolites reticulatus (Brenner, 1963, Pl. 41, 3-4). Based on in situ studies of Friis et al. (2000a) pollen of the Pennipollis-type display an alismatalean affinity and probably represent early monocots. In accordance to the Portuguese results, small, monocolpate pollen types resembling Clavatipollenites minutus have been reported by Doyle and Robbins (1977) from the Lower to Middle Aptian (lower part of Zone I) of the Potomac Group as well as by Doyle et al. (1977) from deposits from north-western Gabon of probably Aptian age. The columellate-tectate forms Asteropollis asteroides and A. cf. asteroides with a distinct tri- or tetrachotomocolpate aperture have previously been reported from post-Aptian deposits (Doyle and Robbins, 1977; Laing, 1975; Singh, 1983; Srivastava, 1977). In our record, Asteropollis cf. asteroides represents a relatively common form in samples from the Lower Bedoulian to Middle Albian interval. This form shows strong similarities to the microreticulate-tectate pollen form with branched sulcus, described by Doyle and Robbins (1977, Pl. 1, Fig. 24, 25) from the Middle Aptian (upper part of Zone I) of the Potomac Group. Based on analysis of in situ Asteropollis-type pollen and the associated floral organs Friis et al. (1999) mention an affinity to the extant genus Hedyosmum of the Chloranthaceae family. Another typical element of most early angiosperm assemblages are the reticulate-semitectate forms of the Retimonocolpites group. Most of the observed forms show distinct differences to previously published taxa. Besides several long-ranging forms, the Late Bedoulian assemblages comprise several so far not described reticulate-semitectate monocolpate pollen grains including the forms Retimonocolpites sp. 2 (Pl. II; 5-6), R. sp. 8 (Pl. II; 9), R. sp. 9 (Pl. II; 12) and R. sp. 10 (Pl. II; 16-17). The observed Retimonocolpites-type forms are not directly comparable to existing pollen records. Only Retimonocolpites sp. 8 is similar to the form Liliacidites textus of Singh (1971, Pl. 29, Fig. 1-4), which has been reported from Lower Albian sediments of the Peace Chapter 5 113

River area. A form comparable to Tucanopollis crisopolensis (T. aff. crisopolensis, Pl. VI; 3) occurs only in this interval. Reported occurrences of Tucanopollis crisopolensis and T. cf. crisopolensis are restricted to the Northern Gondwana floral province ranging in age from Barremian to Early Aptian (Doyle et al., 1977; Regali, 1989; Schrank and Mahmoud, 2002). A further increase in the diversification of angiosperms characterises the Late Aptian angiosperm assemblage. Besides several additional monocolpate forms of the Asteropollis, Pennipollis and Retimonocolpites groups, the FOs of crotonoid forms including Stellatopollis barghoornii and Stellatopollis sp. 1 (Pl. VI; 12) as well as of Clavatipollis cf. tenellis (Pl. 1; 8-9) and Clavatipollis cf. sp. A sensu Doyle and Robbins (1977) (Pl. I; 5-6) are of particular interest for comparison with published records. Representatives of the Stellatopollis group are part of the earliest angiosperm pollen assemblages and have been documented from pre- Aptian deposits in southern England (Hughes et al., 1979; Hughes and McDougall, 1990) and Egypt (Penny, 1986; Schrank and Mahmoud, 2002). Stellatopollis barghoornii and cf. S. barghoornii cover a relatively long stratigraphic interval from possible Barremian (Doyle et al., 1977) to the Aptian (Ibrahim, 1996) to Early Albian (Doyle and Robbins, 1977). Reported occurrences of Clavatipollenites tenellis and cf. Clavatipollenites tenellis from Upper Aptian to Lower Albian sediments (Subzone II-A) of the North Atlantic (Hochuli and Kelts, 1980) and the Potomac Group (Doyle and Robbins, 1977) correspond well with our findings. Singh (1983) reported the occurrence of C. tenellis from the significantly younger deposits of the Cenomanian Dunvegan Formation of western Canada. The coarsely columellate-tectate Clavatipollenites sp. A sensu Doyle and Robbins (1977) represents one of the earliest pollen types in the basal part of the Potomac Group, which is thought to be Early to Middle Aptian age (Doyle, 1992; Doyle and Robbins, 1977).

- Albian

The Early Albian interval is characterized by further diversification of the monocolpates and the first appearance of tricolpate pollen of presumed eudicotyledonous origin. A remarkable event is the consistent occurrence of relatively large, coarsely reticulate-semitectate pollen assigned to the Dichastopollenites group. Our record includes Dichastopollenites cf. ghazalatensis (Pl. IV; 9), D. dunveganensis (Pl. IV; 8), D. sp. 1 (Pl. I; 1-2), D. sp. 2 (Pl. IV; 4- 5) and D. aff. sp. 4. Except for a single occurrence of Dichastopollenites sp. 1 in the Upper Aptian of the Algarve Basin, pollen of this group are restricted to post-Aptian deposits. In previous studies, pollen of the Dichastopollenites group have been reported from Cenomanian Chapter 5 114 deposits of North Africa (Ibrahim, 1996; Schrank and Mahmoud, 2000) and North America (May, 1975; Singh, 1983). The consistent occurrence of Dichastopollenites cf. ghazalatensis in Albian deposits of both studied sections may reflect the influence of the nearby Northern Gondwana floral province. According to May (1975), Dichastopollenites-type forms resemble operculate pollen of the extant Nymphaeaceae. The discovery of various Dichastopollenites types in the Portuguese Lower Albian deposits results in a significant extension of the stratigraphic range of this group. Further diversification includes the Retimonocolpites, Asteropollis and Clavatipollenites groups. Only few of these forms can be compared to published records. The taxa Retimonocolpites sp. 7 (Pl. II; 10) shows similarities to Retimonocolpites dividuus, which represents a common taxa in Late Aptian to Late Albian assemblages (Brenner, 1963; Doyle and Robbins, 1977; Hochuli and Kelts, 1980; Singh, 1971). In contrast to several published probably Barremian to Aptian assemblages (e.g. Doyle et al., 1977; Doyle and Robbins, 1977; Hughes and McDougall, 1990; Ibrahim, 1996; Penny, 1986), the occurrence of unequivocal tricolpate pollen morphologies is restricted to post-Aptian sediments in the studied Portuguese successions. In our material various tricolpate forms including Tricolpites vulgaris, Senectotetradites spp. (Pl. VI; 10-11) and Striatopollis trochuensis (Pl. VI; 2) appear at or near the base of the Albian. Aff. Stephanocolpites fredericksburgensis (Pl. VI; 9) is the earliest polyaperturate form (stephanocolpate) in the studied succession. In accordance to our results, Doyle and Robbins (1977) reported Stephanocolpites fredericksburgensis from the Lower Albian Zone II B of the Potomac Group. The same species has been observed in deposits as young as Cenomanian by Singh (1983). The occurrence of tetrads of tricolpate pollen grains such as aff. Ajatipollis sp. A is documented from the Early Albian by Doyle and Robbins (1977). This form displays strong similarities with Senectotetradites spp. of the Portuguese records as well as with Senectotetradites amiantopollis described from the Albian by Srivastava (1977). Small, striato-reticulate tricolpates of the Striatopollis group represent another regular constituent of many post-Aptian angiosperm pollen assemblages and have been reported from widespread locations (e.g. Doyle and Robbins, 1977; Groot and Groot, 1962; Hochuli and Kelts, 1980; Laing, 1975; Singh, 1971; Srivastava, 1977). Occurrences of Striatopollis spp. of supposed pre-Albian age are restricted to a few sites located in the Northern Gondwana floral province (Doyle, 1992; Doyle et al., 1977). The species Striatopollis trochuensis has been documented by Ibrahim (1996) from the Cenomanian of Egypt. In accordance to our findings from the Chapter 5 115

Early Albian, Tricolpites vulgaris has been reported e.g. from Middle to Upper Albian deposits of the southern United States (Srivastava, 1977) and western Canada (Singh, 1971). The Middle Albian interval of the Portuguese record is characterised by increasing diversity in the monocolpate Dichastopollenites group, reflected in the FOs of 4 additional taxa (incl. Dichastopollenites sp. 4 (Pl. V; 5-6), D. sp. 5 (Pl. IV; 6-7), D. cf. sp. 5 and D. sp. 6 (Pl. V; 3- 4). These relatively large, coarsely reticulate-semitectate pollen types exhibit no clear similarities to published forms and are reported here as informal species. Further diversification is also observed in the Retimonocolpites group (FOs of Retimonocolpites sp. 11 (Pl. II; 13-14) and R. sp. 12 (Pl. III, 7-8). The only additional tricolpate form is represented by a single grain of aff. Retitricolpites vermimurus. Small tricolpates with a vermiculate reticulum have been originally described as Retitricolpites vermimurus by Brenner (1963) from Aptian to Albian deposits of the Potomac Group. According to Doyle and Robbins (1977) the same formation comprises aff. Retitricolpites vermimurus in the Late Aptian to Early Albian Subzone II A.

In general, the composite angiosperm pollen record from the Portuguese sections corresponds well with the published results from the Potomac Group (Brenner, 1963; Doyle and Robbins, 1977) and the North Atlantic Basin (Hochuli and Kelts, 1980). These palynofloras show strong similarities in the composition of the angiosperm assemblage as well as in the temporal appearance of specific taxa. Differences seem to exist considering the first occurrence of triaperturate pollen types, restricted to post-Aptian deposits in the Portuguese successions and reported form Subzone II B from the Potomac Group (Brenner, 1963; Doyle and Robbins, 1977). Originally dated as Middle Albian by Doyle and Robbins (1977) the age of this Subzone has been considered as Early Albian by Doyle (1992). The similarity of the assemblages, in particular the appearance of tricolpate forms in the Middle Albian of Portugal and in Subzone II B of the Potomac Group suggests a similar age for the two records.

8.2. Palaeoecological and palaeophytogeographic implications

The total palynological assemblage of the Luz section reflects a low diversity of the corresponding flora. The large quantities of Classopollis spp. are produced by xerophytic (drought resistant) and thermophytic Cheirolepidaceae, which are considered to reflect well- drained upland environments (Vakhrameyev, 1982) or mangrove-type, coastal vegetation (Watson, 1988). Other conifers (e.g. Araucariaceae, Pinaceae) as well as different types of Chapter 5 116 ferns are of only subordinate importance. The strong dominance of Cheirolepidaceae pollen in tidally-influenced shallow water deposits (Luz Marls and Porto de Mós Formation) points to a presumable habitate of these plants in the vicinity of the palaeo-shoreline. The high number of Classopollis pollen probably overprinted the vegetation signal from the more distal parts of the catchment area. In the Cresmina section, the observed floral pattern is less stable and exhibits several marked shifts. In general, the vegetation was dominated by various conifer types (incl. Cheirolepidaceae, Araucariaceae, Podocarpaceae, Pinaceae, Taxodiaceae) which occurred in varying abundances. Different types of ferns (e.g. Schizaeaceae, Gleicheniaceae, Dicksoniaceae) were only of subordinate importance in the floral assemblage. A significant increase in fern spores during the Lower Albian (LPZ III) might reflect a shift towards increased humidity in the corresponding hinterland (Herrle et al., 2003; Mohr, 1989). The palynofloral composition of the studied sections clearly supports a palaeophytogeographic position near the southern boundary of the Southern Laurasian floral province (Batten, 1984; Brenner, 1976; Vakhrameyev, 1991). A strong Laurasian influence is reflected in the high abundance of various conifer pollen (incl. Pinacea-derived bisaccates) as well as in the common occurrence and high diversity of fern spores. The proximity of the Northern Gondwana floral province adjacent to the south is documented by typical floral elements such as the rare, but consistent occurrences of the taxa Ephedripites spp. and Afropollis spp. as well as sporadic findings of aff. Tucanopollis crisopolensis. Even though strong differences are observed with respect to the overall palynofloral patterns, the Cresmina and Luz sections display strong similarities in the angiosperm records. In both successions angiosperm pollen represent only a subordinate element of the total palynofloral assemblage. Despite their relatively low overall abundance, the incipient radiation of angiosperms is clearly displayed in the consistently increasing diversity of monoaperturate pollen taxa with time. This trend is paralleled by the rise in relative abundance of these pollen types. Whereas monoaperturates occur only sporadically in Barremian assemblages, they account for up to 12% in the Middle Albian. This distribution pattern indicates that angiosperms successively became an important element of the vegetation at least in low- to mid-latitudes from the Late Aptian onwards. In our successions, presumed eudicots, represented by tricolpate pollen types, appear not before the Early Albian. Compared to the record of monoaperturate forms, their diversity (max. 3 taxa per sample) and relative abundance (less than 7% of the angiosperm pollen count) remain low throughout the Lower to Middle Albian interval. This clearly indicates that eudicot plants formed only a minor Chapter 5 117 constituent of the Portuguese angiosperm flora during the Early to Middle Albian. These findings are in broad agreement with the results of Crane and Lidgard (1989), which indicate that a significant rise in non-magnoliid dicots (eudicots) clearly postdates the Aptian to Albian boundary at palaeolatitudes north of 30°N. Due to major differences in pollen production and dispersal between different plant groups (e.g. insect vs. wind pollination), relative pollen abundances can not be directly translated into vegetation patterns. Detailed examination of early angiosperm reproductive structures and pollen morphologies provide strong evidence for insect pollination and consequently for rather low pollen production (Crane et al., 1995; Friis et al., 1994; Wing and Boucher, 1998). An exception is represented by the Asteropollis-type pollen, which have been interpreted to originate from wind pollinated plants (Friis et al., 1999). Considering insect pollination for the majority of early angiosperms, these plants probably represented a significant part of the late Early Cretaceous vegetation flourishing in the hinterland of the study area.

8.3. Implications for the timing of early angiosperm diversification

8.3.1. Angiosperm mesofossil floras from the Lusitanian Basin

According to our palynological and stratigraphic results, the Portuguese angiosperm palynofloras show a stepwise increase in relative abundance and diversity during the Late Barremian to Middle Albian interval. Assemblages from pre-Albian deposits are characterised by low diversity (max. 11 taxa), low relative abundance (> 4 %) and the lack of tricolpate pollen types. These results are in strong contrast to earlier studies, which suggested a Barremian or possibly Aptian age for highly diverse mesofossil floras from the northern part of the Lusitanian Basin (Friis et al., 1997; Friis et al., 1994; Friis et al., 1999; Friis et al., 2000b; Friis et al., 2001). According to these authors the findings from the Lusitanian Basin comprise the earliest angiosperm reproductive structures and thus, the oldest unequivocal evidence for the occurrence of angiosperms in the fossil record. A high number of well-preserved fossil angiosperm remains such as stamens, flowers, fruits, anthers and seeds have been described. In addition, a variety of in situ pollen from reproductive structures of angiosperms have been documented. The assemblages comprise rich and diverse floras and according to Friis et al. (2001) a conservative estimates accounts for a total of ca. 140 to 150 different angiosperm taxa. These mesofloras are obtained form several localities in the Estremadura region Chapter 5 118 including the Torres Vedras, Catefica, Famalicão, Buarcos and Vale de Agua localities (Fig. 1). A detailed description is provided by Friis et al. (1997; 1999). The fossiliferous deposits are mainly composed of varicoloured clays and silts, intercalated within coarse, cross-bedded sandstones, which reflect deposition in fluvial and/or lacustrine settings. Due to the lack of marine deposits and adequate age-diagnostic fossils, the stratigraphic assignment of these deposits remains problematic. So far, a Barremian or possibly Aptian age has been inferred from the studies of Rey (1972) for the Torres Vedras and the Catefica sites. Due to the strong similarities to the fossil floras from Torres Vedras and Catefica, a similar age has been tentatively assigned to the angiosperm mesofloras from the Famalicão, Buarcos and Vale de Agua sites (Friis et al., 1997; Friis et al., 1999).

8.3.2. Evidence for a post-Aptian stratigraphic position of the mesofossil floras

Several lines of evidence, including palynology, sedimentology and biostratigraphy indicate a post-Aptian age for the angiosperm mesofloras of the Famalicão, Buarcos and Vale de Agua localities.

- Palynological evidence

Friis et al. (1999; 2001) reported a relatively high diversity of up to 30 individual pollen taxa, which have been observed in situ within reproductive organs or adhering to fruiting structures. In the in situ assemblages pollen with a tricolpate aperture configuration account for up to 15 %, the rest consists of monocolpate forms. Considering the suggested Barremian or Aptian age of the mesofossil floras, these results contradict the palynological findings of the dispersed pollen assemblages from the Cresmina and Luz sections. Although not directly comparable, the Barremian to Aptian time interval comprises a significantly lower number of dispersed pollen (2 taxa in the Barremian and up to 11 taxa in the Aptian). In contrast, increased diversity of angiosperm pollen (up to 18 taxa) can be recognized in the post-Aptian part of the successions. Another distinct difference to the in situ results is reflected in the post-Aptian appearance of tricolpate forms and their low relative abundance in the dispersed pollen assemblages (less than 7 %) of the Cresmina and Luz sections. As pointed out by Friis et al. (1999; 2001), standard palynological preparation of the mesofossil-bearing sediments yielded very low pollen diversities at the Famalicão, Buarcos and Vale de Agua sites. Similarly, Pais & Reyre (1981) reported only two angiosperm pollen Chapter 5 119 types (Clavatipollenites cf. hughesii and Asteropollis vulgaris) in dispersed pollen assemblages from the Buarcos location. The observed discrepancy between pollen abundance from dispersed and in situ assemblages has been interpreted as preservational bias or low pollen production and reduced dispersial into the environment of deposition. Even though these processes might result in significant differences between the two types of records, they fail to explain the lack of tricolpates in pre-Albian deposits and their low relative abundance in post-Albian sediments. In addition the palynological results from the Buarcos site contrast to the findings of Groot and Groot (1962) who described various tricolpate and tricolporate pollen types from the same locality in the lower part of the Arenitos de Carrascal unit, clearly indicating a younger, post-Aptian age. A correlation between the in situ pollen assemblage of the Vale de Agua location and the dispersed Albian assemblages is further supported by the occurrence of Dichastopollenites- type pollen. The comparison of SEM micrographs showing Dichastopollenites reticulatus (May, 1975, Pl. 2, Fig. 1-6) with the reticulate-semitectate Pollen Type G of Friis et al. (1999, Fig. 86) displays strong similarities considering size, type and size of muri and luminae as well as the configuration of the colpus. Resemblance between Pollen Type G and LM micrographs of Dichastopollenites cf. ghazalatensis in our material (Pl. IV; 9) is evident. The presence of Dichastopollenites–type pollen grains in the Vale Agua material suggest an age not younger than Albian for this assemblage.

- Sedimentological and biostratigraphic evidence

Additional evidence for an Albian age of the mesofossil-bearing sediments is provided by the refined stratigraphic assignment of the siliciclastic Rodízio Formation and the major unconformity (MU) at its base (Fig. 12). The occurrence of the dinoflagellate marker species D. cladoides in sediments directly above the basal conglomerates indicates an Early Albian or younger age for the Rodízio Formation. The MU represents an important angular discordance in the Lusitanian Basin and corresponds to a break-up unconformity (type 1 sequence boundary), which marks the beginning of oceanic opening of the Atlantic sector adjacent to the Lusitanian Basin (Cunha and Pena dos Reis, 1995; Dinis and Trincão, 1995; Hiscott et al., 1990). The duration of the hiatus between the basal conglomerates above the MU and the underlying strata increases from SSW to NNE. In the Lisbon region (e.g. Cresmina section), the hiatus encompasses Early Bedoulian to Late Aptian, whereas north of Nazaré, mid- Chapter 5 120

Cretaceous siliciclastics rest unconformable on Upper Jurassic to Triassic deposits (Cunha and Pena dos Reis, 1995; Dinis and Trincão, 1995). In the northern part of the Lusitanian Basin, the MU corresponds to the lower limit of the Figueira da Foz Formation (informally termed Belasian Sandstone). The Figueira da Foz Formation comprises an up to 500 m thick continental siliciclastic succession, which covers large areas in the northern Estremadura region. The sedimentary sequence is basically composed of conglomerates, sand- and mudstones, which are arranged in two prominent fining-upward cycles. The basal unit of the lower cycle is represented by the coarse conglomerates of the Calvaria Member. Depositional environments of the Figueira da Foz Formation range from prograding alluvial systems to deltaic and prodelta settings (Dinis et al., 2002). A comparison with earlier studies of Rocha et al.(1981), Teixeira and Zbyszewski (1968) as well as with the recent work of Manuppella et al. (2000) indicates that the mesofossil-bearing, fine-grained deposits of the Famalicão, Buarcos and Vale de Agua sites are intercalated within the siliciclastic sediments of the Figueira da Foz Formation. Towards the south, the basal conglomerates of the Calvaria Member correspond to the coarse siliciclastic deposits of the Rodízio Formation (Cresmina section). The identification of several 2nd-order transgressive-regressive cycles in the Lusitanian Basin allows for an accurate sequence stratigraphic correlation between the lower part of the Rodízio Formation in the south and the Calvaria conglomerates in the north. According to (Dinis et al., 2002) this correlation implies that the base of this lowermost sedimentary cycle (corresponding to the MU) must have approximately the same age throughout the entire Lusitanian Basin or might become slightly younger towards the north. The sedimentological observations clearly indicate that the MU marks the base of the Lower Cretaceous deposits in the northern part of the basin. North of Nazaré there is no evidence for the presence of Cretaceous strata below the MU. Based on the correlation of the basal Figuera da Foz siliciclastics with the conglomerates of the Rodízio Formation, a post-Aptian age is inferred for the onset of sedimentation in the entire Lusitanian Basin. Consequently, the angiosperm mesofossil-bearing deposits of the Famalicão, Buarcos and Vale de Agua sites, which are intercalated within in the Figuera da Foz Formation, are not older than Early Albian in age (Fig. 12). The stratigraphic assignment of the angiosperm mesofloras from the Torres Vedras and Catefica locations remains still problematic. According to Friis et al. (1997; 1999), these mesofloras have been collected from strata ranging from supposedly Valanginian to Lower Chapter 5 121

Barremian. Due to the lack of age-diagnostic fossils, the stratigraphic assignment of the corresponding continental deposits is based on lithostratigraphic correlation with marine strata from the SW part of the basin (Rey, 1972). At the Torres Vedras and Catefica sites, the position of the MU can not be determined clearly and therefore, an unequivocal stratigraphic assignment of the fossil-bearing deposits is not possible on the basis of sedimentological arguments. However, the similarities of the mesofloras from the northern sites (Famalicão, Buarcos and Vale de Agua) with those from further south (Torres Vedras and Catefica) has been taken as evidence for a similar age for all five mesofossil assemblages by Friis et al. (1997; 1999; 2001).

SSW NNE Cresmina Torres Vedras Caranguejeira Buarcos section section section section

Cenomanian

Figueira da Foz Upper Albian Formation

Lower - Middle Albian 3 2 1 Rodizio Formation Calvaria Member MU Lower Lower Upper Upper Barremian? Jurassic Jurassic Bedoulian

Fig. 12: Schematic stratigraphic cross-section throughout the northern part of the Lusitanian Basin from the Cresmina towards the Buarcos study site. The distribution of siliciclastic sediments is marked in grey. Presumed stratigraphic positions of different angiosperm mesofossil sites are marked with an asterisk (1, Buarcos flora; 2, Famalicão flora; 3, Torres Vedras flora). Note the increasing age of the strata below the major unconformity (MU) from SSW towards NNE. Modified after Dinis and Trincão (1995).

The various discrepancies in comparison to our palynological results as well as the refined age of the MU indicate that the Portuguese mesofossil floras are significantly younger than previously suggested. The Early Albian age would not contradict with any of the palaeobotanical findings. In contrast, a revised post-Aptian age for the mesofossil flora clears many discrepancies, which occur in comparison with angiosperm remains (incl. pollen, leaves, wood) from other regions of the world. Chapter 5 122

9. Conclusions

(1) The biostratigraphic study of dinoflagellate cyst associations of mid-Cretaceous deposits from the Lusitanian and Algarve Basins results in significant changes of the existing stratigraphic positions of the individual lithological members. (i) In the Lusitanian Basin (Cresmina section), the Cobre, Ponta Alta and Praia da Lagoa Member are assigned to distinctly older ages than previously suggested. The revised Early Albian age for the major unconformity (MU) and the overlying Rodízio Formation is of significant importance with regard to the sedimentary history and tectonic evolution of the Estremadura region. (ii) In the Luz section (Algarve Basin) a detailed survey of the dinoflagellate cyst assemblages resulted in a shift towards younger ages of almost all lithostratigraphic units. An Early Bedoulian (instead of Late Barremian) age is assigned to the Choffatella decipiens Marls, whereas the Porto de Mos Formation holds an Early Albian instead of a Late Aptian age. The refined stratigraphic framework is consistent with chemostratigraphic results.

(2) The changing pattern in the distribution of the pollen and spores assemblages in the studied sections indicates different vegetation types in the corresponding hinterland. The palynological content of the Luz record is strongly dominated by pollen of the Classopollis group, reflecting probably mangrove-type vegetation adjacent to a tidally-influenced, shallow water depositional setting. In the Cresmina section, the more varied palynological composition is essentially composed of various conifer pollen and fern spores. Several significant shifts in the palynological composition suggest changes in the regional palaeoclimatic conditions. The composition of the palynofloral assemblage is consistent with the previously inferred position of the study sites at the southern rim of the Southern Laurasian floral province.

(3) Both sections provide well-preserved angiosperm pollen assemblages which are studied in detail considering composition, diversity and relative abundance. The occurrence of similar angiosperm pollen patterns at the two different study sites indicates that physical environmental factors are of only subordinate importance for the observed changes in the angiosperm pollen assemblages. Monocolpate pollen with reticulate- and columellate- semitectate sculpture dominate the assemblages, whereas tricolpate pollen types are of only subordinate quantitative importance. Comparison of the Portuguese angiosperm pollen records with previously published results from the North American Potomac Group shows Chapter 5 123 strong similarities in the composition of the assemblage as well as in the temporal appearance of particular pollen types. In addition to well-documented pollen species, the Portuguese sediments comprise a variety of previously unreported Aptian to Albian taxa which have been assigned to represent informal species.

(4) The composite Portuguese angiosperm pollen record displays a clear and continuous increase in relative abundance and diversity which primarily reflects the incipient dispersion of angiosperm plants during the Late Barremian to Middle Albian interval. Based on the refined stratigraphic framework, our results imply that early angiosperm pollen were of only subordinate importance in Late Barremian to Bedoulian palynological assemblages of the western and southern Portuguese Basins. With the first occurrence of tricolpate forms and a variety of additional monocolpate pollen in the Early to Middle Albian, a significant expansion and diversification of the angiosperm flora is observed. This trend is paralleled by an increase in relative abundance which displays the rising importance of angiosperm plant communities in mid-Cretaceous floras. However, presumed eudicotyledons represented by tricolpate pollen types, show relatively low diversity and also low relative abundance throughout the Early to Middle Albian interval. This indicates that plants with eudicot affinity were only a subordinate component of the Portuguese angiosperm flora within this interval.

(5) Our biostratigraphic and palynological results contradict previous stratigraphic assignments of the well-known angiosperm mesofossil floras from the Portuguese Estremadura region. The plant-fossil bearing sediments have been assigned to a Barremian or Aptian age and consequently interpreted to bear the oldest unequivocal remains of angiosperms. However, compared to our palynological results, the occurrence of various tricolpate pollen forms as well as of Dichastopollenites-type pollen within the mesofossil floras indicates an Early Albian or younger age. Stratigraphic evidence for a significantly younger position is provided by the revised Early Albian age for the major unconformity in the Lusitanian Basin. In the northern Estremadura region, this unconformity predates the mesofossil-bearing deposits, clearly indicating a post-Aptian age for the angiosperm plant fossils. Chapter 5 124

Genus Clavatipollenites (COUPER)

Species Author Size and shape Exine Columellae Aperture Plate

C. cf. hughesii Couper (1958) 22 columellate-tectate widely spaced monocolpate Pl. 1; circular-elliptical sexine: 0.5 length: 0.5 well-defined Fig. 1-3 nexine: 1.0 C. cf. minutus Brenner (1963) ~20 columellate-tectate densely spaced not visible Pl. 1; circular-elliptical sexine: 0.6 length: 0.5 Fig. 3-4 nexine: 0.4 C. cf. tenellis Phillips & Felix ~30 perforate-tectate very densely spaced not visible Pl. 1; (1972) irregular spherical sexine: 1.0 length: 1.0 Fig. 8-9 nexine: 1.2 C. sp. 1 informal species ~15 columellate-tectate distinct, widely spaced monocolpate Pl. 1; spherical sexine: 1.0 length: 1.0 well-defined Fig. 7 nexine: 1.0 head Ø: 0.5 C. sp. 2 informal species ~30 columellate-tectate very fine monocolpate Pl. 1; circular-elliptical sexine: 1.0 barely visible elongate Fig. 12-13 nexine: 1.0 C. sp. 3 informal species ~40 columellate-tectate densely spaced monocolpate Pl. 1; circular-elliptical sexine: 1.0 length: 1.0 Fig. 10-11 nexine: 0.5 C. cf. sp. A Doyle & Robbins ~22 microreticulate-tectate densely spaced monocolpate Pl. 1 (1975) spherical sexine: 1.0-1.5 length: 1.5-2.0 Fig. 5-6 nexine: 1.0

Genus Asteropollis (HEDLUND & NORRIS)

Species Author Size and shape Exine Columellae Aperture Plate A. cf. asteroides Hedlund & Norris 20-25 columellate-tectate densely spaced trichotomocolpate Pl. 3; (1968) circular-elliptical sexine: 1.5 length: 0.5 Fig. 12-13 nexine: < 0.5 club shaped A. sp. 1 informal species 25-30 columellate-tectate densely spaced not visible Pl. 3; irregular spherical sexine: 1.0 club shaped Fig. 7-8 nexine: 0.5 length: 1.0 head Ø: ~0.5 A. sp. 2 informal species ~50 microreticulate-tectate very densely spaced not visible Pl. 3; irregular spherical verrucate tectum barely visible Fig. 11

A. sp. 3 informal species ~20 perforate-tectate densely spaced trichotomocolpate Pl. 3; circular-elliptical exine: < 1.0 barely visible Fig. 9-10 A. sp. 4 informal species ~25 microreticulate-tectate densely spaced trichotomocolpate Pl. 4; circular-elliptical sexine: 1.0 club shaped Fig. 2-3 nexine: < 0.3 length: 0.5 - 0.8 head Ø: 0.6

Genus Pennipollis (FRIIS, PEDERSEN & CRANE)

Species Author Size and shape Exine Reticulum Muri width Aperture Plate P. sp. 1 informal ~20 reticulate-semitectate lumina: 1.5-2.5 1.2-1.3 monocolpate Pl. 4; species spherical sexine: 1.5-2.0 transverse ridges Fig. 10-11 nexine: 0.5 P. sp. 2 informal 15 - 20 reticulate-semitectate homobrochate 0.5-0.7 monocolpate Pl. 4; species circular-slightly sexine: 0.75 lumina: 1.3-2.8 verrucate Fig. 12-13 elliptical nexine: 0.75

P. sp. 3 informal ~20 reticulate-semitectate homobrochate 0.5-0.8 monocolpate Pl. 4; species spherical sexine: 0.7 lumina: 1.0-2.5 double-row verrucae Fig. 14-15 nexine: 0.8

P. sp. 4 informal ~15 reticulate-semitectate heterobrochate 0.5 monocolpate Pl. 4; species circular-slightly sexine: 0.5 lumina: 1.5-4.0 fine ornamentation Fig. 16 elliptical nexine: 1.0

Genus Dichastopollenites (MAY)

Species Author Size and Exine Reticulum Muri width Columellae Aperture Plate shape D. cf. ghazalatensis Ibrahim 28–31 reticulate-semitectate polygonal 0.8-1.2 widely spaced zono- Pl. 4; (1996) lumina width: 1.5-4.0 Ø: 1.1-1.3 aperturate Fig. 9 D. dunveganensis Singh ~45 reticulate-semitectate heterobrochate; polygonal 1.0-1.3 very widely zono- Pl. 4; (1983) lumina width: 3.5-6.5 spaced aperturate Fig. 8 Ø: 1.2-1.6

D. sp. 1 informal 26–35 reticulate-semitectate heterobrochate; polygonal 0.6-0.8 very widely zono- Pl. 5; species circular- sexine: ~1.5 µm lumina width: 2.0-6.0 spaced aperturate Fig. 1-2 elliptical nexine: 0.5 – 0.7 µm Ø: 0.8-0.9

D. sp. 2 informal 23–27 reticulate-semitectate heterobrochate; polygonal 0.6-0.8 widely spaced zono- Pl. 4; species circular- sexine: 1.5 – 2.0 µm lumina width: 1.0-3.0 Ø: 0.7-0.8 aperturate Fig. 4-5 elliptical nexine: 0.5 µm club shaped

D. sp. 4 informal 45–48 reticulate-semitectate irregular-heterobrochate 0.8-1.0 widely spaced zono- Pl. 5; species circular- sexine: 2.5 – 3.0 µm incomplete meshes Ø: 1.1-1.3 aperturate Fig. 5-6 elliptical nexine: 0.5 µm club shaped

D. sp. 5 informal ~31 reticulate-semitectate irregular-heterobrochate 0.5-0.7 widely spaced zono- Pl. 4; species elliptical sexine: 1.3 – 1.5 µm incomplete meshes triangular Ø: 0.8-1.0 aperturate Fig. 6-7 nexine: 0.5 µm profile spindle shaped

Chapter 5 125

Genus Retimonocolpites (PIERCE)

Species Author Size Exine Reticulum Muri width Columellae Aperture Plate R. cf. Ward ~35 reticulate-semitectate coarse reticulate 0.6-0.8 widely spaced monocolpate Pl. 2; excelsus (1986) nexine: 0.6-0.7 irregular, loosely attached length: ~1.0 Fig. 11

R. sp. 1 informal ~15 reticulate-semitectate heterobrochate 0.5 length: ~1.0 monocolpate Pl. 2; species sexine: 0.6 loosely attached Fig. 1-2 nexine: 0.7 lumina width: 1.0-2.5

R. sp. 2 informal ~25 reticulate-semitectate homobrochate; polygonal 0.8 widely spaced monocolpate Pl. 2; species sexine: 2.0 lumina width: 2.0-5.0 dispersed club shaped Fig. 5-6 nexine: < 0.4 verrucae length: ~2.0 head Ø: 1.0 R. sp. 3 informal ~25 reticulate-semitectate heterobrochate; irregular 0.5 widely spaced monocolpate Pl. 2; species sexine: 2.0 lumina width: 2.0-3.5 few length: ~2.0 Fig. 7-8 nexine: 0.5 verrucae head Ø: ~0.8

R. sp. 4 informal ~22 reticulate-semitectate heterobrochate; polygonal 0.2-0.3 widely spaced monocolpate Pl. 3; species sexine: 1.3-1.5 lumina width: 1.2-3.0 dispersed length: 1.2-1.5 Fig. 3-4 nexine: 0.5 verrucae head Ø: ~0.5

R. sp. 5 informal ~30 microreticulate-tectate irregular microreticulate very dense densely spaced monocolpate Pl. 2; species sexine: 1.0 lumina width: < 0.5 length: 0.5-1.0 Fig. 15 nexine: 0.5 head Ø: ~0.5

R. sp. 6 informal 30-36 reticulate-semitectate heterobrochate; polygonal 0.2-0.3 densely spaced monocolpate Pl. 4; species sexine: 0.5 microreticulate barley visible elongate colpus Fig. 1 nexine: 0.5 lumina width: < 1.0

R. sp. 7 informal 30-33 reticulate-semitectate homobrochate 0.5-1.2 densely spaced monocolpate Pl. 2; species sexine: 1.0-1.5 smaller towards colpus irregular; club shaped long colpus Fig. 10 nexine: 1.0 lumina width: 1.0-3.0 beaded length: 1.0 head Ø: 1.0-1.2 R. sp. 8 informal ~15 reticulate-semitectate heterobrochate < 0.3 widely spaced monocolpate Pl. 2; species very thin sexine loosely attached regular length: 0.3 Fig. 9 nexine: 0.5 lumina width: < 1.0

R. sp. 9 informal ~24 reticulate-semitectate homobrochate 0.5 widely spaced monocolpate Pl. 2; species sexine: 1.5 lumina width: 2.0-3.0 regular club shaped Fig. 12 nexine: 0.8 length: 1.5-2.0 head Ø: ~0.5 R. sp. 10 informal ~16 reticulate-semitectate extremely heterobrochate 0.2-0.3 thin monocolpate Pl. 2; species sexine: 1.0 irregular length: < 0.5 Fig. 16-17 nexine: 0.5 lumina width: < 1.5

R. sp. 11 informal ~21 reticulate-semitectate heterobrochate 0.5-0.8 widely spaced monocolpate Pl. 2; species sexine: 2.0 irregular triangular spindle-shaped, Fig. 13-14 nexine: 0.5 lumina width: 2.0-5.5 profile length: < 2.0

R. sp. 12 informal ~22 reticulate-semitectate homobrochate 0.7 widely spaced monocolpate species sexine: 1.8 smaller towards colpus club-shaped elongate colpus nexine: 0.7 lumina width: 1.0-2.0 length: 0.7-1.0

R. sp.13 informal ~37 reticulate-semitectate homobrochate 0.3 densely spaced monocolpate Pl. 3; species sexine: 1.7 lumina width: < 2.0 club shaped elongate colpus Fig. 3-4 nexine: 0.3 length: 1.7 head Ø: 0.5-0.7 R. sp. 15 informal ~28 reticulate-semitectate heterobrochate 0.3 densely spaced, monocolpate Pl. 3; species sexine: 1.3 loosely attached club shaped Fig. 5-6 nexine: 1.2 lumina width: 0.5-1.7 length: 1.2

R. sp. 16 informal 25-30 reticulate-semitectate homobrochate < 0.3 densely spaced monocolpate Pl. 3; species sexine: 1.2 lumina width: < 1.3 dispersed club shaped elongate colpus Fig. 1-2 nexine: 0.6 verrucae length: 1.2

Table 1: Descriptive data for the observed pollen mentioned in the text. Due to the lack of documentation of comparable forms in earlier studies, most pollen are reported as informal species. All morpholocial specifications are given in µm.

Acknowledgements

We thank R. Gonzales from Algarve University and P. Skelton from the Open University, Milton Keynes for field assistance and determination of rudist bivalves. Financial support from ETH-Project TH-34./99-4 is greatfully acknowledged. Chapter 5 126

References

Batten, D.J., 1984. Palynology, climate and the development of Late Cretaceous floral provinces in the Northern Hemisphere: A review. In: P. Brenchley (Editor), Fossils and climate. John Wiley and Sons, New York, pp. 127-164. Below, R., 1981. Dinoflagellaten-Zysten aus dem Oberen Hauterive bis Unteren Cenoman Süd-West- Marokkos. Palaeontographica, 176, 1-145. Berthou, P.Y., Foucher, J.C., Lecocq, B. and Moron, J.M., 1980. Aperçu sur les kystes de Dinoflagellés de l'Albien et du Cénomanien du Bassin Occidental Portugais. Cretaceous Research, 2, 125-141. Berthou, P.Y. and Leereveld, H., 1990. Stratigraphic implications of palynological studies on Berriasian to Albian deposits from western and southern Portugal. Review of Palaeobotany and Palynology, 66, 313-344. Berthou, P.Y. and Schroeder, R., 1979. Découverte d'un niveau à Simplorbitolina CIRY et RAT dans l'Albien de Guincho (région de Lisbonne, Portugal). Comptes Rendus Hebdomadaires des Seances de l'Academie des Sciences, Serie D: Sciences Naturelles, 288, 591-594. Brenner, G., 1976. Middle Cretaceous floral provinces and early migration of Angiosperms. In: C.B. Beck (Editor), Origin and early evolution of angiosperms. Columbia University Press, New York, pp. 23-44. Brenner, G.J., 1963. The spores and pollen of the Potomac Group of Maryland. Department of Geology, Mines and Water Resources Bulletin, 27. Department of Geology, Mines and Water Resources, Baltimore, Maryland, 215 pp. Brenner, G.J., 1996. Evidence for the earliest stage of angiosperm pollen evolution: a paleoequatorial section from Israel. In: D.W. Taylor and L.J. Hickey (Editors), Flowering plant origin, evolution and phylogeny. Chapman & Hall, pp. 91-115. Chumakov, N.M. et al., 1995. Climatic belts of the mid-Cretaceous time. Stratigraphy and Geological Correlation, 3, 241-260. Clarke, R.F.A. and Verdier, J.-P., 1967. An investigation of microplankton assemblages from the chalk of the Isle of Wight, England. Verhandelingen der Koninklijke Nederlandse Akademie van Wetenschappen, Afdeeling Natuurkunde, 24, 1-96. Costa, L.I. and Davey, J., 1992. Dinoflagellate cysts of the Cretaceous system. In: A.J. Powell (Editor), A stratigraphic index of dinoflagellate cysts. Chapmann & Hall, London, pp. 99-153. Crane, P.R., Friis, E.M. and Pedersen, K.R., 1995. The origin and early diversification of angiosperms. Nature, 374, 27-33. Crane, P.R. and Lidgard, S., 1989. Angiosperm diversification and paleolatiudinal gradients in Cretaceous floristic diversity. Science, 246, 675-678. Cunha, P.P. and Pena dos Reis, R., 1995. Cretaceous sedimentary and tectonic evolution of the northern sector of the Lusitanian Basin (Portugal). Cretaceous Research, 16, 155-170. Damotte, R., Ramalho, M. and Rey, J., 1988. Les ostracodes de l'Algarve occidental (Portugal). Revue de Micropaleontologie, 31, 38-48. Davey, R.J. and Verdier, J.-P., 1974. Dinoflagellate cysts from the Aptian type sections at Gargas and La Bédoule, France. Palaeontology, 17, 623-653. Dinis, J., Rey, J. and de Graciansky, P.C., 2002. The Lusitanian Basin (Portugal) during the late Aptian-Albian: Sequential arrangement, proposal of correlations, evolution. Comptes Rendus Geoscience, 334, 757-764. Dinis, J.L. and Trincão, P., 1995. Recognition and stratigraphical significance of the Aptian unconformity in the Lusitanian Basin, Portugal. Cretaceous Research, 16, 171-186. Doyle, J.A., 1992. Revised palynological correlations of the lower Potomac Group (USA) and the Cocobeach sequence of Gabon (Barremian-Aptian). Cretaceous Research, 13, 337-349. Doyle, J.A., Biens, P., Doerenkamp, A. and Jardiné, S., 1977. Angiosperm pollen from the pre-Albian Lower Cretaceous of equatorial Africa. Bull. Centr. Rech. Explor.-Prod. Elf-Aquitaine, 1, 451-473. Chapter 5 127

Doyle, J.A. and Hickey, L.J., 1976. Pollen and leaves from the mid-Cretaceous Potomac Group and their bearing on early angiosperm evolution. In: C.B. Beck (Editor), Origin and early evolution of angiosperms. Columbia University Press, New York, pp. 139-206. Doyle, J.A. and Robbins, E.I., 1977. Angiosperm pollen zonation of the continental Cretaceous of the Atlantic coastal plain and its application to deep wells in the Salisbury Embayment. Palynology, 1, 43-78. Friis, E.M., Crane, P.R. and Pedersen, K.R., 1997. Anacostia, a new basal angiosperm from the Early Cretaceous of North America and Portugal with trichotomocolpate/monocolpate pollen. Grana, 36, 225-244. Friis, E.M., Pedersen, K.R. and Crane, P.R., 1994. Angiosperm floral structures from the Early Cretaceous of Portugal. Plant Systematics and Evolution, 8 (supplement), 31-49. Friis, E.M., Pedersen, K.R. and Crane, P.R., 1999. Early angiosperm diversification: the diversity of pollen associated with angiosperm reproductive structures in Early Cretaceous floras from Portugal. Annals of the Missouri Botanical Garden, 86, 259-296. Friis, E.M., Pedersen, K.R. and Crane, P.R., 2000a. Fossil floral structures of a basal angiosperm with monocolpate, reticulate-acolumellate pollen from the Early Cretaceous of Portugal. Grana, 39, 226-239. Friis, E.M., Pedersen, K.R. and Crane, P.R., 2000b. Reproductive structure and organization of basal angiosperms from the Early Cretaceous (Barremian or Aptian) of Western Portugal. International Journal of Plant Sciences, 161, 169-182. Friis, E.M., Pedersen, K.R. and Crane, P.R., 2001. Fossil evidence of water lilies (Nymphaeales) in the Early Cretaceous. Nature, 410, 357-360. Groot, J.J. and Groot, C.R., 1962. Plant microfossils from Aptian, Albian and Cenomanian deposits of Portugal. Comunicações dos Serviços Geológicos de Portugal, 44, 133-176. Gübeli, A.A., Hochuli, P.A. and Wildi, W., 1984. Lower Cretaceous turbiditic sediments from the Rif chain (northern Morocco): palynology, stratigraphy and palaeogeographic setting. Geologische Rundschau, 73, 1081-1114. Hasenboehler, 1981. Étude paléobotanique et palynologique de l'Albien et du Cénomanien du bassin occidental portugais au sud de l'accident de Nazaré (province d'Estremadura, Portugal), Paris. Heimhofer, U., Hochuli, P.A., Burla, S., Andersen, N. and Weissert, H., 2003. Terrestrial carbon- isotope records from coastal deposits (Algarve, Portugal): a tool for chemostratigraphic correlation on an intrabasinal and global scale. Terra Nova, 15, 8-13. Herrle, J.O., Kößler, P., Friedrich, O., Erlenkeuser, H. and Hemleben, C., 2004. High-resolution carbon isotope records of the Aptian to Lower Albian from SE France and the Mazagan Plateau (DSPD site 545): a stratigraphic tool for paleocanographic and paleobiologic reconstruction. Earth and Planetary Science Letters, 218, 149-161. Herrle, J.O., Pross, J., Friedrich, O., Kössler, P. and Hemleben, C., 2003. Forcing mechanisms for Mid-Cretaceous black shale formation; evidence from the upper Aptian and lower Albian of the Vocontian Basin (SE France). Palaeogeography, Palaeoclimatology, Palaeoecology, 190, 399-426. Hiscott, R.N. et al., 1990. Comparative stratigraphy and subsidence history of Mesozoic rift basins of North Atlantic. American Association of Petroleum Geologists Bulletin, 74, 60-76. Hochuli, P. and Kelts, K., 1980. Palynology of Middle Cretaceous black clay facies from Deep Sea Drilling Project sites 417 and 418 of the western North Atlantic. In: T. Donnelly et al. (Editors), Initial reports of the Deep Sea Drilling Projects. Texas A & M University Ocean Drilling Program, College Station, TX, United States, pp. 897-935. Hochuli, P.A., 1981. North Gondwanan floral elements in Lower to Middle Cretaceous sediments of the Southern Alps (Southern Switzerland, Northern Italy). Review of Palaeobotany and Palynology, 35, 337-358. Hughes, N.F., Drewry, G.E. and Laing, J.F., 1979. Barremian earliest aniosperm pollen. Palaeontology, 22, 513-535. Hughes, N.F. and McDougall, A.B., 1990. Barremian-Aptian angiospermid pollen records from southern England. Review of Palaeobotany and Palynology, 65, 145-151. Chapter 5 128

Ibrahim, M.I.A., 1996. Aptian-Turonian palynology of the Ghazalat-1 well (GTX-1), Qattara Depression, Egypt. Review of Palaeobotany and Palynology, 94, 137-168. Laing, J.F., 1975. Mid-Cretaceous angiosperm pollen from southern England and northern France. Palaeontology, 18, 775-808. Leereveld, H., 1995. Dinoflagellate cysts from the Lower Cretaceous Rio Argos succession (SE Spain), Utrecht University, Utrecht, 175 pp. Lidgard, S. and Crane, P.R., 1988. Quantitative analyses of the early angiosperm radiation. Nature, 331, 344-346. Lister, J.K. and Batten, D.J., 1988. Stratigraphic and palaeoenvironmental distribution of early cretaceous dinoflagellate cysts in the Hurlands farm borehole, west Sussex, England. Palaeontographica, 210, 9-89. Lupia, R., Crane, P.R. and Lidgard, S., 2000. Angiosperm diversification and mid-Cretaceous environmental change. In: S.J. Culver and P.F. Rawson (Editors), Biotic response to global change: the last 245 million years. Cambridge University Press, Cambridge, United Kingdom, pp. 207-222. Manuppella, G. et al., 2000. Carta geológica de Portugal na escala de 1/50 000. Notícia explicativa da folha 27-A Vila Nova de Ourém. Instituto Geológico e Mineiro, Lisboa, pp. 156. Masse, J.-P. and Chartrousse, A., 1998. Les Caprina (rudistes) de l'Aptian inférieur d'Europe occidentale: systématique, biostratigraphie et paléobiogéographie. Geobios, 30, 797-809. Masure, E., Raynaud, J.F., Pons, D. and de Reneville, P., 1998. Palynologie du stratotype historique de l'Aptien inférieur dans la région de Cassis-La Bédoule (SE France). Géologie Méditerranéenne, 25, 263-287. May, F., 1975. Dichastopollenites reticulatus, gen. et sp. nov.- potential Cenomanian guide fossil from southern Utah and northeastern Arizona. Journal of Paleontology, 49, 528-533. Mohr, B.A.R., 1989. New palynological information on the age and environment of Late Jurassic and Early Cretaceous vertebrate localities of the Iberian Peninsula (eastern Spain and Portugal). Berliner Geowissenschaftliche Abhandlungen (A), 106, 291-301. Monteil, E. and Foucher, J.C., 1998. Cretaceous Biochronostratigraphy. In: P.-C. De Graciansky, J. Hardenbol, T. Jacquin and P. Vail (Editors), Mesozoic and Cenozoic Sequence Chronostratigraphic Framework of European Basins. SEPM Special Publication. Pais, J. and Reyre, Y., 1981. Problèmes posés par la population sporo-pollinique d'un niveau à plantes de la série de Buarcos (Portugal). Boletim da Sociedade Geologica de Portugal, 22, 35-40. Pedersen, K.R., Crane, P.R., Drinnan, A.N. and Friis, E.M., 1991. Fruits from the mid-Cretaceous of North America with pollen grains of the Clavatipollenites type. Grana, 30. Penny, J.H.J., 1986. An Early Cretaceous angiosperm pollen assemblage from Egypt. Special Papers in Palaeontology, 35, 121-134. Ramalho, M.M. and Rey, J., 1981. Reflexions sur la Formation cretacée de Porto de Mos (Algarve, Portugal). Comunicações dos Serviços Geológicos de Portugal, 67, 35-39. Ramalho, M.M. et al., 1981. Carta geológica de Portugal na escala de 1/50 000. Notícia explicativa da folha 34-C Cascais. Serviços Geológicos de Portugal, Lisbon. Regali, M.S.P., 1989. Tucanopollis, a new genus of early angiosperms. Boletim de Geosciências da PETROBRÁS, 3, 395-402. Rey, J., 1972. Recherches géologiques sur le crétacé inférieur de l'Estremadura (Portugal). Memoria (N.S.), 21. Serviços Geológicos de Portugal, Lisbon, 1-477 pp. Rey, J., 1979. Les formations bioconstruites du Crétacé inférieur d'Estramadura (Portugal), L'Urgonien des pays mediterraneens. Géobios, Mémoire spécial, Lyon, pp. 89-99. Rey, J., 1983. Le Crétacé de l'Algarve: Essai de Synthèse. Comunicações dos Serviços Geológicos de Portugal, 69, 87-101. Rey, J., 1986. Micropaleontological assemblages, paleoenvironments and sedimentary evolution of Cretaceous deposits in the Algarve (southern Portugal). Palaeogeography, Palaeoclimatology, Palaeoecology, 55, 233-246. Rey, J., 1992. Les unités lithostratigraphiques du Crétacé inférieur de la région de Lisbonne. Comunicações dos Serviços Geológicos de Portugal, 78, 103-124. Chapter 5 129

Rey, J., Bilotte, M. and Peybernes, B., 1977. Analyse biostratigraphique et paléontologique de l'Albien marin d'Estremadura (Portugal). Geobios, 10, 369-393. Rey, J. and Ramalho, M.M., 1974. Le Crétacé inférieur de l'Algarve occidental (Portugal). Comunicações dos Serviços Geológicos de Portugal, 5. Rocha, R.B., Manuppella, G., Mouterde, R., Ruget, C. and Zbyszewski, G., 1981. Carta geológica de Portugal na escala de 1/50 000. Notícia explicativa da folha 19-C Figueria da Foz. Serviços Geológicos de Portugal, Lisboa. Rocha, R.B., Ramalho, M.M., Antunes, M.T. and Coelho, A.V.P., 1983. Carta geológica de Portugal na escala de 1/50 000. Notícia explicativa da folha 52-A, Portimão. Serviços Geológicos de Portugal, Lisbon. Ruffell, A.H. and Batten, D.J., 1990. The Barremian-Aptian arid phase in western Europe. Palaeogeography, Palaeoclimatology, Palaeoecology, 80, 197-212. Schrank, E. and Mahmoud, M.S., 2000. New taxa of angiosperm pollen, miospores and associated palynomorphs from the early Late Cretaceous of Egypt (Maghrabi Formation, Kharga Oasis). Review of Palaeobotany and Palynology, 112, 167-188. Schrank, E. and Mahmoud, M.S., 2002. Barremian angiosperm pollen and associated palynomorphs from the Dakhla Oasis area, Egypt. Palaeontology, 45, 33-56. Singh, C., 1971. Lower Cretaceous microfloras of the Peace River Area, Northwestern Alberta. Research Council of Alberta Bulletin, 28. Research Council of Alberta, Edmonton, 299 pp. Singh, C., 1983. Cenomanian microfloras of the Peace River area, northwestern Alberta. Alberta Research Council Bulletin, 44. Alberta Research Council, 322 pp. Skelton, P.W. and Masse, J.-P., 1998. Revision of the lower Cretaceous rudist genera pachytraga PAQUIER and retha COX (bivalvia: hippuritacea), and the origins of the caprinidae. Geobios, 22, 331-370. Srivastava, S.K., 1977. Microspores from the Fredericksburg Group (Albian of the southern United States, Université des Sciences et Techniques Montpellier, Montpellier, 119 pp. Stover, L.E. et al., 1996. Mesozoic-Tertiary Dinoflagellates, Acritarchs and Prasinophytes. In: J. Jansonius and D.C. McGregor (Editors), Palynology: Principles and Applications, Volume II. American Association of Stratigraphic Palynologists Foundation, Salt Lake City, Utah, pp. 641-750. Teixeira, C., 1948. Flora mesozóica portuguesa, Part I. Memórias dos Serviços Geológicos de Portugal, Lisboa, 119 pp. Teixeira, C. and Zbyszewski, G., 1968. Carta geológica de Portugal na escala de 1/50 000. Notícia explicativa da folha 23-C Leiria. Serviços Geológicos de Portugal, Lisboa, pp. 99. Torricelli, S., 2000. Lower Cretaceous dinoflagellate cyst and acritarch stratigraphy of the Cismon APTICORE (southern Alps, Italy). Review of Palaeobotany and Palynology, 108, 213-266. Traverse, A., 1988. Paleopalynology. Unwin Hyman, Boston, 600 pp. Vakhrameyev, V.A., 1978. The climates of the Northern Hemisphere in the Cretaceous in the light of paleobotanical data. Paleontological Journal, 12, 143-154. Vakhrameyev, V.A., 1982. Classopollis pollen as an indicator of Jurassic and Cretaceous climate. International Geology Review, 24, 1190-1196. Vakhrameyev, V.A., 1991. Jurassic and Cretaceous floras and climates of the Earth. Cambridge University Press, Cambridge, 318 pp. Watson, J., 1988. The Cheirolepidiaceae. In: C.B. Beck (Editor), Origin and evolution of the Gymnosperms. Columbia University Press, New York, pp. 382-447. Williams, G.L., Lentin, J.K. and Fensome, R.A., 1998. The Lentin and Williams Index of fossil dinoflagellates 1998 edition. American Association of Stratigraphic Palynologists, Contributions Series, 34, 817 pp. Willis, K.J. and McElwain, J.C., 2002. The evolution of plants. Oxford University Press, Oxford, New York, 378 pp. Wing, S.L. and Boucher, L.D., 1998. Ecological aspects of the Cretaceous flowering plant radiation. Annual Review of Earth and Planetary Sciences, 26, 379-421.

Chapter 5 130

Scale bar is 10 µm in all photomicrographs

Plate I

1-2 Clavatipollenites cf. hughesii (Couper 1958), L-13, 106.9 m, (late Barremian-middle Albian) 3-4 Clavatipollenites cf. minutus (Brenner 1963), L-60, 154.6 m, (early Aptian-middle Albian) 5-6 Clavatipollenites cf. sp. A (Doyle and Robbins 1977), L-52, 145.5 m, (early to middle Albian) 7-8 Clavatipollenites sp. 1, A-106, 146.6 m, (late Aptian) 9-10 Clavatipollenites cf. tenellis (Phillips and Felix 1971), A-112, 157.8 m, (late Aptian-early Albian) 11-12 Clavatipollenites sp. 3, L-66, 158.8 m, (early to middle Albian) 13-14 Clavatipollenites sp. 2, A-106, 146.6 m, (early Albian)

Plate II

1-2 Retimonocolpites sp. 1, A-176, 225.8 m, (early to middle Albian) 3-4 Retimonocolpites sp. 4, L-52, 145.5 m, (early Aptian-middle Albian) 5-6 Retimonocolpites sp. 2, L-52, 145.5 m, (early Aptian-middle Albian) 7-8 Retimonocolpites sp. 3, L-52, 145.5 m, (late Aptian-middle Albian) 9 Retimonocolpites sp. 8, A-108, 148.3 m, (early Aptian-middle Albian) 10 Retimonocolpites sp. 7, A-154, 200.7 m, (early to middle Albian) 11 Retimonocolpites cf. excelsus (Ward 1986), L-1, 92.6 m, (early Albian) 12 Retimonocolpites sp. 9, A-106, 146.6 m, (early Aptian-late Aptian) 13-14 Retimonocolpites sp. 11, L-60, 154.6 m, (middle Albian) 15 Retimonocolpites sp. 5, L-31, 121.1 m, (early Albian) 16-17 Retimonocolpites sp. 10, A-79, 116.3 m, (early Aptian-middle Albian)

Plate III

1-2 Retimonocolpites sp. 16, L-66, 158.8 m, (early to middle Albian) 3-4 Retimonocolpites sp. 13, L-66, 158.8 m, (early to middle Albian) 5-6 Retimonocolpites sp. 15, A-193, 245.7 m, (late Aptian-early Albian) 7-8 Retimonocolpites sp. 12, L-55, 149.7 m, (middle Albian) 9 Asteropollis sp. 1, A-193, 245.7 m, (late Aptian-early Albian) 10-11 Asteropollis sp. 3, A-176, 225.8 m, (early Albian) 12 Asteropollis sp. 2, L-40, 130.6 m, (early Albian) 13-14 Asteropollis cf. asteroides (Hedlund and Norris 1968), L-16, 108.9 m, (early Aptian-middle Albian) Chapter 5 131

Scale bar is 10 µm in all photomicrographs

Plate IV

1 Retimonocolpites sp. 6, L-1, 92.6 m, (early Aptian-middle Albian) 2-3 Asteropollis sp. 4, L-37, 125.0 m, (late Aptian-middle Albian) 4-5 Dichastopollenites sp. 2, L-66, 158.8 m, (early Albian-middle Albian) 6-7 Dichastopollenites sp. 5, L-66, 158.8 m, (middle Albian) 8 Dichastopollenites dunveganensis (Singh 1983), L-37, 125.0 m, (early to middle Albian) 9 Dichastopollenites cf. ghazalatensis (Ibrahim 1996), L-1, 92.6 m, (early to middle Albian) 10-11 Pennipollis sp. 1, A-106, 146.6 m, (late Aptian) 12-13 Pennipollis sp. 2, A-79, 116.3 m, (early Aptian-early Albian) 14-15 Pennipollis sp. 3, L-52, 145.5 m, (early to middle Albian) 16 Pennipollis sp. 4, A-108, 148.3 m, (late Aptian)

Plate V

1-2 Dichastopollenites sp. 1, L-16, 108.9 m, (late Aptian-middle Albian) 3-4 Dichastopollenites sp. 6, L-66, 158.8 m, (middle Albian) 5-6 Dichastopollenites sp. 4, L-88, 184.1 m, (middle Albian)

Plate VI

1 Racemonocolpites cf. exoticus, A-188, 240.1 m, (late Aptian-middle Albian) 2 Striatopollis trochuensis (Ward 1986), L-66, 158.8 m, (early to middle Albian) 3 Tucanopollis cf. crisopolensis (Regali 1989), A-79, 116.3 m, (early to late Aptian) 4 Angiosperme inc. sed. 3, A-179, 231.4 m, (late Aptian-early Albian) 5-6 Angiosperme inc. sed. 1, A-162, 206.5 m, (early Albian) 7-8 Angiosperme inc. sed. 2, A-162, 206.5 m, (early Albian) 9 Aff. Stephanocolpites fredericksburgensis (Hedlund and Norris 1968), L-31, 121.1 m, (early Albian) 10-11 Senectotetradites spp., A-188, 240.1 m, (early to middle Albian) 12 Stellatopollis sp. 1, A-188, 240.1 m, (early to middle Albian) Chapter 5 132

Scale bar is 10 µm in all photomicrographs

Plate VII

1 Heslertonia heslertonensis, E-3, 64.9 m, (LO: Early Bedoulian) 2 Pseudoceratium securigerum, D-13, 12.3 m, (LO: Late Bedoulian) 3 Odontochitina operculata, L-19, 110.7 m, (FO: Late Barremian) 4 Subtilisphaera perlucida, E-1, 63.4 m, (LO: Early Albian) 5 Dinopterygium cladoides, K-3, 86.0 m, (FO: Early Albian) 6 Hystrichosphaeridium arborispinum, A-112, 152.9 m, (LO: Late Aptian) 7 Hystrichosphaerina schindewolfii, E-1, 63.4 m, (LO: Middle Albian) 8 Tehamadinium tenuiceras, A-94, 129.9 m, (FO: Late Aptian) 9 Pseudoceratium pelliferum, D-62, 45.0 m, (LO: Late Barremian)

133

Plate I

7

1 3 5

8

2 4 6

9 11 13

10 12 14

135

Plate II

1

3 5 6

2

4 7 8

9

12

11 10

16

13 14 15 17

137

Plate III

1

3 4

2 7 8

9

10 11

5 13

6 12 14

139

Plate IV

2 3

1

10 11

4 5

12 13

6 7

14 15

8 9 16

141

Plate V

1 2

3 4

5 6

143

Plate VI

2

4 3

1

10 7

5

11

8

6 9 12

145

Plate VII

2

1 3

5 6

4

789

Chapter 6 147

Chapter 6

Conclusions

In order to understand the causes and consequences of past environmental change during the mid-Cretaceous, this thesis addressed the role of terrestrial palaeoenvironments during times of major perturbations. Accurate stratigraphy is crucial for the proposed study and therefore, much effort has been put on the establishment of a detailed time framework for the chosen sedimentary archives. The combined approach of palynology, carbon isotope studies and organic geochemistry is demonstrated to be a successful strategy to investigate the response of continental vegetation patterns to major changes of the ocean-atmosphere system. The most important findings of this study include the following conclusions:

• Prominent shifts in the Early Aptian δ13C record can be reproduced in marine carbonates, individual organic compounds of probable marine origin as well as in land plant-derived organic matter allowing for chemostratigraphic correlation on a high resolution. Furthermore, this indicates that fluctuations in the Aptian δ13C record were not restricted to the marine realm but affected the entire ocean-atmosphere system.

• Palynological and geochemical studies of the late Early Aptian OAE 1a black shales in the Vocontian Basin provide no evidence for prominent climate cooling, accompanied

by a significant drop in palaeoatmospheric pCO2 as previously proposed.

• The micropalaeontological results from the Vocontian Basin question the commonly held few of a high-productivity scenario for the formation of OAE 1a black shales. Instead, sea-level fluctuations, probably associated with decreased runoff are suggested to have played a key role for the deposition of black shale.

• Based on chemo- and biostratigraphical results, a revised, more accurate stratigraphic framework for the Upper Barremian to Middle Albian deposits of the Portuguese Algarve and Lusitanian Basins is presented.

Chapter 6 148

• A continuous and well-dated angiosperm pollen record from the Portuguese basins displays the incipient diversification of flowering plants during the Late Barremian to Middle Albian interval on a so far not obtained resolution. Due to the general lack of well-dated angiosperm records, these results are of significant importance for the temporal calibration of existing angiosperm pollen and macrofossil data.

• Based on palynological, bio- and sequence-stratigraphic arguments, a previously supposed Barremian to Aptian age for several well-known angiosperm floras from the Estremadura region is shown to be Early Albian or younger in age. The revised age of the Estremadura angiosperm floras directly influences the current view of the early angiosperm evolution.

These findings contribute significantly to the understanding of past environmental and floral change during the mid-Cretaceous. The presented study emphasizes the value and importance of continent-derived data for a better understanding of the Mesozoic climate and carbon-cycle perturbations and their possible link to major floral changes. Appendix 149

A1: Total organic carbon (TOC), total inorganic carbon (TIC) and carbon isotope results, Luz section

13 13 13 13 13 Sample Hight TOC TIC δ Cbulk OM δ Ccharcoal δ Clignite δ Cleaves δ Ccuticle (m) (dry wt %) (dry wt %) (‰ VPDB) (‰ VPDB) (‰ VPDB) (‰ VPDB) (‰ VPDB)

A-1 46.3 0.1 86.0 -22.9 A-3 46.8 0.0 34.3 -22.6 A-5 49.7 0.1 59.4 -22.7 A-8 52.1 0.1 64.7 -25.1 A-10 52.6 0.1 30.6 -24.7 A-12 54.7 0.1 22.5 -21.1 A-13 54.7 0.3 25.8 -22.1 A-15 55.5 0.2 13.2 -23.3 A-18 56.5 0.0 31.5 -22.0 A-20 56.9 0.1 67.7 -22.6 A-22 58.9 0.2 28.8 -21.2 A-23 62.7 -24.3 A-25 64.9 0.1 22.6 -23.8 A-27 69.4 0.1 35.7 -24.9 A-30 74.5 0.1 0.0 -27.8 A-31 77.1 0.0 1.3 -26.0 A-32 81.0 0.1 0.0 -25.8 A-33 81.5 0.0 0.0 -21.6 -19.4 -20.9 A-35 83.1 0.2 27.3 -25.2 A-36 83.2 0.2 19.5 -23.7 A-37 83.5 0.1 3.9 -21.7 -19.2 A-38 85.6 0.0 6.5 -22.5 A-39 86.9 0.1 7.2 -21.7 A-40 88.2 0.2 56.3 -19.4 A-41 89.1 0.3 43.8 -21.1 -19.6 -20.0 A-43 89.4 0.1 90.9 -24.6 -22.7 A-44 89.5 0.1 46.9 -21.3 A-46 90.2 0.3 8.1 -21.6 A-47 90.5 0.4 84.0 -21.4 A-48 91.8 0.1 92.3 -19.6 A-49 91.9 0.1 51.1 -19.9 -19.2 A-50 92.0 0.1 83.4 -20.9 A-51 92.2 0.0 86.9 -23.1 A-53 93.0 0.1 74.0 -22.1 A-55 94.1 0.0 46.3 -21.4 A-57 95.0 0.1 75.7 -22.1 A-59a 96.7 -20.1 A-62 101.4 0.1 76.9 -22.1 A-63 103.2 0.1 47.1 -21.6 A-65 104.4 0.1 1.1 -22.9 A-66 104.8 0.1 84.4 -23.8 A-68 105.8 0.0 37.3 -23.4 A-70 107.7 0.1 37.2 -22.8 A-71 110.4 -22.4 A-72 112.8 0.1 44.4 -23.5 A-74 114.5 0.1 31.5 -23.4 A-75 115.8 0.3 83.7 -22.3 A-76 115.9 0.1 58.8 -25.3 -23.2 A-77 116.1 0.1 70.1 -25.4 A-78 116.2 0.1 78.1 -24.8 A-79 116.3 0.5 56.7 -25.3 -24.0 -24.1 A-80 116.5 0.1 81.2 -24.3 A-81 116.7 0.3 58.8 -24.6 -23.8 -26.5 A-82 117.3 0.1 72.8 -25.1 A-84 118.8 -21.6 A-85 119.9 0.1 39.5 -23.8 A-87 122.8 0.1 35.5 -20.5 A-88 123.6 0.0 28.2 -24.0 A-90 125.3 0.1 56.4 -22.2 A-91 127.4 0.1 35.0 -21.9 A-93 128.2 0.1 67.7 -22.9 A-94 129.9 0.3 15.5 -22.5 -21.7 -21.8 -19.1 A-95 130.2 0.0 76.7 -22.2 A-97 132.7 0.1 28.6 -22.2 A-98 135.4 0.2 55.8 -22.3 A-99 137.2 0.1 44.8 -21.4 A-100 139.9 0.0 70.1 -21.2 A-101 140.7 0.3 71.5 -22.8 -21.9

Appendix 150

A-102 141.2 0.1 82.8 -23.6 A-105a 141.4 -23.8 -21.3 A-104 142.1 0.4 71.9 -23.1 A-105 145.8 0.1 50.7 -21.7 A-107 147.5 1.0 0.0 -23.1 -20.0 -22.3 -21.7 A-108 148.3 0.3 32.7 -23.4 -22.8 A-109 150.1 0.0 78.7 -23.8 A-110 152.9 0.1 87.5 -22.5 A-112 157.8 0.6 35.4 -23.5 -21.3 -23.0 -22.1 A-112a 158.4 -23.2 -19.8 A-113 158.7 0.4 84.1 -23.1 A-114 159.9 0.2 30.6 -22.4 A-115 162.9 0.3 73.7 -22.7 A-116 166.8 0.1 75.1 -23.3 A-117 167.8 0.4 62.0 -20.8 A-119 170.3 0.1 77.8 -21.6 A-120 170.9 0.5 54.6 -20.8 -19.8 A-121 171.8 0.1 77.6 -22.2 -22.1 A-122 172.6 0.1 87.2 -24.5 A-123 173.1 0.5 89.6 -24.2 A-125 174.1 0.2 90.2 -19.4 A-126 175.3 -21.3 A-128 176.1 0.1 57.9 -19.3 A-129 178.5 0.2 70.4 -21.9 A-130 181.4 0.1 58.9 -22.7 -20.8 A-131b 181.9 -21.0 -21.5 A-131a 182.4 0.1 79.8 -23.2 A-132 183.1 0.1 51.1 -21.7 A-134 184.0 0.3 88.4 -25.8 A-135 185.2 0.2 20.9 -21.5 -24.2 A-137 186.7 0.1 82.7 -24.7 A-138 187.2 0.3 83.4 -23.7 A-140 188.9 0.1 63.8 -22.4 A-141 190.5 0.2 88.2 -21.1 A-142 191.3 0.1 64.4 -21.3 -22.3 A-143 192.7 0.0 88.6 -22.6 A-145 193.8 0.4 91.6 -22.9 A-146 194.3 0.3 63.9 -21.0 A-148 195.3 0.3 22.6 -23.1 -21.6 -23.1 -25.6 A-149 196.7 0.1 91.5 -25.0 A-151 197.2 0.3 80.5 -25.6 -23.0 A-152 198.5 0.1 93.3 -23.9 A-153 200.1 0.3 48.0 -22.2 -20.9 -21.6 A-155 201.4 0.3 91.9 -22.7 A-156 201.9 0.0 93.4 -23.9 A-158 203.2 0.5 73.4 -23.4 -22.5 A-159a 204.5 -18.3 -21.8 A-160 205.7 0.0 94.9 -24.9 A-160a 206.1 -24.3 A-162 206.5 0.8 75.9 -24.2 A-165 208.9 0.1 92.4 -23.0 A-166 210.4 0.3 38.2 -22.8 A-168 212.3 0.1 94.0 -25.1 A-169 213.1 1.2 64.1 -24.9 -26.9 -25.9 A-170 214.4 0.4 89.2 -22.7 A-171 215.3 0.8 71.5 -26.0 -24.2 -26.3 A-172 217.4 0.2 82.5 -20.7 A-173 219.3 0.3 93.0 -21.9 A-174a 221.9 -23.1 A-175 222.9 0.1 93.4 -24.2 A-176 225.9 0.3 48.4 -23.1 -24.8 A-178 229.9 0.2 89.7 -24.0 A-179 231.4 0.1 91.0 -22.7 A-180 232.6 0.1 90.9 -23.0 -20.3 A-182 233.7 0.2 93.3 -23.3 A-183 234.4 0.1 82.6 -22.9 A-185 235.1 0.2 88.7 -23.7 A-186 235.4 0.1 51.7 -22.2 A-187 237.2 0.1 92.2 -24.3 A-188 240.1 0.2 89.4 -24.3 -20.1 A-188a 240.4 -23.5 A-190 242.1 0.2 79.2 -24.0 A-191 242.9 0.2 67.6 -23.2 A-192 244.1 0.2 91.6 -24.4

Appendix 151

A-193 245.7 0.6 66.7 -24.2 -19.6 -21.3 -23.3 -23.5 A-194a 248.2 -19.8 A-195 250.4 0.1 84.5 -23.8 A-196 254.9 0.7 51.0 -22.5 -22.1 -21.6 -21.9 A-198 256.5 0.1 88.4 -24.4 A-199 257.4 0.1 59.1 -24.3 -20.1 A-200 258.9 0.0 85.9 -24.2 -21.1 A-202 264.1 0.0 86.7 -22.1 A-203 265.6 0.3 88.7 -21.5 A-205 268.5 0.0 41.7 -24.4 A-207 271.5 0.1 72.1 -22.2 A-208 273.4 0.3 85.6 -21.8 A-210 275.7 0.0 68.3 -23.7 A-211 277.0 0.1 93.4 -24.8 A-212 280.2 0.1 87.1 -23.3 A-213 282.9 0.2 94.4 -22.0 A-214 284.7 0.1 78.0 -22.1 A-216 285.4 0.3 67.0 -25.6 A-218 289.4 0.0 90.6 -23.5 A-219 291.9 0.4 80.2 -23.2 A-220 294.4 0.3 88.7 -25.4 A-222 296.9 0.2 97.1 -23.0 A-223 298.0 0.0 47.4 -23.0

A2: Total organic carbon (TOC), total inorganic carbon (TIC) and carbon isotope results, Burgau section

13 Sample Hight TOC TIC δ Cbulk OM (m) (dry wt %) (dry wt %) (‰ VPDB)

G-41 0.2 0.1 11.5 -22.3 G-42 2.2 0.1 31.5 -23.1 G-43 4.4 0.1 20.9 -21.5 G-44 5.4 0.2 71.5 -22.6 G-45 7.0 0.0 0.8 -24.5 G-46 9.0 0.0 0.0 -24.4 G-47 12.6 0.1 0.0 -25.1 G-48 18.7 0.0 1.4 -25.0 G-49 20.0 0.1 2.6 -25.9 G-50 21.0 0.0 2.5 -25.1 G-4 21.4 0.0 25.7 -24.9 G-5 22.2 0.1 34.9 -22.7 G-6 23.6 0.1 55.9 -21.3 G-9 24.4 0.3 38.3 -21.4 G-10 24.8 0.1 54.4 -22.6 G-11 25.8 0.2 72.1 -21.3 G-12 26.7 0.1 65.8 -23.1 G-13 27.7 0.2 79.1 -21.6 G-14 28.4 0.1 71.0 -23.4 G-15 31.6 0.1 76.4 -23.9 G-16 38.8 0.1 21.7 -21.3 G-17 40.8 0.1 65.6 -23.1 G-18 41.2 0.3 79.7 -23.5 G-19 41.8 0.1 72.8 -22.5 G-20 51.5 0.6 80.4 -25.0 G-21 52.0 0.0 93.8 -23.4 G-22 57.0 0.1 50.7 -23.0 G-24 61.3 0.1 73.4 -23.1 G-26 67.3 0.0 10.0 -22.7 G-27 71.9 0.0 6.4 -22.9 G-30 77.0 0.1 25.0 -21.1 G-32 81.0 0.2 79.7 -23.3 G-33 85.2 0.1 31.8 -22.7 G-34 87.5 -25.3 G-35 91.0 0.1 30.1 -20.9 G-36 93.3 -24.6 G-37 94.6 0.3 44.3 -20.4 G-39 96.2 0.0 93.1 -26.8 G-40 97.8 -23.2

Appendix 152

A3: Total organic carbon (TOC), total inorganic carbon (TIC) and carbon isotope results, Cresmina section

13 Sample Hight TOC TIC δ Cbulk OM (m) (dry wt %) (dry wt %) (‰ VPDB)

D-1 1.0 0.0 0.0 -23.8 D-3 2.4 0.0 0.0 -23.1 D-4 3.1 0.4 0.0 -22.9 D-5 3.9 0.3 81.9 -22.9 D-7 5.7 0.7 80.7 -24.1 D-8 7.7 0.0 78.8 -22.8 D-9 8.7 0.2 46.5 -23.8 D-12 11.8 0.0 92.3 -24.0 D-16 14.3 0.0 92.7 -23.3 D-19 17.2 0.0 94.2 -24.2 D-26 20.7 0.1 0.0 -24.6 D-27 21.1 0.1 91.6 -24.4 D-33 22.9 0.0 94.3 -21.9 D-34 23.2 0.0 3.4 -22.4 D-35 24.3 0.1 80.2 -23.2 D-38 25.4 0.1 93.9 -21.9 D-39 25.9 0.0 32.2 -22.9 D-42 29.2 0.1 0.8 -21.7 D-44 30.7 0.1 93.2 -23.3 D-45 31.8 0.1 93.0 -24.2 D-47 34.7 0.3 92.6 -23.4 D-48 36.1 0.1 66.1 -23.1 D-49 37.4 0.1 7.5 -23.0 D-52 38.9 0.0 95.6 -21.6 D-57 42.6 0.0 64.7 -23.1 D-60 44.1 0.0 77.3 -22.5 D-62 45.0 0.1 79.2 -24.3 D-65 47.5 0.1 37.6 -23.2 D-68 51.3 0.0 96.7 -23.8 D-72 54.6 0.0 95.4 -23.0 D-74 57.4 0.0 97.3 -21.9 D-77 61.1 0.1 97.5 -21.9 D-79 63.7 0.1 59.0 -22.7 D-83 69.0 0.0 98.2 -21.6 D-87 72.0 0.4 0.0 -21.7

Appendix 153

A4: Total organic carbon (TOC), total inorganic carbon (TIC) and carbon isotope results, Serre Chaitieu section

13 13 Sample Hight TOC TIC δ Ccarb δ Cbulk OM (m) (dry wt %) (dry wt %) (‰ VPDB) (‰ VPDB)

NS-36 11.75 0.8 27.3 -24.1 NS-35 11.35 0.6 25.7 3.9 -23.7 NS-34 10.95 0.7 30.7 3.9 -23.6 NS-33 10.55 0.6 35.7 3.5 -24.1 NS-32 10.15 0.6 22.0 4.1 -23.8 NS-31 9.75 0.8 20.1 4.3 -24.0 NS-30 9.35 1.2 15.5 4.5 -24.2 NS-29 9.15 1.8 12.6 4.4 -24.2 NS-28 9.05 0.9 18.8 4.1 -24.2 NS-27 8.60 1.1 16.7 4.0 -23.8 NS-26 8.20 0.5 21.1 3.8 -23.8 NS-25 7.80 0.4 26.4 3.5 -23.9 NS-24 7.40 0.6 21.0 3.6 -23.7 NS-23 7.00 0.9 23.1 3.6 -24.5 NS-22 6.55 0.9 15.4 3.8 -24.7 NS-21 6.45 0.5 16.8 3.5 -24.7 NS-20 6.25 0.9 21.5 4.0 -24.5 NS-19 6.00 1.1 21.7 3.5 -24.9 NS-18 5.60 1.6 13.5 3.6 -24.1 NS-17 5.50 2.2 25.8 3.4 -25.4 NS-16 5.40 1.3 21.5 3.3 -25.3 NS-15 5.30 1.4 17.4 3.3 -25.2 NS-14 5.20 1.7 10.2 3.5 -25.2 NS-13 5.10 2.1 31.5 3.4 -25.1 NS-12 5.00 1.9 28.4 3.4 -24.9 NS-11 4.90 1.3 15.0 3.4 -24.8 NS-10 4.75 1.3 12.5 3.4 -24.8 NS-9b 4.50 2.2 30.9 3.2 -25.7 NS-9a 4.40 0.8 8.7 3.0 -25.4 NS-8 3.90 1.6 18.2 3.0 -26.2 NS-7 3.35 1.7 15.4 3.0 -25.8 NS-6 2.80 1.7 14.8 3.1 -25.3 NS-5 2.15 0.7 13.3 3.1 -25.2 NS-4 1.65 1.0 12.0 3.1 -25.2 NS-3 1.10 1.8 14.2 3.0 -25.7 NS-2 0.55 1.1 13.2 3.0 -25.3 NS-1 0.00 0.9 14.8 2.4 -25.7

A5: Total organic carbon (TOC), total inorganic carbon (TIC) and carbon isotope results, Tarendol section

13 13 Sample Hight TOC TIC δ Ccarb δ Cbulk OM (m) (dry wt %) (dry wt %) (‰ VPDB) (‰ VPDB)

NJ-13 1.94 0.7 34.9 3.3 -23.2 NJ-12 1.84 1.5 21.2 3.2 -23.2 NJ-11 1.66 1.9 25.8 3.3 -23.1 NJ-10 1.56 1.8 31.9 3.2 -24.2 NJ-9 1.46 0.8 35.4 3.2 -23.6 NJ-8 1.28 1.2 29.6 3.2 -23.8 NJ-7 1.18 1.5 25.8 3.4 -23.8 NJ-6 1.08 2.2 27.5 3.4 -23.4 NJ-5 1.00 1.6 32.0 3.9 -23.7 NJ-4 0.80 0.8 33.0 3.3 -23.4 NJ-3 0.60 1.1 28.4 3.4 -23.3 NJ-2 0.45 0.3 36.1 3.4 -23.3 NJ-1 0.25 0.6 29.6 3.4 -23.5

Appendix 154

A6: Carbon isotope results of individual n-alkane measurements (GC-irm-MS), Serre Chaitieu section

Sample Height δ13C δ13C δ13C δ13C δ13C (m) n- C17 stdev n- C18 stdev n- C23 stdev n- C24 stdev n-C28aaa stdev

NS-36 11.75 NS-35 11.35 NS-34 10.95 -27.8 0.3 -26.96 0.37 -26.71 0.07 -26.94 0.03 NS-33 10.55 NS-32 10.15 NS-31 9.75 NS-30 9.35 NS-29 9.15 -28.6 0.0 -28.24 0.34 -26.93 0.19 -27.09 0.18 NS-28 9.05 NS-27 8.60 NS-26 8.20 -28.6 0.1 -28.78 0.05 -27.08 0.04 -27.61 0.03 NS-25 7.80 NS-24 7.40 -27.8 0.7 -27.30 0.09 -26.60 0.01 -27.06 0.10 NS-23 7.00 -28.9 0.1 -28.81 0.07 -26.97 0.08 -27.45 0.08 -29.53 0.07 NS-22 6.55 -29.5 0.2 -30.06 0.05 -28.25 0.04 -28.82 0.12 NS-21 6.45 -29.8 0.1 -30.24 0.20 -28.65 0.08 -28.77 0.04 -30.40 0.33 NS-20 6.25 NS-19 6.00 -29.6 0.0 -30.29 0.23 -28.62 0.09 -28.88 0.05 NS-18 5.60 NS-17 5.50 NS-16 5.40 NS-15 5.30 -29.5 0.0 -29.35 0.07 -28.11 0.03 -28.12 0.06 -29.38 0.11 NS-14 5.20 NS-13 5.10 NS-12 5.00 -30.00 0.04 -28.30 0.03 -28.56 0.12 -29.78 0.02 NS-11 4.90 NS-10 4.75 NS-9b 4.50 -30.46 0.04 -29.59 0.04 -29.72 0.08 -30.48 0.01 NS-9a 4.40 NS-8 3.90 -31.15 0.01 -29.82 0.03 -30.14 0.08 -31.22 0.04 NS-7 3.35 NS-6 2.80 -31.3 0.0 -31.51 0.13 -29.80 0.08 -30.09 0.04 -31.46 0.13 NS-5 2.15 NS-4 1.65 -30.4 0.3 -29.90 0.04 -28.76 0.05 -29.14 0.07 -30.01 0.21 NS-3 1.10 NS-2 0.55 -30.8 0.3 -30.32 0.22 -28.59 0.08 -29.11 0.08 -30.21 0.10

Appendix 155

A7: Palynofacies results, Luz section

Sample Hight Phytoclasts Phytoclasts Phytoclasts Phytoclasts Membranes Cuticle Sporomorphs Dinocysts other Forams AOM Sum (cm) trans. < 25 trans. >25 opaque <25 opaque >25 cysts

A-201 262.1 5 6 2 6 104 4 16 21 15 28 100 307 A-196 254.9 9 7 12 20 50 2 31 18 51 2 140 342 A-193 245.7 8 18 10 26 45 0 105 39 63 13 23 350 A-188 240.1 8 15 8 21 69 4 57 21 33 2 150 388 A-186a 236.4 5 2 2 3 75 0 14 22 5 20 273 421 A-179 231.4 7 5 9 5 47 1 11 11 13 1 195 305 A-176 225.9 19 9 26 18 38 0 73 6 20 1 155 365 A-172 217.4 3 14 6 20 102 0 57 56 41 9 30 338 A-169 213.1 12 2 10 1 35 0 25 70 0 13 140 308 A-162 206.5 2 26 9 38 145 0 29 5 11 42 22 329 A-154a 200.7 7 21 2 13 92 2 34 45 23 21 70 330 A-148 195.3 18 30 17 22 83 4 40 61 27 0 22 324 A-142 191.3 50 5 90 5 98 0 3 0 36 0 7 294 A-137 186.7 12 20 30 36 46 2 76 15 40 2 46 325 A-134 184.0 3 4 3 7 101 6 6 9 45 55 239 A-125 174.1 15 24 2 37 65 4 21 12 103 1 69 353 A-120 170.9 34 3 263 23 6 1 6 336 A-117 167.8 3 0 4 25 48 35 0 2 65 182 A-115 162.9 11 29 7 56 90 5 68 25 27 1 23 342 A-114 159.9 77 28 97 115 19 12 20 11 1 6 386 A-112 157.8 30 57 15 54 51 3 62 63 22 3 25 385 A-110 152.9 0 0 7 13 4 1 33 1 190 249 A-108 148.3 20 36 25 35 72 6 51 96 36 6 383 A-106 146.7 9 31 6 52 21 2 188 7 15 13 344 A-101 140.7 21 20 15 24 97 11 29 12 20 9 58 316 A-97 132.7 0 2 3 6 9 0 86 31 215 352 A-94 129.9 6 19 12 67 52 7 68 92 39 8 16 386 A-81 116.7 12 12 18 20 80 1 69 54 38 3 54 361 A-79 116.3 32 17 13 15 66 3 101 20 19 46 332 A-59a 96.7 6 18 41 146 50 92 8 7 17 30 415 A-46 90.2 76 81 76 40 17 10 21 3 8 1 333 A-41 89.1 58 27 41 35 39 1 60 32 28 321 A-37 83.5 87 43 81 15 8 3 1 20 3 6 267 A-33a 81.5 11 15 23 30 53 151 21 7 23 334 B-15 20.1 12 0 0 2 190 8 212 B-13 17.9 21 25 12 35 178 68 23 10 7 11 390 B-8 12.0 33 15 19 14 135 99 35 20 2 372

Appendix 156

A8: Ericeira section, lower part

Meters Formation Lithology Samples Sedimentology Fossils Meters Formation Lithology Samples Sedimentology Fossils 50 100 K-3

K-2.1 K-2 Qz 95 45 H-63 H-62 Qz H-61 H-60 Qz H-59 H-58 H-57

H-56 Rodisio Formation K-1 H-55 K K 40 90 K H-54

H-103 H-102

H-101 H-100 H-99 H-98 35 H-53 85 H-52 Qz Qz H-51 H-97 H-50 H-96 H-49

H-48 H-95 H-47 H-46 H-94 30 80 H-93 H-92 H-45 H-44 H-91 H-90 H-43 Cresmina Formation H-42 H-41 H-89 H-88 H-40 H-87 H-39 H-86 H-38 25 Qz 75 H-85 H-37 H-36 H-35 H-84 H-34 H-33 H-83 H-82 H-81

H-32 H-80 20 H-31 Qz 70 Formation de Ribeira Ilhas H-30 Qz H-79 Qz H-29 H-78 H-28 Qz Qz Qz H-27 H-77 H-76 Qz H-26 H-25 H-75 H-74 Qz H-24 H-73 15 65 Qz H-72 Qz H-23 H-22 H-21 H-20 H-71 Qz Qz H-70 Qz H-69 H-19 Qz

H-68 Qz H-18 Qz H-67 Qz 10 60 Qz

H-17

H-16 H-15 H-66 H-14 H-13 Qz H-12 5 55 H-11 H-10 H-9 H-65 H-8 H-7 H-6 H-64 H-5 H-4 H-3 H-2 H-1 50 Regatão Formation 0

Appendix 157

A8: Ericeira section, upper part

Meters Formation Lithology Samples Sedimentology Fossils

140 K-38 K-37 K-36 K-35 K-34

K-33

K-32 K-31 135 K-30 K-29

K-28 K-27

K-26 130 K-25

K-24

K-23 Qz

K-22 125 K-21 K-20 Qz K-19

K-18 K-17

120 K-16

Galé Formation K-15

115 K-14 Qz

K-13 K-12 Qz K-11 K-10 Qz

K-9 K-8 110 K-7

K-6

K-5

105 K-4

100 K-3.1

Appendix 158

A9: Cresmina section , lower part

Meters Formation Lithology Samples Sedimentology Fossils 100 Meters Formation Lithology Samples Sedimentology Fossils 50 L2 D 67

D 66 L1 D 65 Qz

D 64 D 63

D 62 D 61 D 60 Qz Qz D 59 Qz Qz D 58 Qz Galé Formation

D 57 Qz Agua Doce Member Qz D 56 Qz D 55

D 54 40 D 53 90

D 52 D 51 D 50 D 93 D 49

D 92 D 48 Qz

Qz D 91 D 47 Qz D 90 Qz D 46 Cresmina Formation

Qz

D 45

D 44 D 89 30 80 D 43 D 42

Qz Rodisio Formation K K D 41 Qz Qz D 40

D 39 D 38 D 37 D 36 D 35 Qz D 88 D 34 D 33 D 32 Qz D 31 Qz D 87 D 30 Qz D 29 D 86 Qz D 28 D 27 D 85 D 26 D 25 Qz D 24 20 70 D 84 D 23 D 83

Cobre Member D 22 D 20/21

D 19 D 82 K

D 18 K

D 17 Praia da Lagoa D 81 K

D 16 D 80 D 79 D 15 D 78 D 14 D 13 D 12 D 77 D 11 D 76

Qz 10 60 D 75 D 10 Qz

D 9 Cresmina Formation

D 8 D 74

Qz

D 7 D 73 D 6 D 72 D 71 D 5 D 70

D 4 Qz Ponta Alta Member D 69 D 3 D D D 68 D 2 K Qz D 1 0 50 D 67

Appendix 159

A9: Cremina section , upper part

Meters Formation Lithology Samples Sedimentology Fossils

L103

Meters Formation Lithology Samples Sedimentology Fossils L102 150 200 L101 L100 L51 L99 L50

L98 Qz L49 L97 L48 K L96 L47 Qz L95

L94 L93 L46 Qz L45 K L44 K Qz L92 L43 K L91 L42 L90

140 K 190 L89

L88 Qz

L87 L86 L41 L40 K

L85

K

L84

L83 L39

L38 130 L37 180 L82 L36 L81 Qz L35 Qz L80 Qz K Qz L34 ion L79

L33 L32 Qz Qz L31 L30 L78 L29 Qz K

L28 Qz

L27 L77 L76

Galé Format Galé L75

Agua Doce Member L74 L26 Galé Formation L25 Agua Doce Member L73 K L24 L72 L23 120 L22 170 L71 L70

L69 K

L68

L21 L20 L19 L18 Qz L17 Qz L67 Qz L16 L15 Qz L66 L14 Qz L65 L13 L64

L63 K L62 110 160 L61 Qz L12 L60

K

L11 L59

L10 L58 L57 L9 L8 L56 L7 L55 L6 Qz L5 K K L54 L4 Qz

L53

Qz 100 L3 150 L52

Appendix 160

A10: Luz section, lower part

Meters Formation Lithology Samples Sedimentology Fossils Meters Formation Lithology Samples Sedimentology Fossils 50 100

A61 Qz

A60 Qz A59

A58 A57 Qz A56 A55 Qz B 23 A54 A53 B 22 A52 B 21 A51 Upper Luz Marls A49

Qz A47 40 90 A46 A45 A43 Qz A41

A40 B 20

A39 K K

A38 Qz

B 19 A37 A35

Qz

A33 Qz Fe A32

30 K K K K 80 Fe

Qz

B 18 Palorbitolina Beds

A31 Qz Fe

A30

B 17 A29

20 A28 Qz 70 Lower Luz Marls B 16 A27 B 15

A26 B 14 B 13

B 12

A25 A24 K

K B 11 B 10 A23 B 9 B 8 10 B 7 Qz 60

B 6 A22 A21

A20

B 5 A19 Qz A17 B 4 A16 A15 K

K A13 K

B 3 K A11

A10 A9 B 2 A8 K G. trochiliscoides Fm A7 K B 1 0 50

Appendix 161

A10: Luz section, middle part

Meters Formation Lithology Samples Sedimentology Fossils Meters Formation Lithology Samples Sedimentology Fossils 200 150 A159

A158 A157 A109 A156 A155 A108 A154 A107 A153

A106 A152 A105 A151

Porto de Mos Fm. A150 A149

A148 A147 140 A104 190 A146 A103 A145 A102 A101 A144

A143 A100 A142 A141

A99 A140 A139

A98 A138 A137 A136 A135 A97 A135a A134 130 180 A133 A96 A132 A131 A95 A94 A131a K K A130

A93 A92 A91

A129

A90

A128

A89 K K A88 A127 A87 A126 Upper Luz Marls

K A125 A124 K 170 120 K A86 K A123 A122 A85 A121

A84 Upper Luz Marls A120 A119 A83 A82 A81 A80 A79 A118 A78 A76 A117 A75 A74 A74a A116

A73 A72

110 160 A71 A115 Qz

K

K

A70 A114 A113 A69 Qz A112 A68 A67 A66 A111 A65 A64 A63

A62 K K 100 150 A110

Appendix 162

A10: Luz section, upper part

Meters Formation Lithology Samples Sedimentology Fossils

250 A196

A195 Meters Formation Lithology Samples Sedimentology Fossils

A 223

A 222 A194 A 221 A193

A192 240 290 A 220

A191 A190 A 219

A189 A188

A 218

A187 A 217

A 216 A186 A185 A184 A 215 230 A183 280 A 214 A182 A181 A180 A 213

A179

A178 A 212

A 211 A177 A 210 Porto de Mos Fm. A176

A 209 220 270

K A 208 A175 ? Porto de Mos Fm. A 207 A174

A173 A 206

A 205

A172 A 204

A171 A 203 210 260 A170 A 202 A169

A168 A 201

A167 A166

A165 A 200

A164 A199 A163 A162 A161 A198

A197 200 A160 250

Appendix 163

A11: Burgau section, lower part

Meters Formation Lithology Samples Sedimentology Fossils Meters Formation Lithology Samples Sedimentology Fossils 50 100

C 4

Qz

C 3

K G19 G18 G17 C 2 Lower Luz Marls C 1 G16 40 90 Upper Luz Marls

G15

G14 30 80 G13

G12

G11

G10 Qz Qz G7/8/9 Qz G6

G5 G4 K Qz G3/50

G2/49 Qz Qz Palorbitolina Beds G1/48 Qz 20 70

G47

G46 K K 10 Qz 60 Lower Luz Marls G45

G44

G43

G42

K K

G41

0 G. trochiliscoides Fm. 50

Appendix 164

A11: Burgau section, upper part

Meters Formation Lithology Samples Sedimentology Fossils Sedimentary structures 150 G40

Hardground Fining upward G39 G38 Birdseyes Bioturbation G37 Teepe structure Burrow

G36 Small wavy stratification Fossil wood

Current ripples Lithoclasts

G35 Wave ripples Qz Quarz grains

Cross bedding K Calcareous nodules

Nodular bedding D Dolomitic nodules 140 G34 Tabular cross bedding Load cast

Trough cross bedding Channel

G33 Heringbone cross bedding

Flaser bedding

G32 Graded bedding

G31

Fossil content 130 G30 Bioclasts (undifferentiated)

G29 Bivalves Echinoderms

Brachiopods Sponges

Gastropods Corals G28

G27 Ostracods Cyclamminidae

Green algae Agglutinating benthic foram.

Upper Luz Marls Charpophytes Miliolinids

120 Fish debris Undifferentiated benthic foram. G26 Oysters Serpulids

Rudists Bryozoans G25 Stromatoporids Orbitolinids

G24

Lithology

G23 110 G22 Limestone Siltstone

Calcareous marl Sandstone

Claystone Conglomerate

G21 G20

100

Acknowledgements 165

Acknowledgements

First of all I thank Helmi Weissert, Peter Hochuli and Nils Andersen, who initiated the Portugal project, for their enthusiastic support and ongoing motivation during my PhD.

Dear Helmi, thank you for giving me the great opportunity to join your group here at the ETH Zurich as your PhD student. I really enjoyed the collaborative fieldwork along the wonderful Portuguese coasts and the fruitful and inspiring discussions in front of the outcrop. At the ETH, your office door was always kept open for all of your students, entering to discuss scientific results, philosophical hypotheses or personal hardships. This is something I will truly miss and always remember. Thank you for encouraging and motivating me so much during the last four years of my PhD time.

Dear Peter, I am very grateful for the never-ending patience, you exercised in teaching me palynolgy. You never got tired of all my frequent questions and visits, over there in the Palaeontological Museum. Thank you for spending so much time, patience and effort on our project with its demanding stratigraphic issues and challenging paleobotanical problems. Working with you on the microscope was very motivating and instructive and I profited incredibly from your extensive palaeobotanical expertise.

Dear Nils, thank you for your help and technical assistance with the stable isotope measurements and for introducing me to the organic-geochemical methods and the GC-MS analytics. I very much appreciate your corrections and comments on earlier versions of the manuscripts as well as the stimulating and inspiring discussions.

Furthermore I want to thank Judy McKenzie for her interest in this work and for the great opportunity to join her on an exciting field trip to Brazil. I am thankful to Stefano Bernasconi for the ongoing support in the stable isotope lab and the critical discussions of the final results. Thanks to Jens Herrle for the introduction to the sediments of the Vocontian Basin. I really appreciate the ongoing discussions on mid-Cretaceous palaeocanography as well as the helpful and constructive comments on earlier drafts of the manuscripts. Special thanks to my co-worker Stefan Burla for the absolutely amazing times during field work in Portugal - this was for sure the most hilarious part of the project.

Acknowledgements 166

My thanks go to Jorge Dinis, Ramon Gonzales and Martina Bachmann for their support during field work in Portugal and to Peter Skelton for the biostratigraphic support and the helpful suggestions and discussions during our last field campaign. Furthermore, I am very grateful to Stephen Hesselbo, who did not hesitate to join the scientific committee as a co- examiner. Thanks to Luc Zwanc from the EAWAG and Christian Ostertag-Henning from the University of Münster for support with the compound-specific measurements and the identification of particular organic compounds in my samples. Furthermore I want to thank Rui Pena dos Reis, University of Coimbra, for providing the beautiful cover picture.

Moreover, I want to thank all my friends and colleagues in the Geological Institute, who provided a wonderful working atmosphere during my stay here in Zürich. Thanks to all of you for giving me such a great and exciting time, filled with unique humor, sincere friendship and respect. Doing a PhD in the Earth System Sciences group at the ETH was really great fun.

Lastly, I’d like thank my family for their persistent encouragement and support during my studies and my girlfriend Uta for being at my side and sharing the ups and downs of a PhD student’s life. Curriculum Vitae

Ulrich Heimhofer

Date of birth: 19. October 1971 Place of birth: Sonthofen i. Allgäu, Germany Nationality: German

Education

1978-1982: Grundschule Burgberg, Germany 1982-1991 Gymnasium Sonthofen, Germany 1991-1993: Civilian national service at the Red Cross, Immenstadt, Germany

1993-1999: Undergraduate student at the Faculty of Natural Sciences at the Friedrich-Alexander University Erlangen-Nürnberg, Germany 1995-1996: Visiting student at the Department of Earth Sciences, ETH Zurich, Switzerland 1996-1999: Diploma student (Geology/Palaeontology) at the Department of Geology, Friedrich-Alexander University Erlangen-Nürnberg, Germany

2000-2004: Doctoral student and research assistant at the Geological Institute, ETH Zurich, Switzerland

Dissertation: Response of terrestrial palaeoenvironments to past changes in climate and carbon-cycling: Insights from palynology and stable isotope geochemistry

Supervisors: Prof. Dr. Helmut Weissert PD Dr. Peter A. Hochuli Dr. Nils Anderson